Stromatolites (791 p, 1976)

DEVELOPMENTS IN SEDIMENTOLOGY 20

STROM ATOLITES

FURTHER TITLES IN THIS SERIES

1. L. M. J. U. V A N STRAATEN, Editor DELTAIC AND SHALLOW MARINE DEPOSITS

2. G.C. AMSTUTZ, Editor SEDIMENTOLOGY AND ORE GENESIS

3. A.H. BOUMA and A . BROUWER, Editors TURBIDITES

4. F.G. TICKELL THE TECHNIQUES OF SEDIMENTARY MINERALOGY

5. J. C. INGLE Jr. THE MOVEMENT OF BEACH SAND

.6. L. V A N DER PLAS THE IDENTIFICATION OF DETRITAL FELDSPARS

I. S. DZULYNSKI and E.K. WALTON SEDIMENTARY FEATURES OF FLYSCH AND GREYWACKES

8. G. LARSEN and G. V. CHILINGAR, Editors DIAGENESIS I N SEDIMENTS

9. G. V. CHILINGAR, H.J. BISSELL and R. W. FAIRBRIDGE, Editors CARBONATE ROCKS

10. P. McL. D. DUFF, A. HALLAM and E.K. WALTON CYCLIC SEDIMENTATION

11. C.C. REEVES Jr. INTRODUCTION TO PALEOLIMNOLOGY

12. R.G.C. BATHURST CARBONATE SEDIMENTS AND THEIR DIAGENESIS

13 A.A. MANTEN SILURIAN REEFS OF GOTLAND

14. K.W. GLENNIE DESERT SEDIMENTARY ENVIRONMENTS

15. C.E. W E A V E R and L.D. POLLARD THE CHEMISTRY OF CLAY MINERALS

16. H.H. RIEKE III and G.V. CHILINGARIAN COMPACTION OF ARGILLACEOUS SEDIMENTS

11. M.D. PICARD and L.R. HIGH Jr. SEDIMENTARY STRUCTURES OF EPHEMERAL STREAMS

18. G. V. CHILINGARIAN and K.H. WOLF COMPACTION OF COARSE-GRAINED SEDIMENTS

19. W. SCHWARZACHER SEDIMENTATION MODELS AND QUANTITATIVE STRATIGRAPHY

DEVELOPMENTS IN SEDIMENTOLOGY 20

STROMATOLITES

EDITED BY

M. R. WALTER

Bureau of Mineral Resources, Geology and Geophysics, Canberra, A.C.T. (Australia)

ELSEVIER SCIENTIFIC PUBLISHING COMPANY Amsterdam -.I Oxford -- New York 1976

ELSEVIER SCIENTIFIC PUBLISHING COMPANY 335 Jan van Galenstraat P.O. Box 211, Amsterdam, The Netherlands

Distributors for the United States and Canada:

ELSEVIER/NORTH-HOLLAND INC. 52 Vanderbilt Avenue New York, N.Y. 10017

Library of Congress Cataloging in Publication Data Main en t ry under t i t l e :

Strornatolites . (Developments i n sedirnentology ; 20) Bib l iogrqhy: p. Includes index. 1. Stromatolites. I. Walter, M . R. 11. Series.

QE955.S82 561l.93 76-17546 ISBN 0-444-41376-6

Copyright 0 1976 by Elsevier Scientific Publishing Company, Amsterdam

All rights reserved. No part of this publication may be reproduced, stored in a retrieval system, or transmitted in any form or by any means, electronic, mechanical, photocopying, recording, or otherwise, without the prior written permission of the publisher, Elsevier Scientific Publishing Company, Jan van Galenstraat 335, Amsterdam

Printed in The Netherlands

LIST OF CONTRIBUTORS

S.M. AWRAMIK Department of Geological Sciences University of California Santa Barbara, Calif. 93106 (U.S.A.)

E.S. BARGHOORN Biological Laboratories Harvard University Cambridge, Mass. 02138 (U.S.A.)

J. BAULD Department of Microbiology University of Melbourne Parkville, Vic. 3052 (Australia)

J. BERTRAND-SARFATI Facultd des Sciences C.N.R.S., Centre Gkophysique et Gkologique de Montpellier 34060 Montpellier (France)

T.D. BROCK Department of Bacteriology University of Wisconsin Madison, Wis. 53706 (U.S.A.)

A. BUTTON Economic Geology Research Unit University of the Witwatersrand Johannesburg (South Africa)

A.E. COCKBAIN Geological Survey of Western Australia Perth, W.A. 6000 (Australia)

W.E. DEAN Syracuse University Syracuse, N.Y. 13120 (U.S.A.)

J.A. DONALDSON Department of Geology Carleton University Ottawa 1, Ont. (Canada)

E.C. DRUCE Bureau of Mineral Resources, Geology and Geophysics Canberra City, A.C.T. 2601 (Australia)

J.R. EGGLESTON Geologic and Economic Survey Morgantown, W. Va. 26505 (U.S.A.)

K.A. ERIKSSON Department of Geology University of the Witwatersrand Johannesburg (South Africa)

C.D. GEBELEIN Department of Geological Sciences University of California Santa Barbara, Calif. 93106 (U.S.A.)

S. GOLUBIC Department of Biology Boston University Boston, Mass. 02215 (U.S.A.)

R.B. HALLEY U S . Geological Survey University of Miami Fisher Island Station, School of Marine Science Miami, Fla. 33149 (U.S.A.)

VI

L.A.HARDIE Department of Earth and Planetary Sciences The Johns Hopkins University Baltimore, Md. 21218 (U.S.A.)

P.G. HASLETT Department of Geology and Mineralogy University of Adelaide Adelaide, S.A. 5000 (Australia)

LIST O F CONTRIBUTORS

P.F. HOFFMAN Geological Survey of Canada Ottawa, Ont. K1A OE8 (Canada)

H.J. HOFMANN Department of Geology. University of Montreal Montreal, Que. (Canada)

R.J. HORODYSKI Department of Geology University of California Los Angeles, Calif. 90024 (U.S.A.)

D.J.J. KINSMAN DepartHent of Geological and Geophysical Sciences Princeton University Princeton, N.J. 08540 (U.S.A.)

I.N. KRYLOV Geoiogical Institute Academy of Sciences of the U.S.S.R. Moscow Zh. 17 (U.S.S.R.)

L. MARGULIS Department of Biology Boston University Boston, Mass. 02215 (U.S.A.)

D.M. McKIRDY Bureau of Mineral Resources, Geology and Geophysics Canberra City, A.C.T. 2601 (Australia)

F. MENDELSOHN Maritime House Johannesburg 2001 (South Africa)

C.L.V. MONTY Laboratoire de Paldontologie animale Universitd de Lidge 4000 Lidge (Belgium)

G. PANNELLA Department of Geology University of Puerto Rico Mayaguez 00708 (Puerto Rico)

R.K. PARK Phillips Petroleum Company Denver, Colo. 80202 (U.S.A.)

P.E. PLAYFORD Geological Survey of Western Australia Perth, W.A. 6000 (Australia)

W.V. PREISS Department of Mines Adelaide, S.A. 5001 (Australia)

M.E. RAABEN Geological Institute Academy of Sciences of the U.S.S.R. Moscow Zh.17 (U.S.S.R.)

J.F. READ Department of Geological Sciences Virginia Polytechnic Institute and State University Blacksburg, Va. (U.S.A.)

M.A. SEMIKHATOV Geological Institute Academy of Sciences of the U.S.S.R. Moscow Zh.17 (U.S.S.R.)

S.N. SEREBRYAKOV Geological Institute Academy of Sciences of the U.S.S.R. Moscow Zh.17 (U.S.S.R.)

R.C. SURDAM Department of Geology University of Wyoming Laramie, Wyo. 82071 (U.S.A.)

J. THRAILKILL Department of Geology University of Kentucky Lexington, Ky. 40506 (U.S.A.)

R. TROMPETTE htudes gdologique Ouest-africaines C.N.R.S. Centre Universitaire de St-Jdrbme 13013 Marseille (France)

LIST OF CONTRIBUTORS

P.A. TRUDINGER Baas-Becking Geobiological Laboratory Canberra, A.C.T. 2601 (Australia)

J.F. TRUSWELL Bureau of Mineral Rssources, Geology and Geophysics Canberra City, A.C.T. 2601 (Australia)

C.C. VON DER BORCH Flinders University Bedford Park, S.A. 5042 (Australia)

VII

M.R. WALTER Bureau of Mineral Resources, Geology and Geophysics Canberra City, A.C.T. 2601 (Australia)

J.L. WRAY Marathon Oil Company Littleton, Colo. 80122 (U.S.A.)

This Page Intentionally Left Blank

CONTENTS

List of Contributors. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . V

1. INTRODUCTION (M.R. Walter). . . . . . . . . . , . . . . . . . . . . . . . . . . . . . . . . 2. METHODOLOGY AND SYSTEMATICS

1

Chapter 2.1. Basic field and laboratory methods for the study of stromatolites (W.V.Preiss). . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

Chapter 2.2. Graphic representation of fossil stromatoids; new method with improved precision (H.J. Hofmann) . . . . . . . . . . . . . . . . . . . . . , . . . . . .

Chapter 2.3. Biological techniques for the study of microbial mats and living stromatolites (T.D. Brock) . . . . . . . . . . . . . . . . . , . . . . . . . . . . . . . . . .

Chapter 2.4. Approaches to the classification of stromatolites (I.N. Krylov) . . . . Chapter 2.5. Stromatoid morphometrics (H.J. Hofmann) . . . . . , . . . . . . . . . .

5

15

21

31

45

3. ABIOGENIC STROMATOLITE-LIKE STRUCTURES

Chapter 3.1. Calcretes and their distinction from stromatolites (J.F. Read). . . . . Chapter 3.2. Speleothems (J. Thrailkill). . . . . . . . . . . . . . . . . . . . . . . . . . . . Chapter 3.3. Geyserites of Yellowstone National Park: an example of

55

73

87 abiogenic “stromatolites” (M.R. Walter) . . . . . . . . . . . . . . . . . . . . . . . . . 4. BIOLOGY OF STROMATOLITES

Chapter 4.1. Organisms that build stromatolites (S. Golubic). . . . . . . . . . . . . . Chapter 4.2. Taxonomy of extant stromatolite-building cyanophytes

(S.Golubic). . . , . . . . . . . . . . . . . . . . . . . :. . . . . . . . . . . . . . . . . . . , Chapter 4.3. Environmental microbiology of living stromatolites (T.D. Brock) . . Chapter 4.4. Evolutionary processes in the formation of stromatolites

(S.M. Awramik, L. Margulis and E.S. Barghoofn). . . . . . . . . . . . . . . . . . . . Chapter 4.5. Biochemical markers in stromatolites (D.M. McKirdy). . . . . . . . . .

113

127

141

149

163

X CONTENTS

5. FABRIC AND MICROSTRUCTURE

Chapter 5.1. The origin and development of cryptalgal fabrics (C.L.V. Monty) . Chapter 5.2. An attempt to classify Late Precambrian stromatolite

microstructures (J. Bertrand-Sarfati) . . . . . . . . . . . . . . . . . . . . . . . . . . .

6. MORPHOGENESIS

Chapter 6.1. Stromatolite morphogenesis in Shark Bay, Western Australia

Chapter 6.2. Microbiology and morphogenesis of columnar stromatolites

(P. Hoffman) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

(Conophyton, Vacerrillo) from hot springs in Yellowstone National Park (M.R. Walter, J. Bauld and T.D. Brock) . . . . . . . . . . . . . . . . . . . . . . . . . .

Chapter 6.3. Gunflint stromatolites : microfossil distribution in relation to

Chapter 6.4. Biotic and abiotic factors controlling the morphology of Riphean

stromatolite morphology (S.M. Awramik) . . . . . . . . . . . . . . . . . . . . . . . .

stromatolites (S.N. Serebryakov) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

7. STROMATOLITE BIOSTRATIGRAPHY

Chapter 7.1. Experience in stromatolite studies in the U.S.S.R.

Chapter 7.2. Intercontinental correlations (W.V. Preiss) . . . . . . . . . . . . . . . . . Chapter 7.3. Aphebian stromatolites in Canada: implications for stromatolite

zonation (J.A. Donaldson) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

8. RECENT MODELS FOR INTERPRETING STROMATOLITE ENVIRONMENTS

(M.A. Semikhatov) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

Chapter 8.1. Open marine subtidal and intertidal stromatolites (Florida, the

Chapter 8.2. Modern algal stromatolites at Hamelin Pool, a hypersaline barred basin in Shark Bay, Western Australia (P.E. Playford and A.E. Cockbain). . . .

Chapter 8.3. Stratigraphy of stromatolite occurrences in carbonate lakes of the Coorong Lagoon area, South Australia (C.C. von der Borch) . . . . . . . . . . . .

Chapter 8.4. Algal belt and coastal sabkha evolution, Trucial Coast, Persian Gulf (D.J.J. Kinsman and R.K. Park). . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

Chapter 8.5. Textural variation within Great Salt Lake algal mounds

Chapter 8.6. The geological significance of the freshwater blue-green algal calcareous marsh (C.L.V. Monty and L.A. Hardie) . . . . . . . . . . . . . . . . . . .

Chapter 8.7. Freshwater stromatolitic bioherms in Green Lake, New York (J.R. Eggleston and W.E. Dean) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

Chapter 8.8. Hot-spring sediments in Yellowstone National Park (M.R. Walter). .

Bahamas and Bermuda) (C.D. Gebelein). . . . . . . . . . . . . . . . . . . . . . . . . . .

(R.B. Halley) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

9. APPLICATION OF RECENT MODELS TO THE GEOLOGICAL RECORD

Chapter 9.1. The effects of the physical, chemical and biological evolution of the earth (C.D. Gebelein) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

193

251

261

27 3

311

321

337

359

371

381

389

41 3

421

435

447

47 9

489

499

CONTENTS

10. STROMATOLITES IN BASIN ANALYSIS

Chapter 10.1. Use of stromatolites for intrabasinal correlation : example from the Late Proterozoic of the northwestern margin of the Taoudenni Basin (J. Bertrand-Sarfati and R. Trompette) . . . . . . . . . . . . . . . . . . . . . . . . . .

Precambrian Dismal Lakes and Rae Groups, Canada (J.A. Donaldson) . . . . . . Chapter 10.2. Paleoecology of Conophyton and associated stromatolites in the

Chapter 10.3. Lacustrine stromatolites, Eocene Green River Formation, Wyoming

Chapter 10.4. Devonian stromatolites from the Canning Basin, Western Australia (P.E. Playford, A.E. Cockbain, E.C. Druce and J.L. Wray). . . . . . . . . . . . . .

Chapter 10.5. Lower Cambrian stromatolites from open and sheltered intertidal environments, Wirrealpa, South Australia (P.G. Haslett) . . . . . . . . . . . . . . .

Chapter 10.6. Stromatolites from the Middle Proterozoic Altyn Limestone, Belt Supergroup, Glacier National Park, Montana (R.J. Horodyski). . . . . . . .

Chapter 10.7. Environmental diversity of Middle Precambrian stromatolites (P. Hoffman) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

Chapter 10.8. Distribution of stromatolites in Riphean deposits of the Uchur-Maya region of Siberia (S.N. Serebryakov) . . . . . . . . . . . . . . . . . . .

Chapter 10.9. Palaeoenvironmental and geochemical models from an Early Proterozoic carbonate succession in South Africa (K.A. Eriksson, J.F. Truswell andA.Button) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

(R.C. Surdam and J.L. Wray). . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

11. MINERALIZATION ASSOCIATED WITH STROMATOLITES

Chapter 11.1. Mineral deposits associated with stromatolites (F. Mendelsohn). . . Chapter 11.2. Biological processes and mineral deposition (P.A. Trudinger and

F. Mendelsohn). . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

12. CONTRIBUTIONS TO THE HISTORY OF THE EARTH-MOON SYSTEM

Chapter 12.1. Geophysical inferences from stromatolite lamination (G. Pannella) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

XI

APPENDICES

Appendix I. Glossary of selected terms (M.R. Walter). . . . . . . . . . . . . . . . . . . Appendix 11. Table of time-ranges of the principal groups of Precambrian

stromatolites (I.N. Krylov and M.A. Semikhatov) . . . . . . . . . . . . . . . . . . .

517

523

535

543

565

585

599

613

635

645

663

67 3

687

693

Appendix 111. List of available translations of major works on stromatolites

Appendix IV. Selective subject index to the Bibliography (S.M. Awramik) . . . . . (M.R. Walter). . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 695

697

BIBLIOGRAPHY . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 705

INDEX . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 773

INTRODUCTION

M.R. Walter”

It is 68 years since E. Kalkowsky coined and defined the word “stromatolith”, yet there is increasing controversy and confusion as to its use. Usage is not uniform in this book, nor have I attempted to make it so. The original definition is now of historical interest only. The only unifying feature of stromatolites is their genesis. The definition I find most useful is one modified from that presented in 1974 by S.M. Awramik and L. Margulis in the second Stromatolite Newsletter (unpublished) :

Stromatolites are organosedimentary structures produced by sediment trapping, binding and/or precipitation as a result of the growth and metabolic activity of micro-organisms, principally cyanophytes.

Non-genetic definitions are so broad as to include a wide range of nonbiogenic structures, so that the term would cease to be useful. The definition used here embraces the whole range of forms, including stratiform and unlaminated structures that some workers would exclude. It has generated the structure of this book.

For some time the belief that stromatolites are mainly a marine intertidal phenomenon dominated the literature. The distribution of stromatolites is not restricted in this way now, nor was it during the past, particularly during the Proterozoic. Stromatolites can form in almost any free-standing body of water, providing the following conditions are met:

(1) the environment is suitable for the growth of appropriate microorganisms;

(2) the rate of growth of the constructing micro-organisms exceeds their rate of consumption by other organisms;

(3) the rate of sedimentation is sufficient to produce a preservable structure but is not so great as to prevent colonization by micro-organisms;

* Contributions by M.R. Walter in this book, except sections 3.3, 6.2 and 8.8, are published with the permission of the Acting Director, Bureau of Mineral Resources, Geology and Geophysics, Canberra City, Australia.

2 M.R. WALTER

(4) the stromatolites accrete faster than they can be destroyed by boring and burrowing organisms and erosive and other mechanical and chemical forces.

These propositions may seem to be truisms yet they serve to explain the present and past distribution of stromatolites, as is discussed by many authors within this book. I t is now generally accepted that stromatolites may form in a wide range of environments, but another simplistic notion is still prevalent. That is that all stromatolites are built by cyanophytes. This is not so now, nor was it necessarily the case in the past; many ancient stromatolites may have been built by bacteria or eucaryotic algae.

The study of stromatolites is essentially one of morphogenesis. There are several good studies of the morphogenesis of Holocene stromatolites and these provide the basis for the uniformitarian interpretation of ancient stro- .matolites. Such an approach is limited by the small range of stromatolite forms known from the Holocene, in contrast with the great variety described from the rock record, particularly the Proterozoic. Morphogenetic analysis has also been applied to ancient stromatolites with no Holocene homeo- morphs. Both approaches will be found in this book. Neither approach is simple and the effort involved is worthwhile only because of the wealth of subtle palaeoenvironmental and palaeobiological information encoded in the shapes and microstructures of stromatolites. We are only just beginning to crack the code. It is this basic role of morphology in the study of these structures, which are not true fossils but are more than sedimentary structures, that has produced the need for a taxonomy and has thus resulted in much debate. A fruitless part of the debate has concerned nomenclature; the fruit is to be found in discussions of classification, for those who have the fortitude to wade through them. Suffice it to say that it is retrograde to force complex structures into a simple classification.

The study of stromatolites has burgeoned in recent years, and at the same time it has become more interdisciplinary. Microbiologists are making increasingly significant contributions to a field previously dominated by palae- ontologists and sedimentologists. Now there are even geochemists, geo- physicists and computer technologists hovering on the fringes of the field. The resultant abundance of new information and ideas made available over the past 5-10 years has not been assimilated by the geological profession at large and frequently there are false impressions, even among specialists, as to the attitudes of other specialists. These points apply particularly where research is published in languages other than that of the user. For all these reasons, I believe the publication of this book is timely. All contributions were invited by me in an attempt to provide a comprehensive coverage of studies of stromatolites. No attempt was made, however, to make the book geographically representative. The 41 authors have been remarkably co- operative and I have received all but 3 of the contributions originally invited. I have attempted t o balance reviews of tried and tested work with reports of

INTRODUCTION 3

new methods and research, aiming both to reflect accurately the present state of knowledge and to produce a book useful to specialists and non-specialists alike. All authors were instructed to guide readers to the pertinent literature, and the bibliography is, I believe, complete within its defined limits. This, I hope, will compensate for gaps in coverage that readers may detect within the text.

Gaps not covered by the text or the bibliography indicate those areas not yet studied. Some that can be perceived because research to fill them has just begun include the microbiology of Holocene stromatolites, the microfossil content of ancient stromatolites, the effects of diagenesis on stromatolite microstructure, the interpretation of deep-sea manganese nodules as bacterial stromatolites, and the geochemistry of stromatolitic carbonates. Much research remains to be done on the most controversial of all subjects in this field, stromatolite biostratigraphy. This subject is thoroughly discussed in this book, from many points of view, but the reader will not learn here whether or not the method is valid. The available microbiological and micro- palaeontological data suggest that stromatolite biostratigraphy should be possible, but more and more evidence is being uncovered which is inconsist- ent with the simple biostratigraphic scheme used 5-10 years ago. I t may be that the main problem lies in the taxonomies presently in use. We should remember that the history of stromatolite biostratigraphy essentially spans only the last 17 years; compared to other fields of biostratigraphy, this one is in its infancy so we should not be too quick to criticize. Even if stromatolite biostratigraphy eventually proves to be impossible, we will have accumulated abundant data for making much more precise palaeoenvironmental and palaeobiological interpretations than are presently possible.

Editing this publication has occupied a significant part of my time for two years and would not have been possible without the support and assistance of many people. I want to take this opportunity to thank four individuals whose contributions have been indirect yet vital: M.F. Glaessner as my teacher and mentor, Preston Cloud as a source of inspiration and confidence, Brian J. Skinner who by inviting me to Yale University gave me the opportunity to broaden my experience and eventually to meet most of my col- leagues who were to contribute to this book, and Marilyn, my wife, who has shared this burden with me. The Bureau of Mineral Resources, Geology and Geophysics generously allowed me to take on this task while in their employ, and for that I am truly thankful. Almost every paper in this book was scru- tinised by at least one referee and while the authors and I must bear ultimate responsibility, the referees’ contribution. was indispensable. I shall not mention them all by name but several because of their skills and fortunate geographic location during the editing process bore more than their fair share of the work: they are W.V. Preiss, A.L. Donaldson and J.F. Truswell.

2. METHODOLOGY AND SYSTEMATICS

Chapter 2.1

BASIC FIELD AND LABORATORY METHODS FOR THE STUDY OF STROM ATOLITES

W.V. Preiss

INTRODUCTION

The methods employed in the study of stromatolites depend on the aims of one’s investigations and on the size limitations of the stromatolites. Thus, the field geologist will be mainly concerned with observations on the geometry and relationships of stromatolites and the information this provides in determining stratigraphic facings, characterizing mappable rock-units, and interpreting environments of deposition. Although he will not normally require the detailed laboratory investigation necessary for the identification and classification of stromatolites, it should be noted that taxonomy also provides valuable environmental data. Conophy ton, for example, is a stromatolite which largely formed only in subtidal environments (e.g. Donaldson, Ch. 10.2).

FIELD TECHNIQUES

Some of the methods suggested below may not apply in less than ideal exposures. The following check-list is intended as a guide to observations to be made wherever possible. The features described are illustrated in Fig. 1.

Large-scale structures-the nature of stromatolitic beds

(1) What is the thickness of the stromatolitic bed? (2) Is it composed of a single cycle of stromatolitic morphologies and as-

(3) What is the lateral extent of the bed and its constituent units? (4) What is the shape of the bed and its constituent units, and what are the

(5) What are the lithologies of the stromatolites and surrounding sedi-

sociated sediments, or of several superimposed cyclical units?

relationships of these to the surrounding rock types?

ments?

6 W.V. PREISS

MODE OF OCCURRENCE

DOMED

OA L

SLIGHTLY PARALLEL

C G DIVERGENT MARKEDLY DIVERGENT COLUMNS

COLUMN SHAPE AND MARGIN STRUCTURE

PEAKS BUMPS

PROJECTION WALL

LAMINA SHAPE

STEEPLY CONVEX RHOMBIC WRINKLED UNCONFORM lTY CRESTAL ZONES IN

CONOPHYTONS NON-COLUMNAR STROMATOLITES

LATERALLY -LINKED CUMULATE COLUMNAR-LAYERED

Fig. 1. Diagrammatic illustrations of terms used in the description of diagnostic characters of stromatolites. Reproduced from Trans. R. Soc. S. Aust., 96(2); 69, with the permission of the Royal Society of South Australia.

METHODOLOGY AND SYSTEMATICS 7

(6) What is the shape and relief of successive growth interfaces of the stromatolites? (The “synoptic profile” of Hofmann, 1969a). The synoptic surfaces of several successive stages of growth of certain stromatolitic bioherms from Central Australia were constructed by Walter (1972a). Such information is important for palaeoecology since it provides a minimum depth of water by which the stromatolites were covered, at least intermittently.

The mode of occurrence can then be categorized according to the definitions listed in the glossary (see also Preiss, 1972b; Walter, 1972a): bioherm (tabular bioherm, domed bioherm, subspherical bioherm, tonguing bioherm); biostrome (tabular biostrome, domed biostrome).

In term edia te-scale strom a toli t ic structures

(1) What is the nature of the constituents of the stromatolitic beds? Any of the following may be present, in any combination: Stratiform stromatolites (flat-laminated, laterally linked, undulatory and pseudocolumnar), columns and cumuli (cumulate stromatolites).

(2) What are the vertical and lateral relationships of these constituents to one another?

(3) If columns are present, are they branched? If so, note the style and frequency of branching: parallel branching (a-parallel, 0-parallel, y-parallel), slightly divergent branching, markedly divergent branching.

(4) What is the shape of columns in longitudinal sections?, e.g. regular, sub-cylindrical, tuberous, gnarled, straight, or curved.

( 5 ) What is their shape in transverse section?, e.g. round, oval, rounded polygonal, lobate, or elongated.

(6) What are the relative orientations of columns to one another? (7) What is the nature of the column margins?, e.g. smooth, bumpy, ragged,

or with large overhangs. (8) If the margins are smooth or bumpy, are they formed by laminae

bending over and enveloping the lateral surface to form a wall? (9) What is the shape of the laminae in the columns?, e.g. conical (this can

be seen only in axial section), gently to steeply convex, rectangular, rhombic, smooth, wavy or wrinkled.

(10) Are there any lateral or vertical changes in any of these features when traced throughout the bioherm or biostrome, especially for instance, at bioherm margins?

(11) What is the nature of the sediment in the interspaces between columns? This will provide valuable information on the environment of deposition.

(12) The upward convexity of stromatolite laminae is often used as an indicator of stratigraphic facing. Nevertheless, upward concave laminae do sometimes occur (Fig. 2), and care must also be taken to distinguish stromatolitic laminae from laminations in the interspaces, which are often concave

8 W.V. PREISS

Fig. 2. Pseudocolumnar stromatolites, illustrating both upward convex and upward concave lamination, Guck Creek Dolomite, west of Mount Stuart Homestead, Hamersley Basin, W.A. (Photo: M.R. Walter.)

upwards (Hofmann, 1973). Wherever possible, determination of facings should be based on columnar-branching stromatolites where the direction of growth is obvious. The presence of erosional micro-unconformities in the lamination is frequently helpful.

Photography and sampling

Where possible, photographs should illustrate the observations outlined above. Photo-montage of juxtaposed photographs may be necessary, for example, to show elongated structures such as biostromes. Stromatolitic structures are generally readily apparent on clean weathered surfaces free of lichen cover. If, however, there is insufficient contrast for photography, it may be necessary to outline columns with a felt-tipped marking pen (e.g. Walter, 1972a, pls. 19(1), 20(5), 21(3)).

Sampling should be commenced only after photography, and then a sketch


of the photographed outcrop showing location of samples will be useful for future reference. The number and size of samples required will vary with the nature of the stromatolites, and the following is only an approximate guide:

(1) Samples should be selected such that they fit readily into the vice of a slabbing saw, e.g. roughly 25 cm cubes in the case of a 61 cm diameter saw, but longer specimens (to about 50 cm) can be accommodated sideways.

(2) They should show on at least one surface a clear cross-section of several columns.

(3) Different samples should be selected to be representative of the whole outcrop, including samples from the centre and margins of bioherms, and from the base to the top of the bed to illustrate vertical changes.

(4) Single samples should be selected which illustrate vertical changes from one morphology to another (e.g. flat-laminated to columnar) or changes in style of branching.

(5) In the case of stromatolite columns too large to be handled physically, careful field observations, preferably in different orientations, will be needed to show branching style, column shape, lamina shape and margin structure. Specimens can then be collected to show portions of columns, e.g. marginal parts for reconstruction to show margin structure.

Thus, the number of samples required depends on the thickness of the stromatolite bed and the diversity of morphologies represented in it.

LABORATORY TECHNIQUES

For the identification of previously described stromatolites and for the definition of new taxa, the following procedure is required.

Reconstruction

Three-dimensional reconstruction is necessary to show features which do not appear or cannot be differentiated in single sections. The method has been described by Krylov (1963) and Walter (1972a) and will be summarized here. The rock is cut on a saw as described above, parallel to the length of the columns, into slabs up to 6 mm wide, depending on the width of columns. Generally 10-15 cuts are required, each cut providing two sections separated by the blade width (about 2 mm for a 61 cm blade). Two preliminary cuts at right angles to the slabs provide reference surfaces for reconstruction. The slabs are numbered, and their surfaces wetted to allow the columns to be outlined in pencil. Then a column or group of columns is selected for reconstruction and followed from one section to the next. The outlines are then traced on to a block diagram framework on tracing paper, usually drawn with an angle of 45" between the front face and the line representing the top of the side reference face. Successive longitudinal sections are placed

10 W.V. PREISS

SLABS

I I W

/--

A Actual W!dth

Actua l Cut Widlt

Slab 5mm

I 2mm

74-152 De1.C.R.S.

Fig. 3. Method of reconstructing columnar stromatolites from serial slabs (left) and tracing on to a block diagram framework on tracing paper (right).

against this framework, turning over the tracing paper from one section to the next; each slab is displaced along the edge of the reference face by the distance from the previous slab, corrected for perspective by multiplying by the cosine of 45" (0.7). For example, a series of slabs 5 mm thick and 2 mm apart will appear along the reference face 3.5mm and 1.4mm apart, respectively (see Fig. 3). The outlines of columns are traced on to the framework in such a way that only the portions of columns not hidden by the outlines on preceding slabs appear. This basic sketch is then redrafted using shading or stippling to show in three dimensions the features of the columns.

Th in-sec tion studies

Thin sections are cut generally thicker than standard petrological thickness (0.03 mm); a thickness of 0.04-0.06 mm preserves sufficient contrast to


allow the structures and textures of the carbonate rocks to be studied at low magnification. Large thin sections (up to 15 cm square) are ideal for the study of columnar stromatolites. It is important to choose the least altered, representative samples for thin sections, so that the following features of taxonomic interest can be studied in detail.

(a) Lamina shape may be characterized as gently convex, moderately convex, or steeply convex, depending on the ratio of the height of a convex lamina ( h ) to its diameter ( d ) . It is useful to trace from thin sections a number of representative laminae from different parts of columns, and to measure their ratios of h/d . Laminae may also be characterized as domed, hemispherical, conical, sub-conical, rectangular or rhombic (i.e. flat-topped with parallel edges deflexed at about 90°, k d deflexed obliquely, respectively). On a finer scale, laminae may be smooth, wavy or wrinkled.

(b) Margin structure. The behaviour of laminae at column margins is considered to be taxonomically important. If they thin and turn downwards sharply so as to envelop the edges of the column, a structure termed a wall is formed, in which the internal lamination is sub-parallel to the margin. The margins of unwalled stromatolites may be smooth, bumpy, ribbed, ragged, or with pronounced overhangs. Pointed projections and projections set into niches are significant features if present.

(c) Microstructure. This term refers to the thickness, continuity and mutual relationships of laminae, the nature of their boundaries, and any microscopic internal structure. These features are best observed at low magnification under a binocular microscope. Observation under high magnification is usually necessary to determine the textures of the fine-grained carbonate comprising the laminae, but this is probably of less taxonomic significance, being a reflection of the diagenetic history of the rock. The petrographic study of stromatolites is also aided by staining thin sections with alizarin red to distinguish calcite from dolomite. Davies and Till (1968) described a variety of stains useful for distinguishing various carbonates.

(d) If the rock is wholly or partially silicified, there is a good chance of microfossils being preserved. Silicified areas should be scanned at medium magnification and the presence noted of any unicells, colonial unicells, filaments (with or without cross-walls, with or without branching, with or without differentiated cells) or any other organic structures. The orientation of filaments relative to the lamination may be important.

Peels

Acetate peels are a useful adjunct and are easy to prepare rapidly. A smooth-cut rock face is used; it may be necessary to grind it flat if it is rough with saw marks. The surface is etched lightly with 10% hydrochloric acid (for about 10 sec if limestone, 30 sec or more if dolomite). The surface may be stained with alizarin red and is then washed, dried, wetted with acetone, and

12 W.V. PREISS

an acetate sheet (preferably about 0.2 mm thick) is placed smoothly over it, making certain that all air bubbles are pressed out. After drying for about half an hour, the acetate sheet is peeled off bearing an imprint of the structures and textures of the rock. Davies and Till (1968) describe an alternative method of making peels from carbonate rocks, and list additional stains.

Statistical studies

Certain quantitative parameters can be measured, which when treated statistically, give additional criteria for the comparison and discrimination of stromatolite taxa. The following have been successfully applied.

(a) The degree of lamina convexity ( h l d ) , described above, is measured for as many laminae as possible in large thin sections. The results are divided into intervals of 0.1 units, and the percentage frequencies of these intervals calculated.

(b) The thicknesses of laminae may be measured in stromatolites which have sufficiently distinct lamination. A graduated eyepiece is used with moderate magnification, since stromatolite lamina thicknesses generally fall within the range 0.01-2mm. The results may be divided into thickness intervals chosen according to the predominant range of thicknesses (e.g. 0.05-0.10 mm intervals).

(c) In the case of Conophyton, which generally has very distinct laminae, two additional parameters have been used (Walter, 1972a; Preiss, 197313) but the first of these could be applied also to non-conical stromatolites.

(i) The ratios of thicknesses of adjacent light and dark laminae. These can also be represented as contoured frequency plots (thickness of dark lamina plotted against thickness of adjacent light lamina, and contoured for frequency of points per unit area).

(ii) Conophyton is characterized by a thickening and contortion of the conical laminae at their apices. The vertical structure that results from the superposition of these apices is termed the crestal zone. The coefficient of thickening is the ratio of the thickness of a lamina in the crestal zone to its thickness outside the crestal zone.

These parameters can all be plotted as histograms or frequency diagrams, enabling rapid comparisons with other forms. If more than a visual comparison is required, it may be necessary to calculate means and standard deviations for the data.

SUMMARY

Field observations are an essential part of all stromatolite studies, not only for geological mapping, environmental interpretation and determination of stratigraphic facings, but also for detailed taxonomic study. The thickness,


shape, lateral extent, synoptic profile and associated sediments should be determined for the stromatolitic bed. The nature of its constituent units and their vertical and lateral relationships must be described. Columnar stromatolites should be categorized according to shape and orientation of columns, style of branching, and margin structure.

Stromatolitic outcrops should be photographed prior to sampling. The size and number of samples to be taken depends on the size and variability of the stromatolites.

In the laboratory, serial sectioning and reconstruction are needed to show column morphology in three dimensions. Lamina shape, margin structure and microstructure are studied in thin section and acetate peels. Quantitative measurements of features of the lamination aid in the identification and comparison of stromatolites.

ACKNOWLEDGEMENTS

Helpful discussions with Drs. M.R. Walter and H.J. Hofmann are gratefully acknowledged. The paper is published with the permission of the Director of Mines, South Australia.


Chapter 2.2

GRAPHIC REPRESENTATION OF FOSSIL STROMATOIDS ; NEW METHOD WITH IMPROVED PRECISION

H.J. Hofmann

INTRODUCTION

An historical review of the ways of studying and illustrating fossil stromatolites (stromatoids) reveals a number of developments. The earliest methods included verbal description and drawings of structures in outcrop, reproduced by wood cuts and lithographs (e.g. Steel, 1825). Later, the petrographic microscope facilitated the study of microstructure, and the invention of photography and new printing methods opened up new ways for rapid improvement in disseminating information pictorially. The first Precambrian stromatolite to be described, Archaeozoon acadiense (Matthew, 1890a, p. 40), was already illustrated photographically.

The next major innovation did not come until the 1950’s with the method of “graphic reconstruction” from systematically spaced, serial longitudinal sections (Kry!ov, 1959b), adding a new perspective to previously existing techniques. A subsequent variant of the graphic reconstruction method is that of Raaben (1969b), using systematically spaced, serial transverse sections. In both Krylov’s and Raaben’s methods the columns are reconstructed by shading of the surface features, using properly positioned tracings of the serial sections as reference lines. These shaded-in surface features are used to differentiate some groups and forms from others. It has now become evident that the process of shading can be very subjective, varying from one person to the next (Hofmann, 1973), and that consequently the distinction between certain taxa based on this method may be artificial. For these reasons a more objective and precise method of reconstruction is proposed, one which also may eventually lead to automatic representation by computers. This is described below.

PRINCIPLES

Basically, the method involves the preparation of an isometric drawing from systematically spaced, serial longitudinal sections, and may be thought

16 H.J. HOFMANN

of simply as combining the elements of the two methods developed by Krylov and Raaben. However, instead of subjectively shading the reconstructed stromatoid surfaces, the columnar outlines are actually depicted by two sets of lines: systematically spaced vertical profile lines visible in a particular chosen orientation (usually a 45” dextral or sinistral oblique view), and equally systematically spaced horizontal contours. In this manner, the surface of each column becomes outlined by a grid, whose lines and points of intersection are not only accurately placed in space, but which also automatically provides a shading effect that is mechanically produced, and therefore more objective and precise than conventional shading (Hofmann, 1973, p. 364). Further superimposed stippling or shading is optional.

DESCRIPTION OF METHOD

Selection of sample

Samples selected should give at least two or three columns for reconstruction, and include columns as long as possible, showing at least two successive points of branching. However, there is often a limit to the size of the specimen that can be taken, handled, transported, and processed in the saw, and for many columnar stromatoids it is impossible to apply the method because the columns are too large. If several morphologic types are present in a bioherm, a range of variants is sampled. Relevant field observations on the positional, lithologic, sedimentologic, or other aspects relating the structures to enclosing lithosomes are essential for later interpretation of the reconstructed stromatoids.

Preparation of sample

It is advantageous to first make a cut transverse to the columns, to obtain an upper planar “horizontal” surface that will serve as a reference plane (x -y plane) from which later the grid lines for the horizontal contours are measured. This plane also gives the plan outlines of the concentric laminae, which are desirable for completion of the top surface of the reconstruction. A set of two converging, intersecting, or diagonal straight lines are drawn on the surface, with indelible ink markers to permit the exact relative positions of the slabs to be established later on, after sawing. Likewise, it is desirable, but not essential, to have a second, longitudinal section perpendicular to the first one before proceeding with the serial sectioning. This second cut can provide the “front” (x -z ) or “side” ( y - z ) face of the reconstruction and gives additional control for positioning the serial sections, as well as providing a further view of laininar profiles.

After these preliminary cuts have been made (and the surfaces traced


and/or photographed alongside a metric scale for a permanent record), the specimen is cut into slabs of uniform thickness along the third direction. The thickness of slabs chosen depends on the diameters of the columns, and should be selected so that the finished thicknesses of the slices are equal to, or multiples (2,3,4) of the widths sawed and ground away. This will allow later addition of intermediate profile lines by interpolation. For example, if the slab is twice the width of the cut, the two faces of the slab are traced on the final drawing, but later a supplemental line may be drawn intermediate in position between the two actually traced. In this way the uniform spacing between successive profiles is maintained. The caption of the final figure should specify whether the reconstruction is based on actual tracings only, or has supplementary profile lines.

Hints

Carbonate slabs as large as 10 by 30 cm can be cut 4-5 mm thick without breaking if a piece of foam plastic or rubber is placed in the receptacle underneath the slices being sawed off.

For large and heavy samples serial sectioning frequently requires remount- ing of the specimen. The block can be reinserted in the vice with nearly the same orientation if the last-cut face of the remaining portion of the specimen is aligned against the side of the saw blade.

Tracings

The outlines, and prominent laminar profiles of the columns to be reconstructed, as well as all reference lines or points, are transferred from the rock slice onto the drawing surface. The preparation of this tracing will often be facilitated if contrast is enhanced by introducing a film of water between the plastic drafting film and the surface to be traced. If the columns are still not readily visible they should be brought out by special treatment such as etching or staining. Eventually, it may be possible to obtain such outlines automatically with computers with capability for imagedensity evaluation and suitable output peripherals. The tracings are labelled in sequence.

.

Graphic representation (Fig. 1 )

(1) Fasten a sheet of millimeter-paper onto a drafting table. This is the reticle or control grid, which provides the coordinate system for the reconstruction; it should be large enough to fumish orientation for all points represented within the sample.

(2) Select a sufficiently large sheet of transparent drafting film or paper to carry the final graphic representation. This is the working sheet. (3) Place the working sheet over the control grid and fasten it with a piece

18 H.J. HOFMANN

Fig. 1. Method of graphic representation, illustrating reconstruction during stage of trans- ferring fourth profile from tracing 4 onto working sheet. Contour lines have not yet been drawn through lowermost three sets of hypsometric control points.

of drafting tape at each of the two top corners. These pieces of tape act as hinges and allow the working sheet to be lifted for insertion and removal of successive tracings between it and the control grid.

(4) On the working sheet select the position of the origin for the projection (point of intersection of the x - and z-lines of the first tracing used) at the intersection of two of the heavier lines of the control grid, and,mark it in pencil. This point should be so located that all columns to be represented can be accommodated on the working sheet.

(5) For a representation with an obliquity of 45", draw a line at 45" from the origin, to the right if a dextral view is desired, or to the left if a sinistral view is preferable. This line is the y-reference line that gives, in isometric projection, the y-reference points of each of the serial tracings, displaced towards the rear (up and to one side) by a factor of 0.7 (cos 45") of the real distance. Mark off and label these reference points.

(6) Select the columns to be represented by inspecting the serial tracings, and insert the tracing bearing the first oiltline in between the working sheet and the control grid. Move it until the reference point of the tracing exactly coincides with the origin on the working sheet, and the straight horizontal reference line coincides with the underlying grid line passing through the origin. The first tracing is now properly oriented for copying.

(7) Trace the outline(s) of the column(s) of interest onto the working sheet. (If the first face to be depicted is substantially within the column, then the laminae of this surface may also be traced so as to illustrate the laminar profile in longitudinal section.) Note that the horizontal grid lines below the


x-reference line intersect the profiles and are hypsometric lines (contours). Mark or note all points of intersection of the profile lines with uniformly spaced contours, the contour interval equalling the actual uniform spacing between successive vertical profiles.

(8) Remove the sheet bearing the first tracing and insert the succeeding one between the working sheet and the control grid. Superpose the reference points for this section on the tracing and working sheets and orient the tracing until the horizontal reference line is parallel to the corresponding x-coordinate of the underlying grid.

The procedure is considerably simplified if the serial sections are cut and ground so as to be an even number of miIlimetres apart (4 ,6 mm, etc.): This is because the distance between successive horizontal reference lines in the projection chosen is equal to % the true distance as measured on the original sample (cos 45" x cos 45") allowing the horizontal reticle lines to be used directly as guides. Thus, if the sections are 4mm apart in reality, the horizontal x-reference line of the second tracing lies along the grid lines 2mm above the first, but shifted 45' to one side. This is equivalent to moving the reference points 4 x cos 45" = 2.8 mm along the projection's oblique. y - coordinate.

Trace the profile(s) of the same column(s) as in the first tracing onto the working sheet as well as any newly appearing ones of interest, taking care to show only those parts of the profile lines that are visible in the oblique view chosen, and not "behind" the profile(s) drawn for the previous section. Again, mark or note all points where the equally spaced horizontal grid lines intersect the profile lines.

(9) The process is continued for all succeeding profiles of the columns of interest. After the third profile has been drawn, the set of lines joining points of equal elevation can be added. These are contour lines equidistant from the horizontal (x-y) reference plane, following the surface outline of the columns, and approximating the lines that would have been obtained by Raaben's

(10) To complete the view of the upper surface of the reconstruction, a projection of the original tracing of that surface can be used for highest precision. Alternatively, this surface can be reconstructed from the serial vertical profiles by joining correlative points previously positioned on each tracing along the upper horizontal reference lines. A projection of a second, vertical (lateral) section can be drawn in similar fashion if desired.

(11) Where needed, intermediate supplementary profile lines may be added later between those actually traced. These may become necessary where the thickness of the slabs is large. The thickness should preferably be a multiple of the width of the sawed and ground portion of the specimen. In the case where the spacing is double, one supplemental profile line is added by interpolation halfway between those representing the front and back of the slabs; where it is triple, two supplemental lines are inserted.

. method of transverse sectioning.

20 H.J. HOFMANN

SUMMARY

The result obtained by the method is an isometric diagram of vertical profiles and obliquely viewed horizontal contour lines, whose points of intersection lie on a systematic, three-dimensional coordinate system, on two sets of unifornily spaced and mutually perpendicular planes. This allows relative distances and mutual relations between morphological features to be perceived quantitatively on a figure that is largely free from reconstructor’s bias. It also allows for uniformity in presentation, eliminating the possibility that reconstructions made by different persons look different. Moreover, because of the high degree of spatial control, it is expected that the method will be amenable to computerization and lead to eventual automatic reconstruction of stromatoids. Once computerized, the data could conceivably be processed or manipulated in such a way that any particular view can be chosen by electronic reorientation, allowing for the selection of the one or two views which most clearly show the geometric attributes of the columns. On the other hand, the method may be superseded by statistical methods such as described in Chapter 2.5.


Chapter 2.3

BIOLOGICAL TECHNIQUES FOR THE STUDY OF MICROBIAL MATS AND LIVING STROMATOLITES

Thomas D. Brock

INTRODUCTION

Techniques are available for studying the growth and function of microbial mats in nature, and the data obtained with these methods are essential for the environmental interpretation of mat morphology and distribution.. This article briefly reviews these methods. It is hoped that workers will be encour- aged to carry out ecological investigations on living mats, since such investigations will provide much useful information for the interpretation of fossil material.

ESTABLISHING STATIONS

For most purposes, stations should not be established randomly, but should be placed intelligently so that the maximum amount of information possible can be obtained from the study. In marine environments, transects from supratidal to low intertidal or subtidal are preferable, and a series of such transects should be made within the general study area, so that the gener- ality of observations can be ensured. In thermal environments, stations are usually best established at various locations along the outflow channel of the spring, so that a variety of temperatures are obtained. Care should be taken to ensure that the various stations in the thermal effluent have similar flow rates, if it is desired to compare the specific effect of temperature. It should be emphasized that there is a large amount of variability in mats, even those formed by the same species, so that many replicate stations are desirable.

.

ENVIRONMENTAL MEASUREMENTS

In understanding the factors involved in the development of mats, it is ' essential to know the environmental conditions under which growth is

22 T.D. BROCK

actually taking place. This means that environmental measurements should be made periodically at the various stations. It should be emphasized that measurements of dead mats have no biological meaning. It is important to know that the mats under study are actually growing; the procedures by which the existence of growth can be determined will be outlined in a subsequent section. Measurements in growing mats should be made at regular intervals, so that any changes can be detected. Especially in marine mats, measurements should be made over a number of tidal cycles, and because tide also varies with time of the month, measurements should be made over 'several months. In habitats where seasonal changes in temperature or water level occur, it is desirable to make measurements over the whole annual cycle. It might be noted that a series of detailed measurements is essential since only a few days of unusual weather may be sufficient to drastically affect the environmental parameters. Although automatic recording equipment is theoretically the best, reliability is sometimes a problem, and data reduction can require a significant amount of time. In most cases, it is probably simpler and more reliable to make manual measurements at frequent intervals. Once the investigator has obtained some experience with a habitat, he can more intelligently select the required periodicity of measurement.

It should be emphasized that measurements must be made precisely at the locations where the mats are growing. Conditions may be quite different even a few centimeters distant. The most useful parameters are temperature, pH, salinity (in marine habitats), oxidation-reduction potential, oxygen concentration, sulfide concentration, and light intensity. Temperature measurements are most easily made with thermistor probes, which permit measurement at different levels within the mat. The pH can also be measuredcdirectly within the mat, using a combination electrode and a battery-operated pH meter (pH paper is generally unreliable, especially in salty or mineralized waters). It might also be desirable to measure the pH of water within the mat itself, as metabolic processes can result in pH values quite different from those of the surrounding water. When measuring pH in mats, care should be taken to ensure that changes in pH do not take place during expression of the' water. Losses of gases should be especially avoided, since loss of C 0 2 can lead to marked increases in pH. All pH measurements should be made directly in the field, or within a few hours in the laboratory on samples taken to ensure no loss of gases, or oxidation of sulfide. In mats which are completely submerged at all times, salinity of the water within the mat should be the same as the salinity of the overlying water. However, in intertidal mats, exposed to the air part of the day, salinity of the water within the mats may vary widely and may at times be much higher than the salinity of the overlying water. Techniques for measuring chemistry of water within mats have not been published. Water could be expressed by squeezing, or the mat could be removed from the water, drained, and cut into segments for analysis. Doemel and Brock (1976) have used the latter method successfully to analyze for sulfur species in mats.


A key factor in growth and preservation of mats is oxidation-reduction potential. Again, measurement must be made directly in the mat, and at the level in the mat which is of particular interest, since differences in oxidation- reduction potential may be great between surface and subsurface portions of a mat. Although more difficult to measure, oxygen and sulfide concentrations at different levels in a mat might also be measured, since these parameters will influence oxidation-reduction potential and are perhaps more meaningful biologically. Cell biologists use microelectrodes for this purpose which could be easily adapted for use in mats (Lavallee et al., 1969). Light intensity at the surface of mats can vary with depth and with angle of incidence of the sunlight. Especially in rocky habitats,’ there may be quite significant variations in light intensity with angle of incidence. Underwater photometers can be used to measure light intensity at the surfaces of mats in deep water. Light intensity at various levels within the mat are more difficult to measure, but are important in determining the depth in the mat at which photosynthetic activity ceases. Such light intensities can be measured by taking cores and cutting from them discs of various thicknesses. Each disc is placed on top of an opaque mask which contains a small hole through which light can pass. A light sensor such as a radiometer can be placed under the hole to measure light penetration. Because of the marked attenuation of light by even rather thin sections of compact mats, intense light sources must be used, and either sunlight or a high-intensity tungsten light is preferable. Natural sunlight is most environmentally relevant, although not especially subject to the investigator’s control. We have found 96-97% light attenuation by a mat 3 mm thick.

SAMPLING

Cylindrical or square cores can be cut for analysis, and for many purposes cork borers can be used. If the mat is rather loose and lacking in structure, it may fall apart when removed from the coring device. We have found that the structure of such mats can be retained if they are allowed to fall from the coring device gently into a container with melted 2% water agar (at 45-50°C). The agar is allowed to harden without disturbing the container, and then the core can be easily handled, dissected, or otherwise studied. Another way to preserve the structure of a delicate mat is by freezing, preferably on dry ice. Thin sections can be cut from frozen cores upon a block of ice, using a razor blade. Mats can be embedded in resin (we.. have found Durcupan quite satisfactory) and thin sections (around l p m ) made with glass knives using an ultramicrotome. For detailed study of vertical zonation in algal mats, microscopic dissection is useful. We have found it preferable to take a dissecting microscope into the field and perform the dissections right at the habitat.

If a series of replicate samples is to be taken from a single mat (for use,

24 T.D. BROCK

perhaps, in growth rate or photosynthesis studies as described below), it is essential that all samples be from the same population. Micro-environmental influences may result in populations of two or more types developing within a very small area, so that samples presumed to be replicates may not be. The only way to ensure that samples are indeed replicates is by making quantitative measurements on them, using some of the techniques described below.

Sampling of mats with vertical relief (cones, biscuits, nodes, etc.) as opposed to flat mats presents some additional problems. For small cones or nodes, we have taken large cores which include the base of the structures and then excised the individual nodes with micro-scissors. Large cones or biscuits can be easily collected, but each structure is heterogeneous and must be dissected carefully to sort out the various components. Dissection is best done directly in the field under a dissecting microscope.

TAXONOMY AND CULTURE

The taxonomy of the blue-green algae is in a very unsatisfactory state. As emphasized by Stanier et al. (1971), taxonomy of the blue-green algae can only be done by detailed study on a wide variety of pure cultures. Yet, the taxonomic distinctions which have been made and which are the basis of keys and species descriptions are based on rather unsophisticated microscopic examination of natural material. About the only thing that can be really told about blue-green algae from examination of natural material is the major group to which the alga belongs, and even here there is some uncertainty. Unsatisfactory as it must be to geologists, the taxonomy of blue-green algae is so uncertain as to be virtually completely useless. Indeed, some of the described taxa are not even blue-green algae (Brock, 1968) but photosynthetic bacteria (Bauld and Brock, 1973). For environmental and geological purposes, rather than to merely give the names, it is more useful to present a detailed microscopic description of the forms seen. Ideally, photographs of typical forms should be given, preferably using modern microscopic techniques, such as phase contrast or Nomarski interference contrast.

For detailed study of typical forms, cultures are essential, and for many purposes, these must be pure cultures. (There is an excellent recent book describing algal culture methods: Stein, 1973.) However, this is really the domein of the microbiologist, and any team studying a mat-forming system should include a microbiologist.

PIGMENTS

We have described in some detail the methods for measuring chlorophyll pigments in mats (Brock and Brock, 1967; Brock, 1969; Bauld and Brock,


1973) and details of these methods will not be repeated here. Briefly, the technique involves homogenizing mat material in acetone or methanol, centri- fuging, and reading the light absorbance at the wavelength appropriate to the chlorophyll of interest. All chlorophylls absorb in two regions of the spectrum, the blue and the red or infra-red. The absorption peaks which serve to distinguish the various chlorophylls are in the red or infra-red region (Table I).

TABLE I

Absorption peaks of bacterial and blue-green algal chlorophylls

Chlorophyll Absorption peaks

in vivo solvent extract

Algal chlorophyll a 680 nm 665 nm

Bacteriochlorophyll a 820 nm 770 nm

Bacteriochlorophyll c 747 nm 660 nm Bacteriochlorophyll d 725 nm 650 nm

Values listed are for the absorption peaks in the red or infra-red. All chlorophylls also have strong absorption peaks in the blue. Data are from Pfennig (1967), and Clayton (1963).

Bacteriochlorophyll b 1025 nrt,~ -

Blue-green algae have only chlorophyll a, which absorbs in acetone extract at 665 nm. We have found that blue-green algal chlorophyll is stable for several weeks if samples are placed in aqueous formaldehyde (4% final concentration) in the dark, but chlorophyll is not very stable in acetone extract, so that after extraction the absorbance must be measured within a day. If samples cannot be analyzed directly in the field, then cores should be placed in 4% formaldehyde (made up in the kind of water in which the organisms are growing), placed in the dark in a cool place, and extracted and analyzed as soon as possible in the laboratory. Vertical zonation of chlorophyll pigments will tell much about the structure and manner of growth of the mat (Brock, 1969), and can be determined by cutting segments of cores and assaying these separately.

Bacteriochlorophylls may be an important part of the photopigments of a mat, and must be handled somewhat differently from the method described above. In general, the photosynthetic bacteria are usually in layers underneath the algae, although in some Yellowstone hot spring mats this is not the case, the photosynthetic bacteria being right on top with the algae. These bacteria may even be present in the absence of algae (Doemel and Brock, 1974). The absorption spectra of the various bacteriochlorophylls are given in Table 1. Note that the absorption spectrum is different in solvent extract than it is within the cell (in uiuo). It is the in uiuo absorption spectrum that is ecologically relevant, but for quantitative studies the measurement of absorption in solvent extracts is usually carried out. Note, however, that the

26 T.D. BROCK

absorption peaks of two of the bacteriochlorophylls, c and d, are quite close to the absorption peak of algal chlorophyll a. Particularly, it is impossible to distinguish bacteriochlorophyll c from algal chlorophyll a by spectrophoto- metry of solvent extracts. This became important to us because some of the mats we were studying consisted of a blue-green alga and a filamentous photosynthetic bacterium (Bauld and Brock, 1973), the chlorophylls of which absorbed at exactly the same wavelength in solvent extract. Indeed, this fact initially prevented us from knowing that a photosynthetic bacterium was involved (Brock, 1968). Thus, to distinguish these two chlorophylls in mats, in uiuo absorption spectra must be obtained, since in uiuo the two chlorophylls absorb at quite different wavelengths. To obtain in uivo absorption spectra, it is necessary to eliminate in some way the light scattering of the material. We use the technique first developed by Pfennig (1967) of sus- pending the homogenate in a concentrated sucrose solution, which removes most of the light-scattering properties of the homogenate. It should also be noted that the absorption spectra of several of the bacteriochlorophylls are far in the infra-red (Table I), so that a spectrophotometer with a photocell with appropriate sensitivity is necessary.

PHOTOSYNTHESIS

In order to determine the activity of the organisms in mats, photosynthesis measurements are useful. The method of choice is the I4CO2 method, fully described by Brock-and Brock (1967), and widely used in our laboratory (Brock, 1967; Brock, 1969; Bauld and Brock, 1973; Doemel and Brock, 1974). Photosynthesis studies can be carried out on cores of mat, or on dissected segments of cores, if natural conditions should be maintained, or they can be carried out on homogenates. Homogenates have the advantage that a series of replicates can be prepared from the same material. However, for many purposes it is essential that the natural structure of the mat be maintained. As described by Brock and Brock (1967), incubation with 14C02 can be carried out on undisturbed cores, and then after the incubation is complete, the core can be homogenized and subsamples taken for separate determination of radioactivity and chlorophyll content. In this way, the radioactivity incorporated per unit chlorophyll can be determined, permitting normalization of values for different mats. It may be found that rate of photosynthesis per unit chlorophyll (photosynthetic efficiency) may vary markedly from mat to mat. This is a good indication that some of the mats are more active than others, an important piece of geomicrobiological information. Descriptions of how these techniques can be used to study the activity of mats under different situations can be found in Brock (1967) and Brock and Brock (196913).

It should be clear that 14C02 studies such as those just described present


data on the activity of the mat as a whole, but do not distinguish between the activity of individual organisms within the mat. To accomplish this, microscopic autoradiography must be done. This is done by placing samples of homogenized mat after incubation with 14C02 onto microscope slides, covering with liquid photographic emulsion, and allowing the radioactivity in the cells to expose the emulsion (M.L. Brock and Brock, 1968). By microscopic examination of the slides, the radioactivity can be localized over particular organisms. An example of a use of this technique with organisms from a marine environment can be found in Munro and Brock (1968). This technique would be especially useful in distinguishing between bacterial and algal photosynthesis in mats. Another way to distinguish bacterial from algal photosynthesis is by the use of the inhibitor 3-( 3,4-dichlorophenyl)-l,l- dimethylurea (DCMU). This agent inhibits photosynthesis in algae but has no effect on bacterial photosynthesis. At concentrations of 1-5 x lO-’M, virtually complete inhibition of algal photosynthesis occurs (Bauld and Brock, 1973). Because of the possible importance of bacterial photosynthesis in the anaerobic portions of mats, it seems essential to carry out studies with this inhibitor, to determine the relative importance of bacteria and algae.

GROWTH RATE

Probably the most sensitive indicator of environmental influence on mats is growth rate, and for many purposes measurement of growth rate will be desired. The most direct means of measuring growth rate is by carborundum marking, the carborundum layer providing a fixed horizon, on top of which the accreting mat continues to grow. By taking cores through the area marked by carborundum at various intervals, the rate of accretion can be obtained. For simple studies, merely measuring the thickness of the mat above the carborundum layer may be sufficient, but since compression effects may lead to different thicknesses for the same biomass, it is probably preferable to make some chemical measurement, such as chlorophyll content, organic- matter content, or protein content, to determine biomass.

Another way of assessing growth rate is by darkening an area of mat with an opaque shade, and measuring the rate of disappearance of the mat material (T.D. Brock and Brock, 1968). If the original mat was in steady state (probably not common in marine mats), the rate of growth would be balanced by the rate of disappearance (due to decomposition, erosion, and grazing), so that once growth is inhibited by darkening, only the loss processes can occur. Thus, by measuring the rate of disappearance of material, the rate of growth in the steady state can be determined.

28 T.D. BROCK

DECOMPOSITION

It is clear that preservation of mats can only take place if they have not previously decomposed. For geological interpretation, the whole process of mat decomposition is thus of great importance, and needs considerable study. Crude evidence of mat decomposition can be obtained by observation of gas- filled holes in mats, or other signs of degradation of cell material, but for any detailed studies, the rate of decomposition should be measured. This can most readily be done by use of carborundum marking. The procedure involves successive marking of the same mat with layers of carborundum at least twice. When the first carborundum layer is applied, the organisms will move on top and continue to grow, so that the mat will gradually build up .over the carborundum layer. Subsequently, another layer of carborundum is added. Cores are taken at the time the second layer is added, to determine the distance of separation of layers. Afterwards, more cores are taken. As decomposition proceeds, the two carborundum layers move closer together, and may eventually merge. One can assess the decomposition rate by measuring the distance between the two layers, and detailed quantification can be obtained by separating the material between the two layers and determining chlorophyll content, organic-matter content, or some other biologically meaningful parameter.

EXPERIMENTAL MODIFICATIONS

One of the attractive features of studying the ecology of mats is that they can be readily modified experimentally. These modifications reveal something about how the mats can respond to environmental change, and help in understanding the factors involved in the mat-forming processes. Darkening of portions of mat has been already described in a previous section. In addition, we have used neutral density filters to reduce the available light over portions of mat, with the aim of studying the adaptation of the algae and bacteria to reduced light intensities. This work has shown (Brock and Brock, 196913) that the algae respond to light reduction by increasing chlorophyll content. In the case of the photosynthetic bacteria, we have shown (Doemel and Brock, 1974) that if the light intensity is reduced to a sufficiently low level, algal growth is no longer possible, but growth of the filamentous photosynthetic bacteria will continue and a strictly bacterial mat will be formed. In some earlier work, we used experimental channels (Brock, 1969) within which themats could develop. Other possible modifications which may be of interest are: alteration in water-flow rate over the mat, changes in temperature by diverting flows of hot water (in thermal habitats), poisoning of the components of the mat with selective inhibitors, elimination of the mat completely to determine the rate of mat regeneration (see Brock and Brock,


1969a), transfer of portions of mat from one location to another, change in sedimentation rate, change in periodicity of flooding.

ACKNOWLEDGEMENTS

Research of the author in this area has been supported by research grants from the National Science Foundation (U.S.A.) (6B-35046).

NOTE ADDED IN PROOF

To avoid the necessity of measuring chlorophyll absorption spectra in uiuo, Madigan and Brock (in press) extracted mixtures of chlorophyll and bacteriochlorophyll in methanol, separated the chlorophylls by thin-layer chromatography, and eluted the spots. This technique permitted measurement of chlorophyll a and bacteriochlorophyll c in algal-bacterial mats.


Chapter 2.4

APPROACHES TO THE CLASSIFICATION OF STROMATOLITES

I .N. Krylov

IS A CLASSIFICATION NECESSARY?

The necessity to classify and name stromatolites inevitably arises in any attempts to use them for stratigraphic or paleofacies reconstructions. If a researcher thinks that the morphology (and the microstructure) of stromatolites is accidental and has no regular relation to the composition of the stromatolite-forming microbiota, or to their growth conditions, then he will neither classify stromatolites nor name them. That is why Vologdin (1962) and Korde (1950) rejected any classification of stromatolites. They studied and described only algae, describing stromatolites only incidentally as forms of the growth of these algae.

Another group of researchers is of the opinion that the shape of stromatolite constructions can be more or less regularly associated with their formative conditions: under similar conditions similar (or different) algal complexes could build similar constructions. Such constructions can be used for determining environmental conditions, or serve as local stratigraphic markers to correlate deposits within small regions or basins. In both cases the stromatolites should be described in some way and named. In classifying, “useful”, “operating” features are used: the extent of layers or the isolation of nodules and columns (Logan et al., 1964), the shape of the constructions (Robertson, 1960; Donaldson, 1963; Szulczewski, 1968) or a whole complex of features (Maslov, 1960). These workers rejected the system of type specimens and of the nomenclature-priority principle. The constructions may have names, but the nomenclature should not be Linnean. For example:

(1) Uniform constructions may have Latin names, even if they are binomial, but these names should differ from “true” paleontological names, at least in the method of transcription (not italicized or capitalized, but Gothic, cursive, boldface or extended - see Cloud, 1942; Hofmann, 1969a).

(2) Binomial Latin nomenclature should be replaced by a polynomial (Maslov, 1960) or uninomial system (Donaldson, 1963). (3) The name may consist of several letters forming an abbreviated de-

scription of the construction, for instance LLH (laterally linked hemispheroids; Logan et al., 1964).

32 I.N. KRYLOV

(4) Uniform constructions can be represented by letters or figures. Classi- fications of this type are usually temporary, pending the compilation of a fuller classification. So in papers by Korolyuk during the years 1956- 1958, stromatolites were designated by Roman and Arabic numerals (e.g. 11-3,1-2).

Finally, a third group of researchers assumes that stromatolites can be successfully used for biostratigraphy. Such an opinion is based on the fact that in very distant regions we find identical stromatolites in a closely comparable order of superposition. Such identical constructions can have the same names and would enjoy all the rights of Linnean nomenclature. These stromatolites could and should be classified within the framework of formal paleontological classifications including strict observance of the nomenclatural codes. I t is most convenient to use the provisions for fossils in the fnternational Code of Botanical Nomenclature (I.C.B.N.), as is the current practice of the majority of Soviet and many foreign researchers.

The complicated nature of stromatolites as constructions formed by an association of algae and not by single species is no obstacle for the application of binomial nomenclature. Latin names are a part of international scientific terminology. Their only purpose is to give a name that is uniformly understood by all specialists. Binomial nomenclature is used not only for desig- nating biological species, but also for the names of separate parts and organs of organisms (e.g. Musculus biceps), for chemical compounds and medicines (e.g. Oleum ricini), and for physiological processes and illnesses (e.g. Febris undulans). Latin binomial names are also widely used in geographical literature for regular plant associations (e.g. Pinetum myrtillosum).

A Latin binary name is unambiguous. The nomenclature should be stable and the name should have only one meaning, determined by a type specimen for a form and a type form for a group. The priority principle serves to protect this situation; new names are not accepted for previously named taxa. The I.C.B.N. specifically states (recommendation PB6) the necessity of a single binary name for taxa in formal classifications. Unfortunately, this recommendation has been repeatedly violated and one can find six diagnoses for the Gymnosolen group (Steinmann, 1911; Pia, 1927; Raaben, 1960; Krylov, 1963; Raaben, 1964b, 1969b), four diagnoses for the Boxonia group (Korolyuk, 1960c; Komar, 1964, 1966; Semikhatov et al., 1970), and so on. Urgent measures should be taken to re-establish and strictly protect the stability of the nomenclature. Otherwise it will become useless.

WHAT ARE. WE CLASSIFYING: SAMPLES, COLUMNS OR BIOHERMS?

The codes of nomenclature regulate only the rules on the publication of taxa and do not deal with the features on which these taxa have been established. In starting to classify stromatolites, different researchers working with very different material based their classifications on different features.


Fig. 1. Identification of a stromatolite depending on the size of the sample. The entire hand specimen: columnar branching stromatolite (e.g. Minjaria Krylov, Gymnosolen Steinmann); A, C, D = non-branching columnar stromatolite (CollurnnacoZZenia Korolyuk).

As a result several parallel and independent stromatolite classifications have been published.

The fundamental difficulties in the systematics of stromatolites result from the inevitably fragmentary nature of the material classified. As a rule, stromatolite bioherms and biostromes are rather large and of a complicated structure. And yet the classifications do not regard the bioherms as a unit, but treat only their separate fragments. There are no difficulties in classifying individual samples. The difficulties begin when we see the natural relationships between stromatolites of different gross morphology which have been referred to different taxa.

The determination of a stromatolite means that the researcher must give a name to a specimen that is placed on his desk. It is easy to see that the result will depend not only upon the sympathy he feels for a certain classification scheme, but also upon the specimen itself. Fragments A and B on Fig. 1 are parts of one construction. The researcher will call the stromatolite from sample A non-branching and will compare it with Cryptozoon, Colonnella, Collurnnacollenia, etc. The stromatolite in sample B will be called branched (Gymnosolen, Minjaria, Acaciella). If, however, sample B were divided into parts (C and D ) , the stromatolite (according to the observed features) would again be called non-branching. In this way, the application of a certain classification and, correspondingly, the choice of the group name, is sometimes determined simply by the size of the sample studied.

Fig. 2 represents a bioherm from the Lower Riphean Kotuikan suite in the north of Siberia. At its base we see a large dome-like nodule with the features of the group Collenia Walcott. Columnar stromatolites B, C, D, E , would be

34

Fig. 2. Stromatolites of different morphological groups in a single bioherm. The Kotuikan River, Lower Riphean. A = Collenia Walcott; B, C, D, E = Kussiella Krylov; F = Collumn- acollenia Korolyuk; G = Omachtenia Nuzhnov; H , Z = Stratifera Korolyuk.

referred to Kussiellu Krylov, the columnar “non-branching” stromatolite in sample F to Collumnacolleniu Korolyuk (= Colonnellu Komar), the columnar stratifonn constructions in sample G to Omuchtenia Nuzhnov, and <the separate fragments of the base (samples H, I ) to Strutiferu Korolyuk. Even in the monotypic bioherms consisting of columnar stromatolites only (Fig. 3), there are representatives of several formal groups. In the centre the columns are vertical and parallel (Gymnosolen); in the marginal parts they are inclined to horizontal (Tungussia?).

It is a practice to refer to the Conophyton group non-branching (Komar et al., 1965a) or extremely rarely branching (Walter, 1972a) stromatolites. In fact, we almost never see branching in a sample. In outcrops (Fig. 4), however, it is difficult to find a Conophyton without a branch.* Thus, Cono- phyton is both an independent group (if specimens 5-10 cm in size are in- spected) and part of the group Jacutophyton Shapovalova and Georginia Walter (if it is studied in exposures and in large - 50 x 100 cm - hand specimens).

However, the diagnostic features of a stromatolite can change even within one column. Following Maslov (1938), the majority of researchers understand under the name Conophyton, stromatolites consisting of conical layers and having a distinctive axial zone. Within one column, however, we see

* Not correct in my experience: Editor. (See also p. 525.)


Fig. 3. Monotypic bioherm of columnar stromatolites. The Lena River, Upper Riphean. In the centre the columns are vertical and subparallel (Gymnosolen Steinmann), in the marginal part they are inclined (? Tungussia Semikhatov).

(Fig. 5) an alternation of portions with dome-like ( A , C), conical ( B ) and flattened (D) layers. Correspondingly, the constructions in samples A and C should be referred to Collurnnucolleniu Korolyuk, in sample B to Conophyton Maslov, and in sample D to Plunocollinu Korolyuk. At the base and top of such constructions there are sometimes stratiform stromatolites that can be identified as Strutiferu Korolyuk or Irreguluriu Korolyuk.

A researcher will determine without any difficulty stromatolites in individual specimens A , B, C, D, but will be completely baffled when he sees them together. The reason for such confusion is purely psychological. He will be under the impression that constructions that form one column or one bioherm should surely have only one name. For this reason he is ready to change the original interpretation of the taxon and add new features to its diagnosis. To refer to the Conophyton group, for instance, the length from B to C (see Fig. 5), it is necessary to include in the group diagnosis not only the conical, but also the convex laminae. Simultaneously it is necessary to exclude the obligatory presence of the axial zone, since it is absent in sample C. In this way, the combination of the B and C lengths completely changes the original scope of the Conophyton group. Some researchers go far beyond that. In the Conophyton group diagnosis given by Komar et al. (1965a), 22 morphological features are quoted, not one of which is obligatory. This essentially means complete rejection of the use of morphology for the purpose of group distinction.

At the same time, we really can put on the desk identical specimens of Inzeriu tjornusi or Conophyton gurgunicum from Siberia, the Urals, North America and Australia. This means that the recognition of the Inzeriu, Cono- phy ton and Jucu tophy ton groups was legitimate and their rejection deprives

I.N. KRYLOV

Fig. 4. Bioherm with Jacutophyton Shapovalova, Middle Riphean. A. The Lower Tunguska River. B. The Maia River.

us of the possibility of making interesting and useful correlations of coeval deposits.

There is only one way out of the situation. Morphological varieties of stromatolites from different parts of the same bioherm can and must have different names provided the features actually observed in a sample fully correspond to the group diagnoses. In their turn, the group diagnoses must have one possible meaning, and that must be precise and stable.

THE MULTIPLICITY OF OVERLAPPING CLASSIFICATIONS

There are at least 70 valid groups (form-genera) of stromatolites described in the literature. They have been distinguished on the basis of different sets of criteria and belong to at least twelve independent parallel classifications. Each of these classifications has certain requirements of the specimens and


Fig. 5. Identification of portions o Jacutop..yton in different classifications. The entire hand specimen: Jacutophyton; A, C, F = Collumnacollenia Korolyuk; B = Conophyton Maslov; D = Planocollina Korolyuk; E = Baicalia Krylov.

first of all of their size. The determination of stromatolites for each of these classifications needs the use of different special methods. These twelve approaches are outlined below.

Classifications applicable to any stromatolites, independent of the type o f structure

For a determination it is sufficient to see one vertical section of the stromatolite (thin section, cut section, photograph of an exposure) of any size. The groups distinguished in this way are of a vast scope and need further subdivision.

(1) The distinction of groups is based on the presence of a single morphological feature. Thus, for a long time, all stromatolites with branching columns were referred to the group Gymnosolen (as understood by Pia, 1927;

38 I.N. KRYLOV

Cloud, 1942; Raaben, 1960), while to the group Cryptozoon (e.g., Pia, 1927; Rezak, 1957) went all stromatolites with columns that widened upwards. At the present time most such interpretations are of historical interest only, and these taxa (contrary to the rules) have received new and different diagnoses.

(2) All stromatolites, independent of their being columnar or stratiform, are divided according to one common feature - the shape of the lamination (Maslov, 193713, 1938). This led to the definition of the group Collenia (as understood by Maslov, and differing from the original description by Walcott, 1914) with dome-like layers, and Conophyton with conical layers and an “axial zone”. Later Maslov (1960) added an intermediate group Conocollenia.

Other classifications involve subdivision of stromatolites into types and subtypes, within which groups are distinguished. Five types of stromatolites are recognized: columnar, stratiform, nodular (Korolyuk, 1960c), columnar- stratiform and columnar- nodular (Krylov, 1963). There are different criteria for distinguishing groups within each type.

Classifications o f stratiform stromatolites

(3) Samples should include several (not less than two) adjacent tubercles and hollows. The determination is effected in a single vertical section. Groups are distinguished by the morphology of the layers. The best known of these groups are Stratifera, with a regular alternation of convex tubercles and concave hollows, and Irregularia, with an irregular non-inherited morphology of the layers (Korolyuk, 1960~).

(4) The groups are distinguished according to two criteria: lamination morphology and microstructure. Thus, the Gongylina group (Komar, 1964) stands close to Irregularia in morphology, but differs in microstructure. Micro- scopic study of a thin section is required for group determination.

Classifications o f columnar-stratiform stromatolites

( 5 ) The sample should include at least 2-3 columns. The determination is effected in a single vertical section. The groups are distinguished by the morphology of the columns and the shape of their laminae. In this way the following groups are distinguished: Schancharia and Collumnaefacta (Koro- lyuk, 1960c), Parmites (Raaben, 1964b), Omachtenia (Nuzhnov, 1967), Gruneria (Cloud and Semikhatov, 1969b), Dgerbia (Dolnik, 1969 in Dolnik and Vorontsova, 1971), and Tarioufetia (Bertrand-Sarfati, 1 9 7 2 ~ ) .

Classifications o f nodular stromatolites

(6) A vertical section that crosses the central part of the nodule is studied. In this way the groups Colleniella and Paniscollenia are distinguished (Koro- lyuk, 1960c), the difference between them being in the morphology of the laminae.

METHODOLOGY AND SYSTEM-4T:CS 39

(7) In addition to the morphology, consideration is given to the microstructure of the laminae, which demands a microscopic study of a thin section; Nucleella (Komar, 1966).

Classifications of col um nar- nodular stroma to1 ites

The determination requires a large sample that shows a substantial part of the nodule. I t is desirable to have .also a complete photograph or drawing of the outcrop. A vertical section is studied. In this way the groups Tinnia (Dolnik, 1969, in Dolnik and Vorontsova, 1971”) and Gaia (Krylov, 1971) have been distinguished.

Classifications o f columnar stromatolites

There are five independent classifications of columnar stromatolites that differ in principle.

(8) A vertical section that passes across the middle of a column is studied. Groups are distinguished according to the shape of laminae and to the nature of the column margins (Korolyuk, 1960a). This classification includes the following groups: Collumnacollenia, Planocollina, Linocollenia, Sacculia (Korolyuk, 1960a), Boxonia (Korolyuk, 1960c, non Komar, 1964), Cono- phyton (Maslov, 1938), Ilicta (Sidorov, 1960), Tunicata (Sidorov, in Koro- lyuk, 1968), as well as, apparently, Kuternia (Cloud and Semikhatov, 1969b) and Kasaia (Bertrand-Sarfati, 1 9 7 2 ~ ) .

(9) The groups are distinguished by the combination of three morphological features: (a) general shape of the columns (tuberous, sub-cylindrical), (b) type of column margins (smooth, bumpy, ribbed, walled or naked), and (c) character of branching. For group determination a hand specimen is needed that includes 2-3 columns that show the branching pattern. It is studied by the method of “graphical reconstruction” (Krylov, 1959b, 1963). To this classification belong the groups Kussiella, Baicalia, Jurusania, Minjaria, Inzeria, Pseudokussiella, Katavia (Krylov, 1962)t , Gymnosolen (Krylov, 1962; Raaben, 1964b, non Steinmann, 1911), Pitella, Turuchania (Semikhatov, 1962), Linella, Patomia, Vetella (Krylov, 1967a), Anabaria, Kotuikania (Komar, 1964), Svetliella (Shapovalova, in Krylov et al., 1968), Tenupalusella (Golovanov, 1970), Aldania (Krylov, 1969), Eucapsiphora (Cloud and

* Dolnik (1969) is a thesis for the Kandidat degree; see the author’s footnote below (Editor).

t These groups were described in a dissertation fully published in 1963 (Krylov, 1963). Inasmuch as references to these groups had appeared in publications before this paper came into print (Semikhatov, 1962), it has been accepted in the Soviet literature that the data were published in 1962, when the summary of Krylov’s dissertation was ’

published.

40 I.N. KRYLOV

Semikhatov, 1969b), Poludia (Raaben, 1964b), Lenia (Dolnik, 1969, in Dolnik and Vorontsova, 1971), Boxonia (Komar, 1964; non Korolyuk, 1960c), Tilemsina, Serizia, Nouatila, Tifounkeia (Bertrand-Sarfati, 1 9 7 2 ~ ) .

(10) To the three features enumerated above a fourth one is added-the microstructure of the layers. The requirements of sampling are the same. To determine a group, in addition to the study of the gross morphology (graphical reconstruction), it is necessary to study thin sections (Komar, 1966). In this way the groups Microstylus (Komar, 1966) and Glebulella (Dolnik, 1969) have been distinguished and new diagnoses have been given for groups described previously: Kussiella (Komar, 1966, non Krylov, 1962), Boxonia (Komar, 1966, non Korolyuk, 1960c, non Komar, 1964) and Kotuikania (Komar, 1966, non Komar, 1964). ’

(11) The groups are distinguished as regular combinations of morphologically different constructions. For determination in the laboratory, a very large hand specimen is needed, usually in combination with field photographs and drawings; sometimes a “graphical reconstruction” is used, but as a rule one vertical section is sufficient for confident determination of a group. I t is easiest to make such determinations directly in the field. Such determinations include Compactocollenia (Korolyuk, 1960) which is the combination of a nodule with a branching column, Tungussia (Semikhatov, 1962) which is the combination of inclined columns with vertical tuberous branches, and Jacuto- phyton (Shapovalova, 1968) which is the combination of an axial column corresponding to the diagnosis of the Conophyton group with branches of a definite morphology. (12)* The principles of classification used by Preiss (1972a) and Walter

(1972a) are fundamentally different from those outlined above. Any or all of the following characteristics are used in defining groups: mode of occurrence (e.g. shape of the bioherm), column shape, branching style, lamina shape, and microstructure. The definition of groups follows the prior definition of forms. To define a group, those features which characterize a recognizable natural assemblage of forms are employed. All of these features need not be present in every form of the group, or in every individual within a form. The diagnostic value of each characteristic may vary from group to group. This is a polythetic characterization of groups (Mayr, 1969, pp. 82- 83, 143). Groups defined in this way are Acaciella, Alcheringa, Basisphaera, Georginia, Madiganites, and Pilbaria (Walter, 1972a), and Kulparia Preiss and Walter (in Walter, 1972a). These authors are able to use many groups defined by other authors because of agreement about what constitutes a natural assemblage af forms.

So there are 12 classifications. Each of them is based on different features (or their combinations) and is independent. Without violating the rules, these

* This section was written by the Editor, at the author’s request.


classifications cannot be united into a single, general classification. Thus, for instance, to the Gongylina group are referred stromatolites that correspond in their morphology to the Zrregularia group, but which possess a specific microstructure. Consequently, Gongylina is part of the Zrregularia group and not a group of the same rank. To make the groups equivalent in rank, it is necessary to revise simultaneously the diagnoses of the groups Stratifera and Zrregularia adding to them microstructure as an extra diagnostic feature. How- ever, the author of the first description on the basis of the priority principle, has the right not to agree with such a revision. Thus, Komar (1966) has included into the Kussiella Krylov diagnosis (Krylov, 1962) the nature of microstructure, substantially reducing the content of the group. I do not agree with such an interpretation and continue to use the former diagnosis. This means that the literature now has a Kussiella Krylov, 1962 and a Kussiella Komar, 1966 (non Krylov). This violates the stability of the nomenclature and the names lose their meaning. Stromatolites with conical layers and an axial zone are distinguished in the 2 and 8 classifications as an independent Conophyton group; in the 11 and 12 classifications they can be part of the Jacutophyton and Georginia groups. Both Conophyton and Jacutophyton are valid groups, but they cannot be considered within one classification. Such a situation is reminiscent of the independent classifications existing in paleobotany for the different parts of plants: leaves, wood, fruit, roots, spores and pollen. Different parts of the one plant can and do have different names.

In this review I have not discussed the principles for distinguishing form taxa intermediate between group and type (supergroup, subtype of Korolyuk, 1960c; Krylov, 1963; Raaben, 196413; Komar, 1966; Bertrand-Sarfati, 1 9 7 2 ~ ) . The distinction of such taxa will be justified when we can in some way evaluate the taxonomic rank of the diagnostic features. For the time being we cannot do this; supergroups and su.btypes are arbitrary subdivisions distinguished only for the sake of convenience.

UNIFORMITY OR VARIETY?

Persistence of variety

Within one bioherm or in adjacent bioherms in the one bed it is possible to find morphologically varied constructions with the same microstructure that fit the diagnoses of different groups or even different types. This variety, however, is neither limitless nor accidental. In each case the range of variation is not great and is quite regular. For instance, together with Kussiella kussiensis Krylov there occur stratiform stromatolites (Stratifera), columnar-stratiform stromatolites (Omachtenia) and distinctive columnar-branching stromatolites (the description of which will soon be published). In bioherms with Gymnosolen ramsayi Steinmann there are virtually no stratiform and

4 2

Uussiella kussiensis ( M a s l o v ) Kry lov

Baicalia baicalica ( Maslov) Krylov

Inzer ia tjomusi Krylov

Jurusania cyl indr ica K r y l o v

Gymnosolen ramsoyi S te in rnann

Minjaria ural ica Krylov

Linella ukka Kry lov

Abundant Rare

I.N. KRYLOV

Fig. 6. Bioherm series.

columnar-stratiform stromatolites, but in the marginal parts there are distinctive inclined constructions that fit the diagnoses of Tungussia. similarly, in bioherms with Minjaria uralica Krylov, Pseudokussiella and stratiform constructions are present, and in bioherms with Linella ukka Krylov there is also Tungussia bassa Krylov. Best known is the striking persistence of Conophyton and Baicalia combinations in Middle Riphean Jacu tophy ton bioherms.

For the majority of bioherms the intergrading constructions of which they consist occur in a regular arrangement. Such bioherm series comprise all the principle morphological varieties that make up one bioherm or several bioherms of the same type in one bed and which have the same type of microstructure (or set of microstructures). Each bioherm series can be interpreted as an aggregate of morphological modifications all formed by one community or species of blue-green algae. The bioherm series of some leading forms are shown in Fig. 6. They have been compiled on the basis of a study of biohenns from which the holotypes of forms originated and they have been checked in a number of other regions of the U.S.S.R.

The establishment and study of bioherm series provides additional possibilities for comparison of stromatolites from different beds and regions: one can compare not only constructions of the same formal type but also their regular, persistent, stable associations.


CONCLUSIONS

(1) The necessity of classifying and naming stromatolites is imposed by the possibility of using them in practical work; the same constructions can and must have the same names.

(2) Stromatolites can be classified formally and named in the Linnean fashion in conformity with all the rules of nomenclatural codes. Most convenient is the International Code of Botanical Nomenclature in its paleobotanical part.

(3) Morphologically different constructions from different parts of the same bioherm can have different names .in accordance with diagnoses of the taxa described in the literature.

(4) The morphological variety of stromatolites within one bioherm has definite limits. I t consists of regular, persistent associations of definite constructions. These associations (bioherm series) are stable and can be recognized for the majority of the Riphean key forms of stromatolites in various regions of the U.S.S.R.

( 5 ) The use of Linnean nomenclature is meaningful only when the strictest rules are observed to maintain its stability. This is a problem of utmostim- portance in the present state of stromatolite studies.


Chapter 2.5

STROMATOID MORPHOMETRICS

H.J. Hofmann

INTRODUCTION

The purpose of this chapter is to indicate the methods, advantages, and wide possibilities of rapid-automated, quantitative image-analysis in the study of geometric attributes of stromatolites. Technological advances in electronics in the last ten years have led to the development of equipment well suited to quantitative stromatolitology, opening up a whole new approach to stromatoid taxonomy. One can now foresee the possibility of a rapid improvement in precision in identifying, differentiating, and defining stromatoid groups and forms, because such information can be obtained rapidly in numerical terms. This is not to say that more orthodox methods will be abandoned, but the advantages of the new approach are clear. Not only can the degrees of similarity, differences, and variations of specimens and taxa be expressed quantitatively, but such information becomes amenable to computer processing, storage, and retrieval. It also allows for graphic plots of the features of interest. Moreover, inter-language communication is facilitated because numbers and graphs are neutral and do not need translation, whereas ordinary descriptive adjectives or nouns sometimes have no exact equivalent in another language, or have different meanings in the same language in different parts of the world. The method is also applicable to stromatoids which are too large to be handled physically.

PRINCIPLES AND PARAMETERS

Image-analysis is the extraction and processing of quantitative geometric and densitometric data from images such as photographs, drawings, or thin and polished sections. The object is placed on a stage (microscope or epidiascope) and scanned by a television-like camera under transmitted light (thin sections, transparencies) or reflected light (photographs, drawings, polished sections). The converted electrical signals from the scanner are fed to a processor capable of measuring a number of parameters, and then to a

46 H.J. HOFMANN

Pe

AREA A

PERIMETER Pe

VERTICAL FERET DIAMETER

Fh HORIZONTAL FERET D I A M E T E R

VERTICAL INTERCEPT IF, + a + b) I v

HORIZONTAL INTERCEPT ( F h + c + d) l h

M A X I M U M VERTICAL C H O R D C”

F”

ch M A X I M U M HORIZONTAL C H O R D

FULL FEATURE COUNT

E N D C O U N T

N f

Ne

Fig. 1. Basic descriptor functions.

TV-screen, where the feature selected for measurement is displayed pictorially as well as numerically. The instrument has the capability to obtain the basic descriptor functions shown in Fig. 1. It therefore performs its function as an electronic planimeter (A), curvometer (Pe) , tally register ( N ) , and ruler (F, I , C); an additional module allows it to serve as an optical densitometer. Articles on the image-processing and pattern-recognition techniques and their applications are available in the non-geological literature (e.g., Cole, 1970; Gibbard et al., 1972).

In principle, an image is electronically resolved into a large number of equally spaced points on a very closely spaced grid, corresponding to successive positions of the electron beam following horizontal scan lines, and coordinates perpendicular to them. (The active frame on the display screen comprises a total of 500,000 points or positions of the beam.) The situation is comparable to one where an image is placed under a properly oriented, transparent millimeter-grid paper and measurements are made manually by counting or measuring the number of square millimeters or points of intersection. The latter is also a substitute method that can be used to obtain approximately the same results in cases where access to electronic analyzers to perform the measurements is not available.

To convert the electronic point-counts on the instrument to metric units, a conversion factor can be obtained by placing a metric scale on the stage and comparing its image with the instrument scale on the display monitor.

To obtain a value for A , the function switch .on the computer module is set on “AREA” and the total number of points falling within the image are counted electronically. In the manual method one can count the number of square millimeters covered by the image, or if preferred, use a planimeter.

A value for Pe is obtained by electronically counting all points on the periphery, with the function switch on the computer module in the “PERI- METER” position. The length of the periphem can also be obtained manually with the transparent millimeter grid over the image by a laborious counting and angle-correction process, by the use of a pair of dividers, or by applying a standard curvometer (map-measurer).


The Feret diameter (F) is a measure of the maximum dimension of an object in a given direction. If the orientation of the stromatoid column on the stage is made to correspond to that during growth, two values can be assigned to it in a section: Fh and F,. These two diameters constitute the sides of the smallest rectangle (tangent on four sides) containing the image in that particular orientation. For the manual method these measurements are easily made with the millimeter grid or a ruler.

The intercept ( I ) is a measure of the total projected lengths of convexities (projections, branches) or concavities (re-entrants) of a feature in a given direction. It should be emphasized that these are projection lengths in the horizontal or vertical direction, and not actual lengths in any diagonal direction. In the manual method these projected lengths are obtained with a ruler or the millimeter grid, by measuring the lengths of all convexities in a given direction and integrating these values. As with the Feret diameter, two main intercepts are of interest, Ih and I , , along the x- and y-coordinates, respectively.

The size distribution of chords (C), of which the horizontal chords (apparent diameter of stromatoid columns) are of particular interest, is obtained with the programmed automatic sizer. Data for histograms or cumulative curves are generated by consecutively scanning the image with settings of “INTERCEPT SMALLER THAN”, at intervals in arithmetic progression, starting from zero until the maximum chord (called Martin’s diameter in the literature) for that particular orientation has been attained. The results are accumulated at some output peripheral such as a digital desk-top printer, or transferred directly to a graph on an x-y plotter or a computer for statistical analysis. The size distribution of the chords can likewise be found manually by measuring with a ruler or with the millimeter grid at selected intervals.

With the function switch in the “COUNT” position, the instrument can give a tally of the number of features ( N ) observed in the field of view. Two varianits can be selected: full feature count (Nf), where the total number of images in the field of view is measured, and an end count ( N , ) , which measures the number of convexities in a given direction. In the example in Fig. 1, Nf = 1, because only 1 whole feature is shown, and N , = 3, because the feature has 3 ends (convexities or branches).

Ultimately, the objective is to use a combination of two or more of these functions mentioned above for the description and comparison of the geometric features of stromatoids. For example, the ratio A/Pez is a shape index measuring an image’s departure from “circularity”; the ratio 2A/Pe gives an approximation of the “mean width” of a feature; and ratios such as Iv/Fv and and Ih/Fh are reentrance factors, reflecting the degree of branching or accretion vector variability. F,/Fh indicates the over-all elongation or orientation of ordered objects.

48

EQUIPMENT

H.J. HOFMANN

The image-processing unit used in this study is a QUANTIMET 720, manu- factured by Image Analysing Computers (IMANCO) of Cambridge, England. At the time of study, the instrument was equipped with a Vidicon camera, standard detector and TVdisplay with a “live” frame of 500,000 points (800 x 625), an epidiascope, a Leitz Orthoplan microscope, and modules for variable frame and scale, MS3 computer, amender, automatic shading correc- tor, 16-way programmer, output drive, and a Hewlett-Packard 9810 A desk- top calculator. Other modules, including a pattern-recognition system, and an x-y plotter, can be added, but were not available.

APPLICATION

For purposes of illustrating the application of instrumented image-analysis to stromatolitology, the following essential stromatoid characteristics were selected, compatible with the capability of the instrument used in the study: silhouette or profile view (longitudinal section), diameter variations of columns (horizontal chords), plan outline (cross section), laminar profile, and microstructure. I should point out that this study represents a first effort in this field as applied to stromatolites, and that the results presented here are of necessity exploratory in nature and incomplete. A series of parameters is suggested in Figs. 2-6, but other combinations of functions could also have been used. Future work will show which of the parameters are the most efficient for the discrimination of individual stromatoid attributes.

Stromatoid silhouette (longitudinal section)

Fig. 2 presents the results of an analysis of 24 silhouettes selected from the published literature (13), unpublished photographs (3), and artificial images (8), to cover a wide variety of shapes. The abscissa selected is the quotient Fv/Fh, which represents the overall orientation or elongation of the stromatoid portion considered. The quotient Iv& indicates the relative predominance of vertically directed reentrants (or branches) over laterally directed irregularities such as corrugations or oblique branches. As can be seen by comparing the upper and lower halves of Fig. 2, plots near the upper left of the diagram tend towards longiramose and multiple furcate-branching patterns (Kussielka, Boxonia, etc.), whereas those towards the bottom right tend towards breviramose groups ( A rchaeozoon, Jacu tophy ton, and Inzeriu) and columns in which lateral irregularities predominate over vertical ones. Fea- tures without peripheral concavities, such as circles and rectangles, fall on the line having a value of 1 for the combined ratio. It is evident that a scatter plot for many profiles will be required to delineate each taxon statistically, and to locate the center of gravity.


1 2 F, 4 -

B Fh

0 F F I 2 6 3 A F h

Fig. 2. Comparison of stromatoid silhouettes. A = B = c = D = E = F = G = H = I = J = K = L = M = N = o = P = Q = R = s = T = U = V = W = X =

Acaciella australica (Walter, 1972a, p. 117). Anabaria radialis (Komar, 1966, p. 85). Basisphaera irregularis (Walter, 1972a, p. 135). Boxonia lissa (Komar, 1966, p. 81). Acaciella australica (Walter, 1972a, p. 120). Anabaria radialis (Komar, 1966, p. 85). Gymnosolen furcatus (Komar, 1966, p. 89). Kussiella hussiensis (Komar, 1966, p. 76). Tungussia sp. (Hofmann, unpublished). Cryptozoon, schematic. Ottonosia, schematic. circle. Osagia, schematic. Parmites concrescens (Raaben, 1964b, fig. 1, p. 104). Baicalia burra (Preiss, 1972b, p. 82). Kotuikania torulosa (Komar, 1966, p. 87). Archaeozoon acadiense (Hofmann, unpublished). Jacutophyton sp. (Krylov and Shapovalova, 1970b, fig. 5, p. 40). corrugated column. constringed column. Patomia ossica (Krylov, 1963, p. 48, fig. 26-g). Inzeria, schematic. Gymnosolen sp. (Hofmann unpubl.). Colonne lla, schematic.

Diameter variations of columns (horizontal chords)

The same profiles used in Fig. 2 were subjected to chord-size analysis, with the results presented in Fig. 3. The cumulative frequency curves were all standardized in such a manner that the 100% ordinate has the same length for all, except for profiles L and M, which, for economy, are at a 60% reduction. The cumulative curves are arranged according to the decreasing

50

A O o 6 ,

Pe; 04

H.J. HOFMANN

A 2 0 6 M

*++- 9 10 11 12

ii i o 12

I O 2 0 4 F 0 6 l o

5 Fig. 3. Comparison of cumulative frequency curves of horizontal chords for profiles illustrated in Fig. 2. All vertical scales are identical, except for L and M which appear at a 60% reduction, C = chord size; Ch = maximum chord (Martin’s diameter); positive and negative skew are also illustrated.

Fig. 4. Comparison of plan outlines. A = area; Pe = perimeter; I = composite intercept; F = composite Feret diameter.

rapidity with which the value of the maximum chord ( c h ) is attained. For example, the chords in profiles B and W attain their maximum value 10 times faster than those of a circle ( L = 1.0). Steep and uniform slopes represent good sorting; breaks in slopes of curves reflect polymodality due to branching or corrugations; positive skew indicates predominance of smaller diameters over larger ones, and negative skew shows the reverse. A comparison of patterns of the cumulative curves for some groups shows certain similarities which correspond to similarities in the group’s position on Fig. 2 (e.g., 0, R, V ) . On the other hand, there are groups exhibiting similarities in the cumulative curves, but these show differences in Fig. 2 (e.g., B and F ) .

Plan outline (cross-section)

The parameters selected for differentiating plan outlines are A / P e z , which measures the “circularity” of an image, and a composite margin-simplicity factor, F/I (Fig. 4). Comparison of the 12 images with their plotted circularity values shows that a decrease in this value is associated with a departure from the circular shape. For a perfect circular disc the value is nr2/(2ar)’, or 0.0769. This value is plotted in Fig. 4 beside the circle symbol; the value for a square, which is 1/42 (or 0.0625), is also plotted.

The marginal simplicity, FII is a composite ratio. I t is obtained as follows:


__Ic

1 5

I 2

- FV Pe

Fig. 5. Comparison of laminar profiles. F = Feret diameter ( h = horizontal, u = vertical); Pe = perimeter; A = area.

_ - - F v + F4S + Fh + F13S I v + I45 + Ih + I135

where the subscripts 45 and 135 refer to values in the respective 45" and 135" positions attained by rotating the graduated instrument stage. These values are necessary for images with reentrants to compensate for possible erroneous values due to orientation: for example, a cross in the upright position (Greek cross) will give Fv /Iv and Fh /Ih values of l, whereas the same cross in the diagonal position (St. Andrew's cross, or x) yields values smaller than unity for these ratios.

Laminar profile

Laminar profiles can be distinguished on the basis of their Feret diameters and perimeters (Fig. 5). Inasmuch as the laminae are generally thin, the length of the lamina can be taken to be Pe/2. By comparing this ratio, or Pe directly, with the two sides of the smallest circumscribed rectangle (Fh and Fv), an mdication of the lamina's sinuosity is derived. By also plotting the lines of the related, complementary ratio Fh/Fv on the same graph, one can determine its general first-order laminar curvature, which is related to the E-value proposed by Hofmann (1969a) (F,/FV = 2E, for elliptical curvature).

For the ten profiles given in Fig. 5, nos. 2, 6, and 8 (and similarly nos. 3, 7,

52 H.J. HOFMANN

s ! 7

8 Ad Pe2

0 08

006-

Nf .

0 04.

0 0 2 .

.a 10 70

Pe

? 2

- 30

I . 80 100%

’O 100 - Ad Al

Fig. 6. Comparison of microstructures. Ad = area of dark lamellae; A1 = area of light lamellae; Pe = perimeter; Nf = full feature count; J = adjustment factor. The vertical scales in the two graphs are identical and give the dimensionless “average shape” index for dark lamellae. The horizontal scale in the upper graph is calibrated in relative length-units, which in this case, are unmodified point-lengths, the raw data not having been converted to millimeters (i.e., J = 1). The abscissa in the lower graph shows the percentage area occupied by dark lamellae. Microstructures 1-7 represent forms of Conophy ton taken from Komar et al. (1965, p. 22); microstructure 8 is that of Spongiostroma maeandrinum from Gurich (1906, pl. VI, fig. 1).

9, and 10) have approximately the same F, /F , ratios, but they distinguish themselves by the fact that their lengths depend on the crinkliness. The degree of crinkliness of the profiles could also be identified by using other parameters, such as the ratio N,/Nf, where Nf = 1, and N , represents the number of reentrants, but this parameter may be unsatisfactory for profiles with F,/F, values smaller than about 5 , where convexities are laterally directed and not detectable with the instrument if the analysis is performed with only one orientation.

Microstructure

Eight microstructures were selected from the literature for analysis, and some of the results are plotted in Fig. 6. To obtain an expression for the “average shape” of the dark lamellae (subscript d), a variant of the non- dimensional ratio A/Pez for circularity was used, which takes into account the fact that there was more than one feature in the field of view.

The “average shape” of the light lamellae can likewise be determined by


obtaining readings in the instrument’s “complement” mode. A second parameter, (2Ad/Pe)J, is a modified version of the “mean width”, 2A/Pe, which has a dimension (number of points counted, or millimeters). The adjustment factor J in this expression serves to take into account the final optical magnification of the image, and the instrument-constant relating the number of points counted to an actual length in millimeters; it can also be used to make the values obtained with instruments comparable to the L , and L , values previously used in the literature (Komar et al., 1965). For purposes of illustrating the application, all the images in Fig. 6 were considered to be of the same magnification to eliminate the effect of scale and to permit comparison of the results in graphic form. It is clear that the microstructures with the thicker lamellae (nos. 1, 6, 7) fall to the right and the thinner ones to the left, just as suspected by visual examination of the images.

A third parameter used is the percentage of the image covered by the dark lamellae, 100Ad/A, + A , . This function reflects the predominance of either light or dark lamellae, and is somewhat comparable to the L , / L ratio already in use (Komar et al., 1965).

Further analysis, such as thickness-sizing similar to that done for.Fig. 3, for both light and dark lamellae, can make the data directly comparable to conventionally obtained data on L and L , already in the literature.

CONCLUSION

Instrumental image-analysis of stromatoid characteristics is not only feasible but desirable for precision and efficient communication in taxonomic matters. Laborious serial sectioning and graphic-reconstruction techniques for identification may become unnecessary if large, mutually more or less perpendicular surfaces are processed to extract statistically significant morpho- metric data. This is analogous to the current method of describing the geometric characteristics of clasts in arenaceous sedimentary rocks from thin sections, where the graphic reconstruction of individual grains is not made, but where one desires rather to have statistical expressions of a sampled population. Because of the great polymorphism exhibited, even by individual stromatoid groups, the reconstruction of parts of a column, or, for that matter of ten partial columns, rarely allows one to perceive the extent of the morphologic variations. Objective quantitative data offer perhaps a better solution to the taxonomic dilemma. Other branches of geology and paleontology would also benefit from the ”application of quantitative image- analysis.

54 H.J. HOFMANN

ACKNOWLEDGEMENTS

I thank Richard Sackhouse for advice and assistance with the operation of the QUANTIMET image-analyzer, and Marc Grondin, for supplying literature on image-analysis. The study was financially supported by the Canadian National Research Council under Grant No. A7484.


Chapter 3.1

CALCRETES AND THEIR DISTINCTION FROM STROMATOLITES

J.F. Read

INTRODUCTION

This chapter describes characteristics of calcrete (caliche) deposits and outlines criteria that may be used to differentiate these from superficially similar cryptalgal sediments. Calcrete is a secondary accumulation of fine-grained carbonate (typically cryptocrystalline calcite) formed in soil profiles. Struc- tures and fabrics of many calcrete deposits resemble stromatolitic structures and fabrics, particularly flat-lying cryptalgalaminates, laterally linked hemispheroids and oncolites (Logan et al., 1964; Aitken, 1967). It is important that the two types of deposit are differentiated, as cryptalgal sediments are typically tidal, and to a lesser extent, subtidal deposits (Logan et al., 1964; Hoffman, 1974), whereas calcretes are products of subaerial weathering and are associated with unconformity (Multer and Hoffmeister, 1968). As tidal environments are transitional landward into non-marine terrestrial environments, tidal deposits containing cryptalgal structures may be overlain by and interlayered with non-marine calcrete horizons (Multer and Hoffmeister, 1968; Bernoulli and Wagner, 1971; Colacicchi et al., 1975) a feature that may further hinder distinction of cryptalgal and calcrete deposits. The growing list of documented ancient calcretes includes those described by Swineford et al. (1958), Runham (1965, 1969), Thomas (1965), Amsbury (1967), Nagtegaal (1969), Bernoulli and Wagner (1971), Wardlaw and Reinson (1971), and Colacicchi et d. (1975); this list indicates that calcretes do possess diagnostic features that distinguish them from cryptalgal sediments.

SETTING

Calcrete profiles develop as a result of soil-forming processes in the vadose zone. Calcrete is common in arid and semi-arid environments and may also be developed in humid areas with over 150 cm annual rainfall provided the rains are interspersed with periods of high evaporation (Multer and Hoff- meister, 1968; James, 1972). Host sediments are commonly carbonate-rich

56 J.F. READ

(Multer and Hoffmeister, 1968; Dunham, 1969; James, 1972; Read, 1974) but calcrete may also develop on sandstones and other sediments that lack carbonate (Gile et al., 1966; Amsbury, 1967; Logan et al., 1970b, p. 52). Calcrete is commonly developed on subaerially emergent sediments of modem carbonate seas (Newel1 and Rigby, 1957; Multer and Hoffmeister, 1968; James, 197.2; Read, 1974), and because of its stable low-Mg calcite mineralogy and resistance to erosion, it is likely to be preserved following marine transgression (Dunham, 1969). Where calcrete is intercalated with marine horizons, its presence indicates unconformity (Multer and Hoffmeister, 1968).

DEFINITION, MINERALOGY, AND FABRIC

The term “calcrete” is used to describe accumulations of cryptocrystalline carbonate resulting from soil-forming processes. Calcrete (partly synonymous with caliche and kunkar) is typically composed of small (1-4 pm) interlocking crystals of cryptocrystalline low-Mg calcite, that may be interlaminated with flat-lying acicular calcite laminae (James, 1972); Scholle and Kinsman (1974) describe unusual calcrete mineralogies consisting of cryptocrystalline and fibrous aragonite and cryptocrystalline high-Mg calcite. Calcrete occurs as laminated coatings around grains, as laminated sheet-like deposits, as intergranular cements, and it also replaces carbonate grains. I t may be associated with other secondary minerals such as opaline silica and cryptocrystalline dolomite (Swineford et al., 1958; Reeves, 1970; Read, 1974). Insoluble residues from calcrete are composed of angular detrital quartz and a brown amorphous pigment containing iron; the pigment, which occurs as disseminated particles that are concentrated in certain laminae, is responsible for the laminated appearance and brownish color of much calcrete (Siesser, 1973; Read, 1974); in other calcretes, color-banding has been ascribed to concentration of organic material in calcrete laminae (Multer and Hoffmeister, 1968; James, 1972).

Cements (other than cryptocrystalline forms) associated with calcretes are commonly radial granular to blocky sparry calcite. James (1972) describes unusual cements including randomly oriented calcite needles, tangential calcite needles, and “flower spar” (bunches of elongate single crystals of calcite 20-40 prn long). Aragonitic and high-Mg calcite calcretes described by Scholle and Kinsman (1974) are associated with coarse fibrous aragonite cements that may form thick “microstalactitic” coatings on undersides of grains or occur as “meniscus” cements; later cements consist of microcrystalline to slightly pelletal high-Mg calcite.

Although most calcrete calcite is relatively stable, recrystallization to microspar (6-10pm mosiac) may occur (James, 1972). Some ancient calcretes still retain many primary features even after dolomitization (Dunham, 1969; Wardlaw and Reinson, 1971).

ABIOGENIC STROMATOLITE-LIKE STRUCTURES 57

I BARBADOS SHARK BAY F LORI DA a,----:.. ..ir* Pi sol i tic, r

I E I c9

u r y a , , , c I lC l l

clayey to organic poor

Pleistocene reefal limestone grainstone and mudstone

..:.:.:........ ....... oolitic .

......................... ............. 1 i thocl asti C

.......................... grainstone ...................... .......................... .......................... ............. ............... A Organic rich

Pleistocene Key Largo Lst. and Miami 001 i te

Calcreted limestone with ......... 0 secondary ‘mudstone’ fabrics Laminar calcrete Original host sediment

Fig. 1 . Typical calcrete profiles; adapted from James (1972), Read (1974) and Multer and Hoffmeister (1968). Loose soils in upper parts of profiles commonly absent.

CALCRETE PROFILE AND MORPHOLOGY

Calcrete formation commonly results in diagenetic modification of primary sediment fabrics and structures; calcrete structures include calcrete pisolites, calcrete ooids and calcrete pellets, laminated calcretes, lithoclast breccias and structures resembling Stroma tactis. Calcrete formation may also result in development of secondary “mud-support” fabrics. Among the various calcrete structures, it is mainly the calcrete pisolites and the laminated calcretes that most resemble cryptalgal and stromatolitic structures.

The principal types of calcrete phenomena may be related to position in a weathering profile that ranges from less than a meter to several meters in thickness. In some inland areas, calcrete profiles may reach 50 m or more in thickness (Brown, 1956). The calcrete deposits show considerable variation on a local scale and between deposits formihg in widely separated localities; Quaternary calcrete profiles from Barbados, Shark Bay (Western Australia), and Florida are shown in Fig. 1; unfortunately, there is little published information concerning ancient profiles. Profiles are commonly complicated as a result of multiple profile development, where zones are repeated due to old calcrete profiles being buried by prograding mobile dunes or by transported

58 J.F. READ

PLATE I

Calcrete pisolites and ooids i. Thin section, calcrete-ooid grainstone cemented by sparry calcite; crossed nicols. Shark

aa,r. B. Thin section, calcrete ooids, plastic impregnated sample; plane light. Shark Bay. C. Calcrete pisolites exposed in intertidal rock platform. Shark Bay. D. Polished slab, calcrete pisolites (1-4 cm diameter). Nuclei are mainly lithoclasts of calcreted eolianite, some of which are blackened. Inter-pisolite material consists of calcrete ooids with intergranular cements of calcrete and sparry calcite.

soils. Calcrete development in the younger sediments tends to cause repetition of horizons or formation of complex superimposed zones. Erosion may cause superimposed horizons by development of a new profile of weathering in an older truncated profile. Incomplete profiles may develop where durationof weathering has been insufficient to form all horizons, where local factors such as ground slope have inhibited profile development or where material in upper parts of profiles has been eroded.

Particulate calcrete grains (calcrete pisolites; calcrete ooids, calcrete pellets) may undergo some transportation by downslope movement where the soils


are developed on slopes (Read, 1974). Unless soils become indurated, they are highly susceptible to erosion or reworking following marine transgression, resulting in their being mixed with skeletal material and incorporated into marine units; these reworked calcrete pisolites and calcrete ooids may be mistaken for marine cryptalgal oncolites and marine ooids. Laminar calcrete horizons tend to resist erosion and commonly remain as a hard caprock following removal of loose soil. With marine submergence, the laminar calcretes may be overlain by marine units containing reworked calcrete pisolites, calcrete ooids and pellets.

Calcrete pisolites, ooids, pellets

These are abundant constituents of many calcrete profiles (Swineford et al., 1958; James, 1972; Read, 1974). They occur in loose soils (calcrete-ooid grainstones, -pisolite grainstones and lithoclast breccias) in upper parts of calcrete profiles (Read, 1974) or in brecciated and leached “chalky” horizons having secondary “mud-support” fabrics and which are interlayered with laminar calcretes (James, 1972) (Fig. 1). Pisolitic soils and breccias commonly show upward coarsening or reverse grading (Dunham, 1969; Bernoulli and Wagner, 1971; Read, 1974). Also, calcrete pisolites commonly show in situ brecciation (Dunham, 1969).

Calcrete ooids (Plate I, A, B) are round, sand-size particles composed of an outer envelope of calcrete around a skeletal grain, lithoclast or detrital quartz nucleus (Siesser, 1973; Read, 1974). Most envelopes are faintly laminated and are composed of a few indistinct concentric layers of unoriented cryptocrystalline calcite and tangential calcite needle layers (James, 1972; Read, 1974). The envelopes are up to 0.05mm or more thick and they thin over sharp terminations of tabular nuclei. Some carbonate nuclei are altered tc‘ sparry or cryptocrystalline calcite.

Calcrete pisolites are spherical to ellipsoidal concretions commonly from 2mm to 5cm diameter (Plate I, C, D). The smaller pisolites are similar to calcrete ooids, but larger gravel-size pisolites consist of a thick (2 mm to 1 cm) outer envelope of laminated calcrete around a lithoclast, calcrete fragment or less commonly a skeletal grain nucleus (Plate I, C, D; Plate 11). Envelopes may be of uniform thickness or show “downward thickening” described by Bretz and Horberg (1949), Swineford et al. (1958) and Dunham (1969) or “upward thickening” of Berqoulli and Wagner (1971) (Plate 11, A, B). Fitted polygonal structure of adjacent pisolites may also be developed (Dunham, 1969). Lenses of host sediment between laminae are common, and may be perched in upper parts of pisolite envelopes where pisolites were not rotated during growth (Dunham, 1969). The laminated appearance of pisolite envelopes described by Read (1974) is due to alternation of thin light-colored layers (0.05mm) with brown layers (0.02-0.2 mm). Brown layers have indistinct boundaries, and are composed of cryptocrystalline calcite with disseminated pigment and

60 J.F. READ

PLATE I1

Calcrete pisolites A. Laminated calcrete coating upper surface of lithoclast; note resemblance to bun-shaped cryptalgal structure. Field approximately 8 em. Shark Bay., B. Laminated calcrete coating lithoclast of laminar calcrete. Field approximately 8 em. Shark Bay. C. Calcrete-coated mollusk valve (1 cm diameter). Shark Bay.


angular silt-size fragments of quartz whereas light-colored layers are composed of cryptocrystalline calcite, microcrystalline calcite and opaline silica and are free of pigment and detrital quartz; the layers commonly have sharp boundaries and matching borders and resemble concentric shrinkage cracks (Swine- ford et al., 1958) that have been filled by calcite cement. Calcrete pisolites described by James (1972) have dark laminae of tangential calcite needles whereas light-colored laminae are composed of cryptocrystalline calcite. Con- stituent grains within pisolite nuclei are commonly coated with calcrete or replaced by calcrete; remaining intergranular voids are filled by sparry calcite or calcrete. Pisolites described by Scholle and Kinsman (1974) have envelopes of cryptocrystalline aragonite and interlayered coarse radial-fibrous aragonite.

Calcrete pellets (0.03- 0.5 mm) of structureless cryptocrystalline calcite are dominant components in upper parts of calcrete profiles in Barbados; they form by alteration of skeletal grains in the calcrete profile (James, 1972). Similar pellets also occur in intergranular voids in calcreted marine limestone in Shark Bay (Read, 1974) and in calcreted limestones described by Estaban (1974). Pelleted textures are common in laminar calcretes of high-Mg calcite mineralogy (Scholle and Kinsman, 1974).

Laminar calcretes

Laminar calcretes (Fig. 1, Plate 111) commonly underlie soils that may be pisolitic (Swineford et al., 1958; Read, 1974) or organic-rich (Multer and Hoffmeister, 1968), or they may be interlayered with pisolitic, brecciated and leached horizons beneath leached soils (James, 1972)

Laminar calcretes are generally a few centimeters thick and composed of indurated sheet-like deposits of finely laminated calcrete in the form of flat-laminated structures and small laminated domes and saucers up to 30 cm or more in diameter (Plate 111) that are similar to cryptalgal structures. Laminar calcrete commonly follows irregularities in the underlying surface and tends to fill depressions (Multer and Hoffmeister, 1968); less commonly it truncates structures in the underlying material (Read, 1974) or may thicken over underlying highs (Bernoulli and Wagner, 1971, their fig. 6). Contacts between laminar calcrete and underlying sediment may be marked by rinds of red-brown unlaminated to poorly laminated calcrete; these rinds also surround rock fragments that are enclosed by laminar calcrete (Multer and Hoffmeister, 1968; Read, 1974). Sinkholes and’fissures are lined with laminar calcrete and may be filled with material from overlying soils. Laminar calcrete also floors large secondary (brecciation or solution) voids in sediments beneath the laminar horizon; void roofs are irregular’ with numerous pendant projections. The remaining void space may be filled with infiltrated soil or by later sparry calcite. Bedding surfaces in laminar calcrete are commonly brecciated (Swineford et al., 1958) or they may show sinuous root impressions (Read, 1974).

62 J.F. READ


Laminar calcrete is a dense to porous, red-brown to gray, finely laminated deposit. Laminar calcretes from Shark Bay (Read, 1974) have laminae (commonly 0.02 0.2 mm thick) with indistinct, finely undulating boundaries (Plate 111, D); laminae are texturally similar (little grain-size variation), the laminated appearance of the calcrete being mainly due to fine disseminated pigment concentrated in darker layers. Angular silt-size fragments of detrital quartz occur in the laminae (Plate 111, D) and lenses of carbonate grains and detrital quartz grains are commonly enclosed by adjacent laminae (Plate 111, B). Micro- unconformities formed by truncation of older laminae by younger layers also occur. Laminar calcretes in Florida (Multer and Hoffmeister, 1968) are of two, commonly interlayered types: (1) porous laminated crust, laminae 1- 5 mm, discontinuous, with poorly defined boundaries and composed of stained calcrete calcite, with common horizontal root molds, organic debris and lithoclasts; and (2) dense-laminated crust, persistent laminae 0.5- 1.0 mm with sharp boundaries; laminae consist predominantly of stained calcrete; other constituents are minor. Laminar calcretes in Barbados (James, 1972) consist of alternating light and dark bands of microcrystalline calcite (0.1-0.15 mm thick) and thinner laminae (0.06 mm) of flat-lying calcite needles oriented parallel to laminae. Laminar calcretes of high-Mg calcite in the Persib Gulf are composed of pelleted cryptocrystalline high-Mg calcite. Lamination is poorly defined in thin section and is due primarily to color-banding (Scholle and Kinsman, 1974).

Horizons underlying laminar calcretes

In calcrete profiles in Shark Bay, laminar calcretes may overlie massive calcretes that grade down into mottled calcrete horizons overlying country rock (Fig. 1); more commonly they directly overlie mottled calcrete horizons. The massive calcrete horizon (rarely up to 3 m thick) is characterized by dense structureless calcrete with secondary “mud-support” fabric in which calcrete coats, displaces and replaces carbonate grains and fills intergranular voids. Fine tubular molds of root hairs (0.2 mm in diameter) which cut through grains and cements are locally common. The mottled horizon is characterized by calcrete mottles that are irregular in shape (and 1-5 cm wide), due to localized distribution of calcrete in the host sediment. Mottles

PLATE I11

Laminar calcretes A. Oblique view of sawed laminar calcrete showing similarity to cryptalgal laterally linked hemispheroids. Florida. B. Polished slab of flat-laminated calcrete with small-scale undulations. Small lens of included sediment in lower left. Florida. C. Polished slab of laminar calcrete from domal portion of undulating calcrete bed. Florida. D. Thin section, laminar calcrete. Shark Bay.

64 J.F. READ

are patches where calcrete replaces grains and fills intergranular voids to form a secondary “mud-support” texture; void-filling sparry calcite is confined to root casts. The calcrete mottles grade into patches with grain-support texture in which calcrete coats grains and only partly fills intergranular voids.

Laminar calcretes of Barbados (James, 1972) (Fig. 1) are interlayered with and grade down into chalky carbonate horizons consisting of altered host carbonates that are intensely brecciated, fractured and leached. These chalky horizons contain coated grains and pellets and have secondary “mudstone” fabrics with pelleted textures or recrystallized (microspar) fabrics; grains are commonly separated from enclosing matrix by fracturing along grain boundaries. Host rocks underlying laminar calcretes in Florida (Fig. 1) may show evidence of leaching, but Multer and Hoffmeister (1968) do not mention any calcrete effects. Bernoulli and Wagner (1971) describe vadose diagenetic effects that include dripstone cements, separate generations of internal sediment, and fracturing post-dating cementation in host sediments that underlie laminar and pisolitic calcrete horizons in Lower Jurassic rocks, Italy. Amsbury (1967) notes microcrystalline cements and vadose internal sediments in caliche profiles in Lower Cretaceous rocks of Texas.

Associated features

Several types of macro- and microstructures are common to both calcrete deposits and cryptalgal deposits.

Tepee structures or “caliche (calcrete) anticlines” vary from broad gentle swellings to steep-sided structures (1.5- 17 m in diameter) and are cored with soft powdery calcrete; they have been described from calcretes in Texas (Blank and Tynes, 1965). Grossly similar structures have been described in Recent tidal deposits (Kendall and Skipwith, 1968; Davies, 1970) and from subtidal, cemented sediments (Shinn, 1969).

Fenestral fabrics are commonly associated with cryptalgal fabrics ( Aitken, 1967; Logan et al., 1974), but may also be developed in soils and calcretes. Fenestral fabrics in soils are described in Brewer (1964). Amsbury (1967) notes fenestral fabrics in calcrete profiles in the Cretaceous of Texas; tubular fenestrae (root molds) have been noted in many Recent calcrete deposits (Multer and Hoffmeister, 1968; James, 1972; Read, 1974).

Microscopic tubular structures ranging from 5 to 20 pm, that are similar to filament molds of blue-green algae, are abundant in some calcrete deposits (James, 1972; Ward, 1970). Some tubules penetrate grains, and others form an interlocking meshwork between grains. The tubules have been interpreted as calcified filaments of blue-green algae (James, 1972) or the root hairs of angiosperm dune plants (Ward, 1970). Estaban (1974) described problematic organic (?) structures termed Microcodium from Eocene calcretes, Spain. The structures consist of spherical, elliptical, sheet- and bell-like clusters of petal- like calcite prisms up to 1 mm long and of polygonal cross-section. Locally, these structures are the main constituent of the calcretes.

ABIOGENIC STROMATOLITE-LIKE STRUCTURES

ORIGIN OF CALCRETES

65

Chemical and biological aspects

Calcrete formation results from processes in the vadose zone involving solution and reprecipitation of carbonate. The chemistry of the system CO, - H,O-CaCO, is discussed in Bathurst (1971) and Sweeting (1972) and the reactions may be summarized by:

CO, + H 2 0 + CaC0, 2 Ca2+ + 2HCO; Dissolution of CaCO, (reaction proceeds t o right) may be caused by increased CO, partial pressure, decreased temperature or low pH; precipitation (reaction moves to left) may be caused by decrease in CO, , evaporation or by biological activities. Rapid changes in composition of fluids favoring alternating periods of solution and reprecipitation are common in the vadose zone during periods of rain that are followed by long periods of evaporation (James, 1972).

CO, partial pressure is probably the major factor influencing solution of carbonate, which is only slightly soluble in pure H,O. Sweeting (1972) notes that average rain has a pH of 5.0-5.5 with a tendency for more acid rain- waters in coastal areas (due to sea spray and addition of SO',- and C1- ions), but considers that the role of atmospheric CO, in rainwater may be overrated, as C02 partial pressure in the soil zone is commonly many times greater than in the atmosphere or in the underlying bedrock (Bathurst, 1971; Sweeting, 1972). Consequently, the most likely place for marked increase in CO, partial pressure is in the soil. High CO, partial pressures in soils may result from activities of higher plants (particularly their roots) and micro-organisms. Carbonate dissolution may also be promoted by organic acids formed by plant growth and decay; for example, root hairs produce chelating acids which dissolve CaC03 (Sweeting, 1972). Solution may be aided by rapid percolation of meteoric water during heavy rains (an increase in unsaturated solvent flow past a solid increases rate of dissolution); rapid percolation would be common in sands but would tend to be inhibited in fine-grained sediments. Warm temperatures in tropical soils prior to rainfall also may accelerate the rate of solution. Finally, unstable mineralogies (e.g. abundant aragonite) of host sediments would make them more susceptible to solution.

Precipitation of carbonate may result from evaporation, removal of CO, , temperature changes, pH changes or biologic activities. Evapotranspiration causing rapid loss of vadose water, is probably an important cause of precipitation; James (1972) points out that the small crystal size of calcrete is probably related to precipitation from rapidly evaporating solutions. Evapo- transpiration may also bring about precipitation by causing CO, loss due to decrease in pressure of pore water as a result of increase in suction pressure (Netterberg, 1971 in Siesser, 1973). Increased temperature may also cause

66 J.F. READ

C 0 2 loss. Carbonate precipitation may be promoted by pH increase downward in the soil profile. Multer and Hoffmeister (1968) note minimum pH of 5.6 in soils increasing to a maximum of 8.2 in underlying host sediments. Precipitation may also be promoted by microfloras associated with calcretes; Krumbein (1968) reported that cultured microfloras were able to cause precipitation of large amounts of CaC03, precipitation being due not only to pH change, but also possibly related to transfer of Ca2+ due ‘to chelating substances generated by the microflora. Blue-green algae, fungi or root hairs of dune plants may be involved with calcite precipitation in calcretes in Barbados and Yucatan (Ward, 1970; James, 1972).

The chemistry of vadose waters has an important influence on the carbonate phases and morphology of crystals precipitated (Reeves, 1970; James, 1972). M.ost vadose waters involved in calcrete formation are “fresh”, containing low concentrations of dissolved salts; consequently, calcite is the common calcrete mineral formed. Crystal form of the precipitated carbonate is strongly influenced by dissolved ions and by intense evaporation which may cause a high degree of supersaturation and formation of abnormal crystal forms (needle and whisker crystals; James, 1972). In contrast to “fresh” vadose waters and associated calcite mineralogies, vadose waters in some arid coastal areas are hypersaline and the calcretes are composed of aragonite and high-Mg calcite. These unstable calcrete minerals are probably related to high concentrations of dissolved ions including Mg2+ and Sr2+( Scholle and Kinsman, 1974).

Profile development

Calcrete profile development is discussed in Gile et al. (1966). However, as calcrete profiles are highly variable, profile development can only be discussed in general terms. The calcrete horizons are characteristic of distinctive stages of calcrete accumulation and reflect increasing deposition of diagenetic carbonate. Carbonate is leached from upper parts of soil profiles and re- precipitated in place or carried down the profile. Initially, the precipitated carbonate forms coatings around grains to form calcrete ooids and pisolites, the coatings tending to push the grains apart (Siesser, 1973). Lower in the profile, the calcrete coats and replaces grains as well as plugging interparticle voids to form local patches of “mud-supported” carbonate. In the zone of most frequent wetting by unsaturated flow, void space in the sediment is obliterated and grains are replaced by carbonate to form extensive secondary “mud-support” fabrics (massive carbonate horizons; Read, 1974). The massive calcrete horizon, being relatively impermeable, inhibits percolation and a thin zone of free‘ water accumulates periodically above the horizon. At this stage laminar calcrete develops. Laminar calcrete may also form directly on bedrock surfaced by thin cryptocrystalline calcite films, believed by Multer and Hoffmeister (1968) to form during subaerial exposure of bedrock. Laminar calcrete may also be precipitated on walls of horizontal to subvertical fissures and irregular cavities.


Most coastal calcrete profiles are rarely more than a few meters thick. Thick inland profiles tens of meters thick probably formed by eolian aggra- dation of the profile (Brown, 1956). Loose, carbonate-rich oolitic and pisolitic soils in Shark Bay probably reflect low rainfall (20-30 cm/year) whereas leached carbonate-poor, organic-rich soils of Florida and Barbados relate to higher rainfall (110-150 cm/year). Oriented growth of pisolites indicates in place deposition of coatings without rotation of grains, the sense of the orientation (upward or downward thickening) possibly being related to climatic conditions; for example, downward thickening probably results from collection and evaporation of water films on undersides of pisolites; absence of downward growth in Shark Bay pisolites is probably due to low rainfall and high evaporation, inhibiting localization of water films on undersides of pisolites, as well as to rotation of pisolites by soil creep. Upward coarsening of pisolites may relate to more rapid deposition of carbonate in upper parts of profiles or may be due to increase in supply of coarse lithoclastic nuclei following stripping of loose soils from local highs (Read, 1974).

INTIMATE ASSOCIATION OF CALCRETES AND CRYPTALGAL STRUCTURES

Tidal environments are transitional to landward into terrestrial environments dominated by subaerial weathering processes. Calcrete ooids and pisolites from soils may be carried by sheet flooding or winds into tidal flats where they may be bound by algal mats and incorporated into stromatolitic sediments (Kendall, 1969). Also, depending on relative movements of sea level, calcretes may overlie, underlie or be complexly interlayered with cryptalgal sediments (Multer and Hoffmeister, 1968). With regressive offlap or progradation, cryptalgal sediments will tend to be overlain by calcretes (Col- acicchi et al., 1975) particularly where this is accompanied by lowered sea level. Alternatively, with marine submergence, cryptalgal structures may be developed on laminar and pisolitic calcretes formed on underlying beds during emergence; such a situation occurs in Shark Bay, Western Australia, where columnar stromatolites are developed on laminar calcretes overlying Pleisto- cene sediments. Finally, alternation between shallow-marine/tidal-flat conditions and periods of subaerial emergence may result in complexly interlayered cryptalgal sediments and calcretes; for example, in Lower Jurassic rocks in Italy, subaerial horizons of calcrete pisolites and laminar calcretes are interlayered with marine units containing cryptalgal oncolites (Bernoulli and Wagner, 1971).

DISTINGUISHING CALCRETES AND CRYPTALGAL SEDIMENTS

Laminar and pisolitic calcretes are similar to stratiform stromatolites, laterally linked hemispheroids and oncolites (Logan et al., 1964). I t is difficult

68 J.F. READ

to define a set of distinguishing criteria, partly because of the great variation in weathering profiles, structures and fabrics of calcrete deposits and considerable variation in cryptalgal structures and fabrics. However, the fact that ancient calcretes have been recognized as such indicates that distinctive criteria do exist.

General features

Certain diagnostic features are common to both laminar and pisolitic calcretes. These include:

( 1) Calcretes are developed within distinctive weathering profiles that may be characterized by certain vertical successions of calcrete structures and fabrics, and by the presence of vadose diagenetic effects (leach fabrics, cavity formation, internal sediments of crystal- and pellet-silt, and floored voids or Stromatactis structures). Also, calcrete deposits (and subjacent sediments) commonly show abundant in situ brecciation effects that include fractured pisolitic and laminar calcretes, and brecciated, jointed and fractured host sediments in which vertical and horizontal fissures are lined or floored with laminar calcretes, pisolitic or oolitic soils or spar. (2) Lithologies of calcrete deposits are dominated by calcrete ooids, pisolites, and pellets, laminar calcretes, angular lithoclast breccias, and secondary “mud-support’’ fabrics due to calcrete deposition. These should be distinguishable from typical sediments associated with cryptalgal deposits (marine oolitic, pelletal and intraclastic carbonates, flat-pebble breccias and tidal-channel deposits; and structures such as fenestral fabrics, cross-bedding and ripple marks). (3) Calcrete deposits tend to follow microtopography, dipping into steep-walled depressions and sinkholes. (4) Calcrete deposits commonly contain micro-unconformities or solution unconformities marked by sinkholes and collapse structures. These solution surfaces occur on exposed calcretes or underlie laminar calcretes, truncating structures and fabrics of underlying indurated calcrete deposits (pisolitic/oolitic sediments, breccias, laminar calcretes). With renewed calcrete deposition following burial by aggrading soils or prograding dunes, they may be incorporated into calcrete deposits. Micro-unconformities are also developed in cryptalgal deposits, but these may show evidence of mechanical erosion. (5) Lamination in many calcretes is largely due to color differences (Read, 1974), although some may reflect differences in crystal size and morphology of the precipitated carbonate (James, 1972); also, calcrete laminae commonly have finely undulating, indistinct boundaries. Cryptalgal laminae may show size differences in the bound particulate carbonate. However, particulate grains in cryptalgal laminates commonly alter to, and are cemented by, cryptocrystalline aragonite, which tends to destroy the particulate texture (Logan, 1974). (6) Calcretes of relatively young geological ages may contain traces of terrestrial organisms (land snails, insect pupal cases, and impressions of larger plant roots lined with laminar calcrete). (7) Thickness of the deposits may also give a clue to origin. Most cryptalgal laminates of substantial thickness develop in settings where submergence is balanced by shoaling due to sedimentation. Where this balance is maintained for long periods, thick cryptalgal sequences may develop. Although thick calcrete profiles may develop under exceptional conditions (such as continued eolian supply of carbonate), individual laminated or pisolitic horizons are rarely more than a few centimeters or meters in thickness.


Calcrete pisolites and cryptalgal oncolites

Calcrete pisolites resemble cryptalgal oncolites (Logan et al., 1964), ranging in form from concentrically stacked spheroids (analogous to concentrically laminated calcrete pisolites of Read, 1974) to inverted stacked hemispheroids (analogous to downward or upward elongated pisolites of Dunham (1969) or Bernoulli and Wagner (1971)). Features that are characteristic of calcrete pisolites (Swineford et al., 1958; Dunham, 1969; Wardlaw and Reinson, 1971) include: (1) Calcrete pisolites generally lack mechanical sedimentary structures, but may be reverse graded (upward coarsening), nuclei are mainly lithoclasts or pisolite fragments, and associated grains are calcrete-coated; in contrast, cryptalgal pisolites (oncolites) commonly are associated with sedimentary structures such as cut-and-fill and cross-bedding characteristic of current-laid deposits, are associated with numerous uncoated grains, and many have skeletal nuclei. (2) In-place downward growth is characteristic of some calcrete pisoliteo, and may result in polygonal-fitted structure and perched inclusions; however, other pisolites may show upward elongation (Bernoulli and Wagner, 1971) or concentric growth (Read, 1974). Also, fitted structure of calcrete pisolites may be simulated by fibrous-calcite overgrowths on marine cryptalgal oncolites, as suggested by Kendall(l969). (3) Calcrete pisolites commonly show repeated in situ brecciation during growth, in which the calcrete coating the fracture is continuous with the outer pisolite coatings; matching halves of a pisolite may show little movement from their original position. Fracturing is commonly associated with episodes of leaching, solution-cavity formation, cementation and internal sedimentation.

Laminar calcretes, cryptalgal structures and laminates

Laminar calcretes simulate: (1) flat-laminated cryptalgal sheets where the laminar-calcrete surface is planar; (2) laterally linked hemispheroids where the calcrete surface is developed as low domes and saucers; and (3) columnar- stacked hemispheroids where the surface is developed into domes and intervening small sinkholes (Logan et al., 1964). Also, laminar calcrete that fills depressions may simulate stacked saucers (Kendall and ' Skipwith, 1968). Aitken (1967) summarized criteria for recognition of cryptalgal origin and Multer and Hoffmeister (1968) outlined some criteria for distinguishing between cryptalgal sediments and laminar calcretes. These and other criteria are critically discussed below: (1) Cryptalgal features tend to thicken over highs. In contrast, most laminar calcretes tend to fill depressions, although some laminar calcretes may also thicken over highs (Bernoulli and Wagner, 1971, their fig. 6). Most laminar calcretes tend to follow closely underlying microtopography, bending into collapse structures, sinkholes, lining vertical walls of fissures and flooring voids in underlying sediments to form Strornatactis structures. (2) Grain-size alternation in cryptalgal laminates where it is not destroyed during early diagenesis (Logan, 1974), contrasts with the color-lamination in many laminar calcretes. Also, cryptalgal laminates may be finely graded (upward fining) (Shinn et al., 1969), contain thin cross-stratified or rippled layers, consist of alternating layers of dolomite and calcite (Gebeleinand Hoffman, 197 3), be interlayered with thin bituminous films (Laporte, 1967), or.show evidence of convolution of flexible mat layers (Davies, 1970b).

70 J.F. READ

(3) The presence of abundant algae in mats does not help in distinguishing cryptalgal sediments, as algal filaments are rarely preserved (Logan et al., 1964). Also, the occurrence of abundant tubules resembling algal filaments in calcretes from Barbados and elsewhere (James, 1972) suggests that traces of algae, where preserved, should be used with caution in interpreting laminated carbonates as cryptalgal. However, certain arrangements of filament casts may be characteristic of some cryptalgal deposits (e.g. palisade structure: Davies, 1970b). (4) Presence of intense bioturbation effects, considered to be characteristic of Florida algal mat terrains by Multer and Hoffmeister (1968), is not characteristic of many ancient cryptalgal deposits. Well-preserved cryptalgal lamination in sediments actually reflects lack of burrowing; intense bioturbation destroys lamination and masks the algal-mat origin (Logan et al., 1974). However, scattered marine sediment-filled burrows in cryptalgal sediments may give a clue to algal origin. In contrast, laminar calcretes that underwent marine submergence are likely to contain organic borings rather than burrowings, reflecting their indurated character. (5) Distinctive fenestral fabrics are commonly associated with cryptalgal sediments, and may occur in characteristic vertical sequences (Logan et al., 1974). Fenestral fabrics also may occur in soils (Brewer, 1964). calcretes (Amsbury, 1967) and travertines (Tebutt et al., 1965), but with care it should be possible to distinguish these from cryptalgal fenestral fabrics that have distinctive form and vertical successions. (6) Whereas laminar calcretes are likely to be associated with weathering profiles and vadose diagenetic phenomena, cryptalgal laminates may be associated with distinctive domal, digitate, columnar or polygonal stromatolitic structures and their associated marine sediments.

SUMMARY

Calcretes are typically associated with distinctive weathering profiles and associated vadose diagenetic phenomena, and are characterized by distinctive calcrete products. The calcrete deposits are strongly controlled by underlying topography, and may contain solution surfaces and traces of terrestrial organisms. These features contrast with the typical lithologic associates of cryptalgal deposits.

Pisolitic calcretes differ from cryptalgal oncolites in that they lack mechanical sedimentary structures and show evidence of in situ growth that may be interrupted by periods of in situ brecciation, leaching, cementation and internal sedimentation. Laminar calcretes may simulate a wide variety of cryptalgal deposits (laterally linked hemispheroids, stacked hemispheroids, flat-lying sheets). However, in contrast to cryptalgal deposits, laminar calcretes tend to fill lows, floor voids, line walls of fissures, and they lack distinctive cryptalgal desiccation features. Calcrete laminations also lack characteristic features of cryptalgal laminations. Finally, features such as fenestral fabrics, tepee structures and traces of algae occur in both cryptalgal and calcrete deposits, and should be used with caution in identifying these deposits.


ACKNOWLEDGEMENTS

Thanks are due to Brian W. Logan for supervision of the study on the Shark Bay calcretes during 1967- 1970 and for providing several photographs of calcrete. Helpful discussions on calcrete were had with various people including R.J. Dunham, N.P. James, R. Colacicchi, and M. Estaban. I wish to thank J.F. Truswell, Malcolm' Walter, D.A. Hewitt and G. Grover for critically reading the manuscript and providing helpful suggestions and L. Groover for typing.


Chapter 3.2

SPELEOTHEMS

John Thrailkill

INTRODUCTION

A variety of crystalline cave deposits (speleothems) resemble stromatolites and related forms in their gross morphology and internal structure. Such common forms as stalagmites (Fig. 1) will often have an internal structure that is at least superficially identical to described stromatolites. Stalactites seen in thin section (Fig. 2) have a regularity of shape that may suggest a biogenic origin.

It is clearly impossible in limited space to describe all speleothems that may mimic or contribute to the understanding of stromatolites or other biogenic deposits, and this discussion will center on the two forms that are the most important in this regard: cave popcorn and unattached speleothems. Cave popcorn is also referred to as cave coral, but the latter term seems to be more misleading to the uninitiated and should be avoided. Because of its probable biogenic nature and relationship to the above speleothems, moonmilk will be discussed briefly. The term moonmilk, although a mistranslation from the German, is widely used in the U.S. and Britain to apply to any soft subaerial speleothem, although much of what is called moonmilk is actually a residue from the destruction of other speleothems, rather than a primary deposit.

Because of the implication in the placement of this section that speleothems are of abiogenic origin, an examination of this question would seem to be in order at the outset. A number of studies have indicated that a varied microflora, including blue-green and red algae, is present in caves (e.g., Claus, 1955; Palik, 1960; Nagy, 1965; Jones, 1965, gives a good review). Although sample location descriptions are often sketchy, it seems likely that algae that are normally photosynthetic do live in complete darkness, and a role of such organisms in speleothem deposition cannot be ruled out. Friedman (1955, 1964) has found filaments of two species of blue-green algae in a cave in Israel to be encrusted with calcium carbonate.

74 J. THRAILKILL

Fig. 1. Polished sections of stalagmites from (left) Kentucky (collected by D.P. Beiter) and (right) Florida, U.S.A.

Fig. 2. Thin section (cross polarized light) normal to axis of calcite stalactite from Colorado, U.S.A. Width of photograph is 3.5 mm.


BIOGENIC ASPECTS OF SPELEOTHEM DEPOSITION

It is usually assumed that the principal processes responsible for the deposition of the better-known speleothems (stalactites, stalagmites, flowstone,. and related forms) are abiogenic. The apparent absence uf microorganisms on growing surfaces and the lack of included organic material seem to support this view. Fig. 2 is a photomicrograph of a section normal to the length of a small calcite stalactite in which no included organic material can be seen (there are a few opaque detrital grains). The faint banding near the outer margin and the outline of the central tube is the result of ghost rhombohedral faces. Although the portion of this stalactite outside the central tube has recrystallized, it seems unlikely that such recrystallization could have so effectively excluded organic matter (and left the detrital grains). The author has examined numerous other stalactites and stalagmites which apparently contain no organic matter.

At the other extreme, it is nearly certain that some carbonate speleothems are of biogenic origin. The strongest case has been made for soft masses which coat the walls or other speleothems in many caves. These deposits, called moonmilk, have a pasty texture when wet and are powdery when dry. When the original deposit is undisturbed it often has a nodular outer surface, similar to some forms of cave popcorn. A variety of carbonate minerals has been reported from moonmilk and moonmilk-like material (Warwick, 1962; Moore and Nicholas, 1964; Thrailkill, 1971). In thin section, moonmilk is seen to consist largely of very fine carbonate crystals in finely laminated sheets and as loose masses partially filling open pores (Fig. 3). Examination of various occurrences has revealed the presence of fungi (Ulrich, 1938), bacteria, algae, and protozoa (Hoeg, 1946; Caumartin and Renault, 1958; Williams, 1966). Williams (1959) isolated a bacterium from moonmilk which deposited calcium carbonate in a laboratory experiment.

By its very nature, moonmilk is unlikely to be preserved and, except for the botryoidal appearance of its outer surface in some occurrences, it bears little resemblance to stromatolites. In Carlsbad Caverns, however, there often appears to be a gradation between moonmilk and cave popcorn. Commonly, the stalagmite deposited at the point of impact of a ceiling drip is surrounded by cave popcorn which grades to moonmilk with distance from the drip. Moonmilk was found growing on cave. popcorn (Fig. 5) and some dry cave popcorn may actually be indurated moonmilk (Thrailkill, 1971).

,

CAVE POPCORN

This form, while seldom the most spectacular, is one of the more common speleothems. It consists of nodules, usually in clusters, developed on bedrock or on other speleothems. The clusters are oriented approximately normal to

76 J. THRAILKILL

Fig. 3. Thin section (cross polarized light) of calcite moonmilk from Colorado, U.S.A. Width of photograph is 3.5 mm.

Fig. 4. Cave popcorn growing on stalagmite in Carlsbad Caverns, New Mexico, U.S.A. Pencil at bottom is 130 mm long.


the surface to which they are attached, and although they may depart somewhat from this orientation, they seldom are elongated downward (as a stalactite) or upward (as a stalagmite). Fig. 4 shows a group of cave popcorn clusters on a large stalagmite in which the growth seems to be toward the left side of the photograph and slightly upward. This photograph also illustrates the common phenomenon of cave-popcorn clusters tending to form on one side of a stalactite or stalagmite, or on other surfaces. In some cases, a group of stalagmites will all have cave popcorn developed on the same side.

As the term is used here, all cave popcorn is deposited subaerially, i.e., from a thin film of water in an air-filled'cave. Some subaqueous speleothems, such as those lining the walls of deep pools, superficially resemble cave popcorn but are usually coarsely crystalline and lack the branching nodule structure and fine laminae. I t should be pointed out that it has been thought by some that the lack of geotropic orientation in cave popcorn indicated a subaqueous origin, but even a casual inspection indicates that it is now being actively deposited in areas that have not been flooded.

There is a wide variation in the size of individual nodules, both between localities and, to a lesser extent, within a single cluster (Fig. 5 ) . The largest nodules (which are clearly not stalagmites) have diameters of about 100 mm and the smallest on the order of 2 or 3 mm. There is likewise great variation in nodule shape, from nearly perfectly spherical to contorted masses. The morphology shown in Fig. 6 is quite common, but many individual nodules are more elongate (Fig. 7) and some are cylindrical (Fig. 9).

The mineralogy of cave popcorn has been little studied. Most of it consists of calcite and/or aragonite, but hydromagnesite and dolomite occur in cave popcorn from Carlsbad Caverns, New Mexico (Thrailkill, 1968, 1971) as does chalcedonic quartz. Huntite and magnesite have been reported from moonmilk (Pobeguin, 1960) and probably also occur in cave popcorn. In addition, various detrital minerals (quartz, clay, etc.) are commonly present in small amounts.

Cave popcorn usually consists of irregularly alternating clear and dark layers of crystalline aragonite or calcite (Figs. 7- 12). Even in highly ellipsoidal or cylindrical forms the layers are commonly nearly hemispherical, with the elongation of the nodule resulting from rapid thinning of the layers away from the axis (Fig. 10). In a few specimens the layering departs significantly from the hemispherical shape (Fig. 11). Reentrants between nodules may have acute angles and the layering may be partially discontinuous between nodules (Fig. 10) and curved concavely (Fig. 11). The former probably represent the joining of two nearby nodules by lateral deposition.

The growth of cave popcorn begins often from a smooth layer of flowstone or the surface of a stalactite (Fig. 12) or other speleothem. Fig. 1 2 shows another common phenomenon, early small nodules being covered with smooth flowstone, then later the development of cave popcorn. I t should be noted that all cave popcorn examined appears to grow by the addition of surface

78 J. THRAILKILL

Fig. 5. Cave popcorn growing on flowstone and stalagmites on wall in Carlsbad Caverns, New Mexico, U.S.A. Finest material is moonmilk. Pencil on right is 130 mm long.

Fig. 6. Cave popcorn from Carlsbad Caverns, New Mexico, U.S.A. Mineralogy is calcite and aragonite with small amounts of dolomite.


Fig. 7 . Polished section of cave popcorn from Colorado, U.S.A.

Fig. 8. Thin section (plane polarized light) of cave-popcorn nodule growing o n flowstone (Kentucky, U.S.A.). Width of photograph is 3.5 mm.

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Fig. 9. Polished section of cave popcorn from Carlsbad Caverns, New Mexico, U.S.A.

Fig. 10. Thin section (plane polarized light) of specimen in Fig. 9. Width of photograph is 3.5 mm.


Fig. 11. Thin section (plane polarized light) of inner layers of specimen in Fig. 7 . Width of photograph is 3.5 mm.

Fig. 12. Thin section (cross polarized light) of cave popcorn (left) growing on wall of stalactite (right) from Kentucky, U.S.A. (collected by D.P. Beiter). Width of photograph is 3.5 mm.

82 J. THRAILKILL

layers; there is seldom any evidence of disruption of the structure by substantial volume increases below the surface.

At present, it is not possible to describe unequivocally the processes responsible for the origin of cave popcorn. As stated earlier, it is clear that it is formed subaerially and it seems equally clear that it forms on surfaces which are covered only with a thin film of water. Its non-geotropic growth pattern is probably due to the fact that only a €ilm of water is present. Stalactites and similar speleothems grow downward because of rapid deposition in the pendant drop; when there is only a film of water growth takes place in directions determined by other factors. I t is not uncommon to see stalactites (carrying a pendant drop) growing from a clump of cave popcorn nodules in areas where the amount of water supplied to the surface has increased.

It is immaterial how the water film is supplied to the surface. Moore and Nicholas (1964) believe that the water seeps out from the surface behind the cave popcorn, but the writer’s observations are that the water may flow down from above, be pulled up from below by surface tension, seep out from behind, or be splashed on to the surface from an adjacent drip.

Seepage water entering a cave is normally (but not invariably) in equilibrium with a partial pressure of carbon dioxide (Pco,) considerably higher than that of the cave atmosphere and has such a high concentration of calcium, bicarbonate, and possibly magnesium ions that it will become supersaturated with respect to aragonite, calcite, and possibly dolomite and huntite if the Pco, falls to that of the cave atmosphere or if a significant amount of water evaporation occurs (Thrailkill, 1968, 1970, 1971). Deposition as stalactites and stalagmites occurs where such seepage is relatively concentrated and sufficient evaporation or carbondioxide loss occurs to cause precipitation. Because such precipitation is relatively rapid and the precipitation kinetics for the magnesium-containing carbonates are slow, usually only aragonite and calcite are precipitated in stalactites and stalagmites, with the former more likely when the Mg/Ca ratio is high (Murray, 1954; Thrailkill, 1968, 1971).

Because of the thinness of the water film on cave popcorn, accurate determination of Pco, or saturation with respect to carbonates have not yet been achieved. I t seems likely, however, that in Carlsbad Caverns at least, such water has a Pco, near that of the cave atmosphere and is very nearly saturated with respect to the carbonates. I t has been proposed, therefore, that the deposition of cave popcorn is tied to carbon dioxide and/or water vapor escape from the thin film, with fastest deposition occurring where such escape is rapid (Thrailkill, 1965, 1968). Thus deposition will tend to occur where the film is thin, as on nodule tips where the convex curvature is great, and be inhibited where the film is thick, as in reentrants between nodules and on the lower surfaces (where gravitational thickening occurs). Slight irregularities on a surface would produce a local thinning of the film and the budding of a new nodule. Other factors are ventilation, which would also favor deposition


on nodule tips and inhibit it in reentrants, and water supply (i.e., nodules sometimes seem to grow toward a drip), but since too much water thickens the film (or causes pendant drops to form), this is of doubtful importance. Because aragonite and calcite deposition will increase the Mg/Ca ratio (Thrail- kill, 1968), evaporation will increase the Mg concentration, and the water may be in contact with the cave popcorn for a long time, Mg-containing minerals may precipitate.

As discussed earlier, algae, bacteria, and other microorganisms have been found inhabiting many cave surfaces. If the above explanation for cave popcorn deposition has any validity, it is obvious that the presence of a microbiota in the water film which utilizes earbon dioxide may have a profound effect on carbonate deposition. That such may be the case is suggested by the fine dark laminae in much cave popcorn (e.g. Figs. 7-12) which, may be organic-rich layers. This might also explain the abundance of hydromagnesite in cave popcorn from Carlsbad Caverns. According to current thermodynamic data, seepage water could reach saturation with respect to this mineral only at impossibly high Mg/Ca ratios (- 10’) or at Pco, values several orders of magnitude bslow that of the cave atmosphere. I t has been suggested by G.W. Moore (personal communication, 1972) that its precipitation is caused by C02 depletion by microorganisms. I t should be emphasized, however, that much cave popcorn (e.g. Figs. 8, 9) contains thick layers of clear calcite or aragonite with no indication of organic involvement.

UNATTACHED SPELEOTHEMS

Cave pearls are rounded unattached speleothems which are occasionally found in caves. They have drawn attention because of their attractiveness and the analogy of their supposed origin with that of pearls. They often occur in small basins under active ceiling drips, and those that have been sectioned usually are found to consist of layers of calcite or aragonite surrounding a nucleus of some type (Fig. 13). The term cave pearl is usually reserved for specimens that are approximately spherical and whose outer surface is smooth enough to be reflective when wet.

There is a second type of unattached speleothem, termed a pool accretion (Thrailkill, 1963) which, although more common than cave pearls, has attracted much less attention. It is either crudely spherical or ellipsoidal (Fig. 14) and has a rough outer surface. In thin or polished section, pool accretions are porous and have a crude concentric layering (Fig. 15). Fig. 16 shows two other features common to most pool accretions: scattered calcite crystals tending to be arranged radially or dendritically, and wavy laminae.

Several cave pearls which have been sectioned by the writer have the internal structure of a pool accretion, with a few thin laminae on the outside. In the unattached forms examined by the writer it is the presence of these laminae

84 J. THRAILKILL

Fig. 13. Thin section (cross polarized light) of cave pearl from Colorado, U.S.A. Clear calcite layer surrounding microcline crystal nucleus (upper left) is covered by finely laminated outer layers. Width of photograph 3.5 mm.

Fig. 14. Sectioned pool accretion from Kentucky, U.S.A. (collected by G. O’Dell).


Fig. 15. Polished section of specimen in Fig. 14.

Fig. 16. Thin section (cross polarized light) of pool accretion from Colorado, U.S.A. Width of photograph is 3.5 mm.

86 J. THRAILKILL

(usually alternating dark and clear) that gives some of them the “polished” appearance of cave pearls, although many cave pearls commonly contain layers of radial calcite crystals more than 1 mm thick (Fig. 13).

It is commonly supposed that cave pearls are unattached because they are rotated by*falling water, but it is clear that this is not the case in numerous occurrences (Thrailkill, 1963; Moore and Nicholas, 1964). The smooth surface layers of most cave pearls are so similar in appearance to those’of cave popcorn (compare Figs. 10 and 13) that it seems likely that they may both be deposited under similar conditions. I t is also unlikely that pool accretions are unattached because of agitation. The ellipsoidal specimen in Figs. 14 and 15 is 9 cm in diameter but the layering is nearly concentric with the center. De- position must have occurred simultaneously on the bottom and top. All of the pool accretions known to the writer are found either in shallow pools or in positions*where they could have been washed from these pools, and it is likely that they are formed when in contact with standing water. The mechanism responsible for their formation is unknown, but the resemblance of their internal structure (Fig. 16) to that of moonmilk (Fig. 3) suggests some biogenic processes are involved.

CONCLUSIONS

A comparison of these speleothems with illustrations in papers discussing stromatolites reveals many similarities. Some examples are Fig. 9 (this paper) with figure 31 in Walter (1972a) or figure 20 in Hofmann (1969a); Fig. 12 (this paper) with plate 7-1 in Monty (1967); and Fig. 15 (this paper) with plate 2F in Logan et al. (1964).

Some tentative criteria which may be used to distinguish speleothems from stromatolites are: (1) association with other speleothems such as stalactites; (2) scale, since it is doubtful that cave popcorn nodules attain diameters greater than about 100 mm (although stalagmites may be much larger); (3) presence of nodules or layers in nodules which are apparently abiogenic, but not due to much later recrystallization (although these may be found in stromatolites as well); (4) presence of initially deposited or early diagenetic magnesium minerals such as dolomite, huntite, or hydromagnesite; ( 5 ) indication that most of the deposition was due to precipitation from solution rather than entrapment of detritus; (6) non-vertical orientation; and (7) lack of filaments or other biogenic structures. Criterion 1 may be useful for distinguishing pool accretions from unattached stromatolites (oncolites), but because it is distinctly possible that such pool accretions are stromatolites deposited in a cave environment, other criteria which might be suggested are likely to stem largely from our ignorance of the origin of these speleothems. The same may be true of many occurrences of cave popcorn as well.


Chapter 3.3

GEYSERITES OF YELLOWSTONE NATIONAL PARK: AN EXAMPLE OF ABIOGENIC “STROMATOLITES”

M.R. Walter

INTRODUCTION

The siliceous geyserites of Yellowstone National Park (Wyoming, U.S.A.) are morphologically very similar to stromatolites, but are abiogenic. This study was undertaken because geyserites in the geological record could easily be mistaken for stromatolites, leading to erroneous palaeoenvironmental interpretations. Moreover, the recognition of fossil geyserites would provide useful data on the chemistry of palaeogroundwaters and on palaeoenvironmental settings. During the Precambrian, before the advent of silica-secreting organisms, geyserite-like sediments may have formed in a wide range of environments. The Y ellowstone geyserites are here described and features which allow their differentiation from stromatolites are noted.

The word geyserite is used in the sense of White et al. (1964) to mean microbanded, botryoidal, opaline silica that is deposited in the proximity of spring vents and fissures; as used here it also includes abiogenic, stratiform sinter. Geyserite is only one of several forms of siliceous sinter (White et al., 1964; Walter, Ch. 8.8). The spring and geyser deposits of Yellowstone Park have been described previously by Peale (in Hayden, 1883), Weed (1889a) and Allen and Day (1935). Walcott (in Anonymous, 1916) has fig- ured Yellowstone geyserites. The chemistry of the hot spring silica has been discussed by White et al. (1956) and by Morey et al. (1964). Chemical analyses of Yellowstone thermal waters have been summarized by Rowe et al. (1973). Yellowstone geyserite has been compared with a particular type of Precambrian banded iron formation “stromatolite” (Walter, 1972b). Geyserite morphogenesis is discussed in this last paper and has been commented on by Oehler (1972). Austen (1873) has discussed the formation of opaline pisolites in Yellowstone hot springs.

88

DISTRIBUTION OF GEYSERITE

M.R. WALTER

The distribution of geyserite is determined by the location of vents, the temperature of and rate of heat loss from the issuing water, the concentration of silica in the water, and the pH of the water. The first factor is discussed elsewhere (Walter, Ch. 8.8). The temperature and rate of heat loss determine the distance from the vents of the 73OC isotherm; hotter waters are almost sterile and the silica deposition is probably unaffected by biological activity, but at temperatures below about 73°C alkaline waters are abundantly populated by the unicellular cyanophyte Synechococcus and the filamentous bacterium Chloroflexus, and the silica deposits (sinter) take on various biogenic forms (Walter et al., 1972; Doemel and Brock, 1974; Walter, Bauld and Brock, Ch. 6.2). The pH determines the rate of polymerization and precipitation of the dissolved silica: both of these processes are extremely slow in acid waters but rapid in alkaline waters (see the summary by Walter, 197213). As a result, most acid springs and geysers have no geyserite; an exception is Echinus Geyser, with a pH of 3.5 (Rowe et al., 1973), which is surrounded by ferruginous spicular geyserite (but this is unlike that in alkaline waters; see below). Some waters that are acid at the point of issue become alkaline in surface pools owing to C02 loss (Rowe et al., 1973). Thus, geyserite is largely restricted to areas submerged in or splashed by silica- charged alkaline water hotter than 73°C. Since the 73" isotherm is rarely more than 5- 10 m from the spring or geyser vents, geyserite is very restricted in occurrence.

MORPHOLOGY OF GEYSERITE

The shape of spring and geyser vents varies from chimney-like (Fig. 1) through mound-shaped (Fig. 2) to saucer-shaped (Figs. 3, 4, 5 ) . A section through a mound-shapedvent is shown in Fig. 6. Sections of other vent forms have not been observed, but I have previously published a hypothetical cross- section of a saucer-shaped vent (Walter, 1972b). The form of the vents is presumably related to the rate of deposition of the silica and the rate and mode of discharge of the water.

Geyserite has the following range of forms: stratiform (Fig. 7), spicular (Figs. 8, 9), columnar (Figs. 10, ll), pseudocolumnar (Figs. 12-14), and pisolitic and oolitic (Figs. 15-17). Stratiform geyserite frequently has a gently undulatory form and locally contains ripple mark-like structures (see Walcott in Anonymous, 1916 his Figs. 21, 22). Spicular geyserite is abundant in Echinus Geyser (where it is aluminous and ferruginous, unlike other Yellow- stone geyserites: M.R. Walter and J.B. Jones, unpublished). There it is the only form of geyserite. Elsewhere, spicular geyserite occurs frequently at the innermost margins of pool rims; the rims are dominated by columnar or


Fig. 1. Lone Star Geyser.

pseudocolumnar geyserite. The spicules are mostly 0.5-1.0 mm wide and are up to 10 mm long. In Echinus Geyser they form radiating clusters, like pins in a pin cushion; elsewhere they are mostly near-vertical. The columns of columnar geyserite range from about 1.0mm to 30mm wide, but most are about 10 mm wide. They have up to 25 mm of vertical relief and are near- vertical or in radiating clusters. In transverse sections they are subcircular, polygonal, or lobate. Branching of columns ranges from infrequent to frequent and varies in form from a-parallel to slightly divergent (see glossary). Pseudocolumnar geyserite is Similar to columnar geyserite, but has only a few millimetres of intercolumn relief. The ooids and pisolites range in width from about 0.5 mm to 50 mm and in shape from subspherical through polygonal to irregular. They have smooth or rugose surfaces; the rugose forms have a fine columnar or pseudocolumnar internal structure.

Spicular, columnar and pseudocolumnar geyserite frequently form rounded mounds several tens of centimetres wide (e.g. Fig. 18).

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Fig. 2. Minute Man Geyser, Shoshone Geyser Basin. Hammer for scale.

Fig. 3. Steep Cone Spring, Sentinel Meadows. The area shown is on the crest of a steep. sided mound of sinter. Hammer for scale.


Fig. 4. Geyser in the Rustic Geyser group, Heart Lake Geyser Basin. 10 cm scale on left.

Fig. 5. Twin Geyser, Shoshone Geyser Basin.

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erosional margin_ flL+!-

I 1 metre I

W S t r a t i f a r r n geyserite / Columnar geyserite m Fig. 6. Sketch of a natural erosional section through Mortar Geyser, Upper Geyser Basin. No vertical exaggeration.

Fig. 7 . Subaqueous stratiform geyserite between logs of wood encrusted with columnar geyserite (where they project above the water). Rustic Geyser, Heart Lake Geyser Basin. Scale 10 em.

Geyserite lamination forms are illustrated in Figs. 19-26. The laminae vary from gently to steeply convex. They may extend down column margins as a coating (wall), may end abruptly at the margins, or may project past the margins as cornices. Former open spaces within the geyserite are marked by


Fig. 8. Spicular geyserite in Echinus Geyser. Norris Geyser Basin. Scale 10 cm.

Fig. 9. Spicular geyserite on the innermost margin of Pearl Geyser, Norris Geyser Basin. Scale 10 cm.

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Fig. 10. Columnar geyserite, on the rim of a spring near Flat Cone Spring, Sentinel Meadows (vent to the left). Nail marks the site of a deposition rate experiment. Scale 10 cm.

Fig. 11. Columnar geyserite, on the rim of Pearl Geyser, Norris Geyser Basin (vent towards the bottom). Scale 10 cm.


Fig. 12. Top view of a rim of pseudocolumnar geyserite, unnamed small spring uphill from Twin Geyser, Shoshone Geyser Basin (vent towards the top). Scale 10 cm.

Fig. 13. Side view of the same rim as in Fig. 12, showing the pseudocolumnar structure and rhythmically distributed cornices (looking away from the vent). Scale 10 cm.

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Fig. 14. Top view of a rim of pseudocolumnar geyserite, geyser near Steep Cone Spring, Sentinel Meadows (vent towards the top). Scale 10 cm.

patches of unl inated silica, in places enclosed by concentrically laminated silica (Fig. 27)~icrocrosslamination, particularly characteristic of geyserite, is frequently present in spicular, columnar and pseudocolumnar geyserite (Figs. 28-30), but is not ubiquitous. Microdisconformities occur infrequently (Fig. 27), as do ridges and spicules on laminae (Fig. 29).

MICROSTRUCTURE OF GEYSERITE

The basic lamination consists of laminae 0.5-4.0 pm thick (Fig. 31) which are laterally extensive and have more or less parallel, abrupt boundaries (a banded microstructure). These are commonly grouped into darker and lighter macrolaminae 10- 150 pm thick (Fig. 32). The macrolaminae are discon-


Fig. 15. Ooids and pisolites in a little pool on the flank of Bead Geyser, Lower Geyser Basin. Scale 10 cm.

Fig. 16. Part of the area shown in Fig. 15. All of the grains, including the large platy polygonal forms on the left, are concentrically laminated. Scale 10 cm.

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Fig. 17. Pisolite gravel which surrounds a small spring vent near Bead Geyser, Lower Geyser Basin. These pisolites have a columnar internal structure. Scale 10 cm.

Fig. 18. Mounds of columnar geyserite on the flank of Minute Man Geyser, Shoshone Geyser Basin. The pools between the mounds contain ooids, pisolites and intraclasts. Scale 10 cm.


Fig. 19. Thin section of stratiform geyserite, Rustic Geyser, Heart Lake Geyser Basin. Transmitted light. Scale bar 1 cm,

tinuous (forming a streaky microstructure). Very fine cracks perpendicular to the laminae are common (Fig. 32); they possibly formed during drying pre- paratory to thin-sectioning (but see Nolan and Anderson, 1934, p. 222).

One specimen of subaqueous geyserite has been studied in thin section. This has some layers with a finely banded microstructure but much of it is only coarsely laminated (Fig. 19).

CAUSE OF SILICA PRECIPITATION

This discussion refers to waters hotter than about 73’C. Above this temperature only non-photosynthetic bacteria live in the Yellowstone springs (Bott and Brock, 1969; Brock and Freeze, 1969; Brock, 1970; Brock and Darland, 1970) and these are relatively sparse and sporadic; so sparse indeed that these waters were once thought to be sterile: “In the great majority of the hot alkaline springs the clear waters are entirely free from visible organisms, and Dr C.B. van Niel, microbiologist, who, with Mr Lewis A. Thayer, made observations in the Park in 1929, pronounced them sterile” (Allen, 1934, p. 383). Very infrequently the bacteria form brownish coatings and

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Fig. 20. Thin section of spicular geyserite encrusting weathered igneous rock, Echinus Geyser, Norris Geyser Basin. Transmitted light. Scale bar 1 cm.

white stringy masses in the springs. I t is likely, therefore, that silica precipitation at temperatures more than 73°C is inorganic (Allen, 1934).

Even at temperatures lower than 73°C microorganisms play only a minor role, if any, in silica precipitation. “The evidence is clear that algae are, at best, only a minor factor in causing polymerization and precipitation of silica at Steamboat Springs” (White et al., 1956, p. 39). Silica precipitation to form geyserite apparently is due to cooling and evaporation of the spring and geyser waters (Krauskopf, 1956; White et al., 1956). I have previously summarized the literature which demonstrates that both polymerization and precipitation of silica is very much faster in alkaline than in acid solution (Walter, 1972b). The processes involved are not well known. However, it is known that the silica in hot springs is in true solution (as monosilicic acid).


Fig. 21. Thin section of columnar geyserite, “Column Spouter” of Walter (Ch. 6.2), Fairy Creek meadows. Transmitted light. Scale bar 1 cm.

Fig. 22. Thin section of columnar geyserite, spring near Flat Cone Spring, Sentinel Meadows. Transmitted light. Scale bar 1 cm.

Subaqueous deposition results from polymerization of the dissolved silica and precipitation of the polymer. In subaerial deposition, where evaporation is rapid, the larger polymeric molecules may not have time to form (White et al., 1956). This may account for the observed microstructural differences between subaqueous and subaerial geyserite.

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Fig. 23. Columnar geyserite with cornices, unnamed spring uphill from Twin Geyser, Shoshone Geyser Basin. Transmitted light. Scale bar 1 cm.

Fig. 24. Thin section of pseudocolumnar geyserite, spring near Steep Cone Spring, Sentinel Meadows. Transmitted light. Scale bar 1 cm.


Fig. 25. Thin section of a pisolite. The nucleus is.a fragment (intraclast) of geyserite. Bead Geyser, Lower Geyser Basin. Transmitted light. Scale bar 1 cm.

Fig. 26. Thin section of pisolites with a columnar internal structure, from around a small spring near Bead Geyser, Lower Geyser Basin. Transmitted light. Scale bar 1 cm.

ORIGIN OF GEYSERITE LAMINATION

Geyserite lamination observable with an optical microscope is 0.5- 4.0 pm thick. Preliminary observations with a scanning electron microscope suggest that even thinner laminae are present (M.R. Walter and J.B. Jones, unpublished). Super-laminar depositional periodicity is indicated by the frequent presence of regularly spaced overhanging groups of laminae (cornices) on the margins of geyserite columns (Figs. 13,23, 30,33). The cornices are usually spaced 0.5-1.5mm apart.

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Fig. 27. Thin section showing microdisconformity and concentrically laminated silica surrounding open spaces (some of which are partly filled by unlaminated deposits of silica); pseudocolumnar geyserite; unnamed spring uphill from Twin Geyser, Shoshone Geyser Basin. Transmitted light.

Fig. 28. Thin section showing microcrosslamination in columnar geyserite. The “foresets” dip quaquaversally away from the column axis. Pearl Geyser, Norris Geyser Basin. Trans- mitted light.


Fig. 29. Thin section showing microcrosslamination and spicules within layers (lines perpendicular to laminae are cracks). Unnamed spring uphill from Twin Geyser, Shoshone Geyser Basin. Transmitted light. Scale bar 1 mm.

In an attempt to discover the temporal significance of geyserite lamination, deposition-rate experiments were conducted at the following springs and geysers: Pearl Geyser, Echinus Geyser, a spring near Steep Cone Spring, a spring near Flat Cone Spring, an unnamed geyser on the bank of the Firehole River near the Fountain Freight Road (several hundred metres upriver from Ojo Caliente), a small geyser near Bead Geyser, a geyser in Fairy Creek meadows (“Column Spouter” of Walter, Ch. 6.2), Bronze Geyser and a geyser uphill from Twin Geyser. The experiments lasted from 29 to 91 days, terminating in July 1972. In the experiments, columnar, spicular, and pisolitic geyserite were marked in situ with dyes (felt Den ink or methylene blue) and were grooved with a small triangular file. The dyed and grooved geyserites were then sampled after a known period, and thin-sectioned for microscopic examination; the dyes and grooves are clearly visible in thin section and provide a precise datum. The laminae above the markers were counted and their numbers were compared with the numbers of days elapsed (Table I).

In all cases there are more laminae than days elapsed, in one case more than five times as many. The number of visible laminae is fewer than would be expected if they represented individual eruptions (Table I). The very thin (uncounted) laminae observed with an electron microscope may represent single eruptions. On the basis of the average deposition rate and the average distance between cornices, it was estimated for one sample of geyserite that the cornices probably form roughly once a year.

TABLE I

Comparison of numbers of laminae and numbers of days elapsed of some Yellowstone geysers

Geyser or Spring Terminal date Days Visible Average no. visible Eruption (1972) elapsed laminae laminaelday period

Pearl 21 July Near Steep Cone 20 July Near Flat Cone 20 July Near Fountain Freight Road 21 July “Column Spouter” 20 July Bronze 23 July Uphill from Twin 23 July

- 31 174 5.6 90 168 1.9 90 140,124,132 1.5 -

31 50 1.6 - 91 100 1.1 70 min 29 90,101 3.3 10-20 min

-

29 50,70 2.1 -


Fig. 30. Thin section showing microcrosslamination, open spaces and cornices in columnar geyserite. Locality as for Fig. 29. Transmitted light.

It is probable that columnar geyserite accurately records the eruption history of geysers and may be used to study their past periodicities. This may provide data on Earth tides and earthquakes, both of which have been shown to influence geyser eruption intervals (Rinehart, 1972a, b).

GEYSERITE MORPHOGENESIS

The distribution of forms of geyserite provides some clues to their morphogenesis. Columnar and spicular geyserites are exclusively subaerial and generally occur in splash zones around geysers and springs; although some occur in areas of intermittently flowing water (where it is intermittently exposed to the atmosphere). Oolitic and pisolitic geyserites occur in turbulent waters,

108 M.R. WALTER

Fig. 31. Thin section showing the typical finely banded lamination of geyserite. Columnar geyserite, Pearl Geyser, Norris Geyser Basin. Transmitted light.

Fig. 32. Thin section showing coarse streaky lamination in columnar geyserite (structures perpendicular to lamination are cracks). Columnar geyserite, Pearl Geyser, Norris Geyser Basin. Transmitted light. Scale bar 250 pm.


Fig. 33. Columnar geyserite, near Steep Cone Spring, Sentinel Meadows. Specimen from inner margin of pool rim, showing regularly spaced cornices and bridging laminae. Also shown are tiny spicules directed towards the spring vent. Scale bar 1 cm.

water droD

Fig. 34. Diagrammatic illustration of the deposition on geyserite columns of asymmetric deposits of silica, eventually leading to the formation of microcrosslamination. A silica- charged drop of water adhering to the top of a column is deformed by the force of gravity; it evaporates, leaving a lens of silica.

and may be permanently or intermittently subaqueous. Most stratiform geyserite is subaqueous. It is apparent from these distributions that the morphogenetic processes producing columnar and spicular geyserite only operate subaerially . A result of these processes is the localization of silica deposition; generalized deposition would produce a featureless deposit (stratiform geyserite). Observations in areas of active columnar and spicular geyserite formation showed that during eruption water splashed over the columns and spicules. The water either remained as droplets or spread out over the geyserite. It drained away from between the columns and spicules, except in

110 M.R. WALTER

some tiny “enclosed basins”. Between eruptions the water evaporates. Evap- oration of droplets produces localized silica deposits. I t is likely, therefore, that localized deposition is due to the formation of these water droplets, a surface tension effect only possible above water. The cause of repeated localization near the same point, and therefore column accretion, is not apparent (see also Lebedev, 1967, pp. 59-60). Water droplets adhering to the tops of geyserite columns presumably evaporate to give tiny asymmetric silica deposits that accrete upwards to produce the microcrosslamination (Fig. 34).

The reasons for interpreting the morphogenetic processes as being noa- biological have been discussed previously (Walter, 1972b). They are the occurrence of the geyserite in environments that frequently are virtually sterile (discussed above), the absence of mineralized mats of algae or bacteria in thin sections of geyserite, and the rarity of algae and bacteria in the residue produced by hydrofluoric-acid dissolution of geyserite. To these reasons may be added the fact that in twenty specimens of geyserite studied with a scanning electron microscope, moulds of algal and bacterial filaments are rare both within columns and in the intercolumnar sediment (M.R. Walter and J.B. Jones, unpublished). Such moulds as were observed appeared to be of short fragments of algal filaments (i.e. probably organic detritus).

POTENTIAL RANGE OF ENVIRONMENTS OF FORMATION

Inorganic deposition of silica is presently greatly restricted since in water cooler than about 30°C diatoms, radiolaria, silicoflagellates, siliceous sponges and other silica-secreting organisms efficiently extract any available silica, keeping concentrations far below the saturation level for amorphous silica (Krauskopf, 1956; Siever, 1957). In the absence of silica-secreting organisms surface waters would become supersaturated with respect to silica (silicate synthesis in sedimentary environments is not quantitatively significant as an inorganic silica sink; Calvert, 1968). Silica-secreting organisms appear to have been present throughout Phanerozoic time, but the rapidly expanding knowledge of Precambrian fossils provides no evidence for their existence then (Glaessner, 1962; Cloud, 196813; Schopf, 1970). During the Precambrian there was apparently no organic silica sink, so silica would have precipitated inorganically (Holland, 1971) and geyserite-like sediments could have formed in a wide range of environments. Silica-deposition rates were probably very, low (as they are now in thermal waters), so geyserite-like sediments would have formed only where little other sediment was being deposited. Under these conditions, spicular and columnar geyserite-like deposits would not be restricted to areas of hydrothermal activity, but would be characteristic of deposition in any type of splashing water environment where the water was rich in silica. One such environment may have been along the shorelines of the Precambrian banded iron formation basins. Columnar geyserite-like struc-


tures do occur in the Gunflint Iron Formation of Ontario (Walter, 197213). These are microstructurally distinct from the true stromatolites of the Gun- flint Iron Formation.

An appropriate term for geyserite-like deposits preserved in the geological record is “stiriolite” (from the latin stiriu, a frozen drop, alluding to the genesis of columnar and spicular geyserite from drops of water in splash zones). This term should be used if the environment of deposition is unknown, or if the deposit formed in some environment other than around geysers and hot springs. Thus the geyserite-like deposits in the Gunflint Iron Formation at Mink Mountain, Ontario (Walter, 1972b) may be called stiriolites (columnar, stratiform, etc.). On the other hand, deposits demonstrably resulting from hydrothermal activity should be termed geyserites.

CRITERIA FOR DISTINGUISHING GEYSERITES FROM STROMATOLITES

Three main criteria distinguish geyserites from stromatolites. Firstly, geyserite occurs only around hot springs and geysers, immediately adjacent to the place of water discharge. Ideal outcrops of fossil geyserites would be expected to show geyser vents. Secondly, in my experience, microcrosslamination is a feature found only in geyserite. Finally, the extremely thin, regular lamination of geyserite is much thinner and more regular than the lamination of any assured stromatolites known to me (it is comparable to the lamination of Holocene carbonate ooids: Bathurst, 1971, p. 78). In addition, some stromatolites show a distinct biogenic fabric (e.g. oriented filament moulds; Bertrand-Sarfati; Ch. 5.2) which does not occur in geyserite (though geyserite may contain detrital organic matter). All of these criteria, except perhaps distribution, could also be used to distinguish stiriolites from stromatolites.

SUMMARY

Geyserite is opaline silica deposited nonbiogenically within and around hot springs and geysers. The deposits have a variety of shapes, each characteristic of particular environments. The shapes strikingly resemble those of stromatolites, but the morphogenetic processes are nonbiological. Columnar and spicular geyserite forms in subaerial splash zones; stratiform geyserite forms subaqueously; oolitic and pisolitic geyserites form in turbulent water, and may be continually or only intermittently submerged.

Geyserite is distinguished from stromatolites by its distinctive distribution around points of water discharge, by microcrosslamination that is common but not ubiquitous, and by its banded lamination with laminae less than about 4pm thick (ranging down to less than 0.5 pm thick).

Before the advent of silica-secreting organisms (probably during the Cam-

112 M.R. WALTER

brian) nonbiologic deposition of silica was probably widespread in oceans, lakes and rivers, and geyserite-like deposits could have formed in a wide range of environments. Spicular and columnar forms could have formed in many types of water-splash zones, such as along shorelines. The word stiriolite is introduced for geyserite-like deposits preserved in the geological record.

ACKNOWLEDGEMENTS

The U.S. National Park Service permitted field work in Yellowstone Park. Professor T.D. Brock provided laboratory facilities in West Yellowstone. Financial support was provided by the Department of Geology and Geo- physics, Yale University.

4. BIOLOGY OF STROMATOLITES

Chapter 4.1

ORGANISMS THAT BUILD STROMATOLITES

Stjepko Golubic

INTRODUCTION

Most stromatolites, fossil and Recent, are interpreted as organosedimentary structures formed by sediment trapping and binding and/or mineral precipitation within prostrate microbial communities termed algal mats. Algal mats are formed, modified or destroyed by an array of different microorganisms and under diverse environmental conditions. I t is therefore necessary to discuss the organisms that build stromatolites within the ecological context of their activity. In the account that follows I will: (1) relate algal mats as Recent microbial ecosystems to stromatolites; (2) discuss basic mechanisms and adaptations involved in the microbial construction of stromatolites; (3) discuss mechanisms involved in microbial stromatolite destruction; and (4) illustrate the distribution of algal mat types using two marine intertidal environments as examples.

ALGAL MATS AND STROMATOLITES

An algal mat can be viewed simultaneously as a morphological entity, as a microenvironmental modifier, as a differentiated microbial community, and as an ecosystem.

Algal mats form more or less cohesive fabrics of intertwined filaments and/or gelatinous matter produced by both filamentous and coccoid microorganisms. These algal fabrics spread over loose sediment, trap and bind particles deposited on them, and thus stabilize the sediment-water interface. Algal mats range in cohesiveness from loose mucilaginous films to hard lithified crusts. They, and the associated sediment, display colors including all shades of bluegreen, brown or pink. They display various surface textures ranging from smooth and velvety to mamillate or tufted.

Biologically important environmental factors such as light, oxygen supply, and ionic composition of the interstitial water vary along steep gradients arranged perpendicularly to the mat surface. Metabolic activities of various

114 S. GOLUBIC

components of the mat community are largely responsible for stabilization and intensification r>f these ecological conditions. In this sense, algal mat communities exhibit a marked capacity for homeostatic regulation and maintenance of the immediate environment, niche differentiation, diversity regulation and induction of successional changes. This biological control of the microenvironment affects directly or indirectly the trapped and underlying sediments as well as the conditions leading to mineral precipitation or dissolution.

Differentiation of the microbial community frequently follows the pattern of microgradients and results in a biological stratification of the mat. Light normally penetrates the upper few millimeters of the mat establishing a euphotic zone where organic matter is produced photosynthetically. Pri- mary production of Lyngbya conferuoides (C. Agardh) Gomont mat in the lagoons of the Texas coast is relatively low at the fully illuminated mat surface, reaches its maximum within 1-4mm below the surface, and then declines toward the dark interior of the mat (Sorensen and Conover, 1962). This mat appears to be similar to large-scale aquatic environments, such as lakes and oceans. However, no inhibition of photosynthesis at the mat surface was found in the algae of the Yellowstone thermal springs, which are adapted to high light intensity (Brock and Brock, 196913). In addition to scattering and quantitative light attenuation, photosynthetic microorganisms within the euphotic zone of the mat have a filtering effect on specific wave lengths of the passing light. They are well adapted to absorb unused spectral portions of attenuated and qualitatively altered light by the combination of their photosynthetic pigments. The concentrations of C02 and 02, pH and Eh are affected by varying rates of photosynthesis uersus respiration along the light attenuation gradient. Light and redox potential largely determine the sequence of microorganisms in the mat.

Specialized microbial populations occupy zones of their ecological optima within the mat producing biological strata expressed in cross-section as colored bands. A flat mat from the intertidal zone of the Persian Gulf (Golubic, 1973a, p.458) serves as an example of a well-differentiated, stratified algal mat (Fig. 2c). It is composed of the following layers each approximately 1 mm thick: (1) brown surface layer; (2) blue-green layer; (3) salmon- pink layer; (4) purple-pink layer; and (5) black layer. In the surface layer, Lyngbya aestuarii (Mertens) Liebman seems protected from excessive solar radiation by a brown extracellular sheath pigment “scytonemine” (Kylin, 1937), and in turn provides a shield for the entire community below. The blue-green layer is occupied by Microcoleus chthonoplustes Thuret which here achieves maximal biomass, but lacks the protective sheath coloration. Below these layers, in the transitional zone between the aerobic and anaerobic portion of the mat, is the salmon-pink layer of pigmented filamentous bacteria, followed by the purple sulphur bacteria in the anaerobic purple- pink layer below. The black iron sulfide-stained layer contains mainly non- photosynthetic, anaerobic bacteria.

BIOLOGY OF STROMATOLITES 115

Algal mats are therefore microbial ecosystems which occupy and modify a microenvironmerit several millimeters deep at the sediment-water interface. At the primary production level, algal mats are dominated by cyanophytes and to a lesser extent by eucaryotic algae and photosynthetic bacteria. In rapidly flowing streams and waterfalls travertines are formed by mosses as well as algae. The mats may support (Monty, 1972) or be controlled by (Garrett, 1970b) populations of protozoa, insects and larger animal grazers. However, under extreme environmental conditions (e.g., salinity, temperature, and desiccation) the consumer level of algal mats is reduced or absent. A complex and diverse but still poorly known flora of bacterial, and to a lesser extent fungal decomposers is always present in algal mats. Both aerobic and anaerobic decomposition of o’rganic matter regularly take place within algal mats.

Algal mats are found in environments characterized by various rates of sedimentation, encrustation and burial. Several microorganisms adapted to these conditions respond by upward growth and/or movement, rapidly colonizing newly deposited particles. Under continuing sedimentation, the entire community becomes gradually displaced upward, adjusting their position in relation to the constantly shifting surface. Under such dynamic conditions the competing rates of organic production uersus sediment deposition determine the residence time of the mat in a given sediment layer, i.e. the time during which any biogeochemical interactions may affect algal mat morphology. The difference between primary production and total decomposition rates determines the amount of organic-matter residue, potential material for preservation and fossilization.

The fossilized, cumulative record of algal mat activities is left behind in the sediment as stromatolites. Variations in species composition and microstructure of algal mat communities an,d/or of sedimentation-affecting conditions in the course of stromatolite formation may be recorded as laminae. The configuration of individual laminae, or synoptic profile (Hofmann, 1969a) reveals past positions of actively accreting stromatolite surface. It is the live algal mat community, several millimeters thick which in time moves up the vertical stromatolite profile.

MICROBIAL STROMATOLlTE CONSTRUCTION

The presence of microorganisms within a Recent or fossil stromatolite docs not necessarily imply any causal relationship with the genesis of the structure (Hofmann, 1973). With respect to their function microorganisms may contribute to the construction of the structure, be passively present, or contribute to its destruction. Both organic and inorganic components of a stromatolite as well as the microchemical conditions within it are potentially subject to microbial influence. Microbial photosynthesis and respiration may

116 S . GOLUBIC

TABLE I

Principle adaptational types of microorganisms (see also drawings in text)

A. Movement ~

By unidirectional gel By gliding: production: coccoid forms filamentous forms

1. Entophysalis (cyanophyte) 5 . Chloroflexus (chlorobacterium) 2 . Cyanostylon (cyanophyte) 6. Microcoleus (cyanophyte) 3. Cymbella (diatom) 7. Lyngbya (c yanoph yte) 4 . Oocardium (desmid) 8. Phormidium (cyanophyte)

B. Growth

Radial or upright Prostrate or reticulate

9. Rivularia (cyanophyte) 10. Gardnerula (cyanophyte) 11. Calothrix (cyanophyte) 12. Gongrosira (chlorophyte)

13. Geitleria (cyanophyte) 14. Scytonema (cyanophyte) 15. Caulerpa (chlorophyte) 16. Vaucheria (xanthophyte)

interfere with carbonate equilibria of interstitial waters and cause carbonate precipitation and dissolution. Only dominant or very abundant organisms can be expected to play any significant role influencing the microenvironment or the structure of stromatolites. The presence of rare species, however, contributes to the species diversity of a microbial community, which in turn may reflect the ecological conditions of the habitat. Generally, species diversity is inversely proportional to the harshness of environmental conditions. Highly fluctuating intertidal environments, for example, are dominated by few, exclusively procaryotic species. Under more favorable conditions in permanently submerged environments, the species diversity increases and both procaryotic and eucaryotic organisms are present.

Stromatolite-building microorganisms colonize both hard and loose substrates preparing an extraordinarily broad spectrum of benthic environments for potential stromatolite construction. The explanation for the limited occurrence of Recent stromatolites probably includes competition for space by eucaryotic algae, aquatic plants, or reef-building organisms and competition for mineral support by various skeleton-producing eucaryotes (Fischer, 1965; Monty, 1972), as well as destruction by grazers (Garrett, 1970b) and bacteria (Brock, 1967; Golubic, 1973a).

Once established, the stromatolite growth follows one basic common scheme: By serving as traps for suspended particles or mineral crystallization, the microorganisms become buried and encrusted. This condition supplies selective pressure for an evolution of efficient escape mechanisms. The evolution of such adaptations was based either on strategy of movement or of growth out of sediment, and included both coccoid and filamentous forms. The principal adaptational types are exemplified in Table I. A large


number of microbial species can be classified within these types according to their stromatolite-building activity. Drawings in the text representing some of these types carry the same numbers. The scale bar is 10pm unless stated otherwise. The older literature has been reviewed by Pia (1934).

The simplest and presumably most primitive microorganisms are non-motile coccoid unicells. It is difficult to conceive how such a morphotype could cope with conditions of perpetual sediment burial, yet several adaptations evolved which enable unicells to function as stromatolite building agents. Entophysalis major Ercegovid (1 in Table I; Plate I, 1) is a coccoid cyano-

phyte which forms brown, gelatinous, pustular or mamillate mats and lithified domal stromatolites in the Bahamas, Persian Gulf and Shark Bay (Figs. le, 2b) (Golubic, 1973a; Golubic and Awramik, 1974; Logan et al., 1974). Morphologically equivalent counterparts built stromatolites in the > 1.8 x lo9 years old Kasegalik Formation of the Belcher Islands (Hofmann, 1974; Golubic and Hofmann, 1976). Massive gel production provides a coherent sediment coating. The gel excretion is predominantly unidirectional causing cells to be arranged in vertical rows and growth of “mamillae” on the surface of the mat. Sediment is trapped in depressions between mamillae and

then overgrown by the mat. When sedimentation rate increases, free cells are released by the colony which recolonize the surface through passive transport by currents.

A more pronounced case of unidirectional gel production is represented by Cyanostylon species (2; Plate I, 4) on Bahamian intertidal flats. The alga produces a basal

Ecologically equivalent morphotypes which all provide a unicellular organism with properties of a basally attached filament, evolved convergently in several genera of diatoms: in a marine Stephanodiscus in Shark Bay, and in Gomphonema and Cymbella (3).in freshwater (Hufford and Collins, 1972). In travertine depositing environments gelatinous stalks of diatoms become

encrusted with calcite (Wallner, 1935; Golubic, 1957). Most specialized of these sessile unicells is the desmid Oocardium stratum

Niigeli (4) in calcareous springs and travertine waterfalls (Golubic, 1957). This alga, ecologically unique among the desmids, attaches to the substrate and produces a gelatinous collar which &mediately calcifies (Golubic and Marcenko, 1958). As in Gomphonema and Cymbella (but unlike Stephano- discus), the cell. division results in ramification of the stalks. This is determined by the plane of cell division in relation to stalk production. Ultimate- ly, the progeny of a single Oocardium cell produces a calcified “tree” of

@&- \... &- . /?.i ‘%, .;, :: .% ’ ..

gelatinous stalk elevating itself out of the sediment.

118 S. GOLUBIC

PLATE I

Representative stromatolite-building microorganisms (cyanophytes); scale bar in the middle is 10 pm long. 1. Entophysalis m i o r . 2 . Phormidium hsndersonii. 3 . Sheaths of Phormidium hendersonii with entrapped sediment particles. 4 . Cyanostylon sp. 5. Schizothrix splendida. 6 . Lyngbya aestuarii.


branched stalks maintaining an optically uniform crystallization pattern. In this manner, a biologically highly modified calcite monocrystal is formed for each clonal line of the alga (Wallner, 1933).

Gliding motility of filamentous bacteria and cyanophytes is the most efficient way t o escape burial by sediments. Forms which do not produce cohesive gelatinous

species of Oscillatoria (cyanophyte), Beggiatoa (sulphur- oxidizing bacterium), and various flexibacteria, are present in most algal mats and can produce susbstantial amounts of organic matter (Anagnostidis and Schwabe, 1966; Castenholz, 1968), but their contribution to the formation of permanent organosedimentary structures is limited. The non-sheathed filamentous chlorobacterium Chloroflexus (5) forms stratiform laminated mats in Yellowstone thermal springs and provides a base for colonization of Synechococcus, a unicellular non-sheathed cyanophyte (Doemel and Brock, 1974). When associated with the sheathed filamentous cyanophyte Phormid- ium and with silica precipitation, Chloroflexus participates in formation of lasting conical stromatolites (Walter et al., 1972).

Sheathed filamentous cyanophytes represent a well- adapted type of stromatolite-building microorganism. They were the first stromatolite builders recognized (Cohn, 1862) and were long thought to be the only kind capable of laminated stromatolite construction (Schopf et al., 1971). The algae colonize the substrate in a free gliding state (hormogonia). Subsequently the growing

trichomes produce a firm gelatinous sheath which remains in the sediment, lending a more permanent cohesiveness to the structure. Microcoleus chthonoplastes (6) is the most common species in marine algal mats. A combination of high motility, production of cohesive fabric and a pronounced tolerance for salinity fluctuation and microaerobic conditions accounts for its dominance in a variety of stromatolitic structures, particularly in vertically differentiated stratiform mats in stagnant tidal pools and channels (Fig. 2c). The trichomes of Microcoleus form twisted bundles retaining their motility within a common gelatinous sheath.

A similar type with fewer, less motile trichomes in a common sheath is represented by several species of Schizothrix.’Schizothrix cresswellii Harvey, S. gracilis Golubic and S. spzendida Golubic (Plate I, 5 ) form finely laminated structures in well-drained, periodically air-exposed algal mats (Fig. 2 d ) (Walter et al., 1973, Golubic, 1973b). -Schizothrix lateritia (Kutzing) Gomont, S . coriacea (Kutzing) Gomont, S . penicillata (Nageli) Gomont and S. fasciculata (Kutzing) Gomont form hard and concentrically laminated colonies in freshwater in conjunction with carbonate precipitation (Geitler, 1932).

Lyngbya aestuarii (7; Plate I, 6) is motile in its hormogonal state. Settled

120 S. GOLUBIC

trichomes produce individual sheaths which may be thick and layered. In response to intensive illumination these sheaths are stained yellow to brown. This species forms a protective surface layer to vertically differentiated algal mat communities of the Persian Gulf (Figs. 2c ,d , e ) and Shark Bay. By a combination of growth and movement L. aestuarii protrudes up to 20mm out of

the sediment forming radiating bunchlets which collapse into pointed tufts during low tides. This “paint-brush” arrangement acts as an efficient sediment trap resulting in the tufted and reticulate mat formation of Shark Bay (Fig. If) . Other species of Lyngbya also form mats: L . majuscula Harvey in subtidal, and L. semiplena J. Agardh (Golubic, 1973a) and L. confervoides (Sorensen and Conover, 1962) in intertidal environments. Hard calcareous crusts in springs and rivers of carbonate-rich regions are formed by L. aerugi-

Phormidium hendersonii Howe ( 8 ; Plate I, 2, 3) forms cartilaginous finely laminated biscuits by sediment entrapment on the hemispherical smooth surface of the colonies (Fig. 2a) (Golubic and Focke, 1976). It was first described by Howe (1918) and later referred to under various names (P. crosbyanum Tilden em. Drouet, by Drouet, 1942; Symploca laete-viridis Gomont em.

Drouet, by Ginsburg and Lowenstam, 1958; Schizothrix calcicola (Agardh) Gomont em. Drouet, by Monty, 1965b; and Schizothrix sp., by Golubic, 1973a). The movement of trichomes and subsequent production of common hard gel results in an alternate arrangement of vertical and horizontal filaments which together with entrapped sediment produce noctidiurnal laminations (Monty, 1967). The freshwater species Phorrnidium incrustaturn (Nageli) Gomont forms hard lithified crusts and oncolites in calcareous streams (Fritsch, 1950b; Kann, 1973; Golubic and Fischer, 1976) similar to those formed by Lyngbya aerugineo-coerulea.

1 mm Non-motile filamentous forms cope with intensive sedimentation rates by means of a correspondingly rapid growth. Radiating filaments of Rivularia haernatites (De Candolle) Agardh (9) in calcareous streams and waterfalls are embedded in common gel which becomes a microenvironment of autochthonous calcite precipitation, and ultimately hardens into a concentrically

laminated stromatolitic structure (Jaag, 1938, 1945, p. 411; Freytet and Plaziat, 1965). I t was first noted by Lamarck and De Candolle (1806) as Batrachospermurn haematites, probably the first record of a freshwater stromatolite-building organism. Marine species of Rivularia generally do not form organosedimentary structures.

Another marine rivulariacean, Gardnerula coryrnbosa (Harvey) De Toni

tzing) Gomont (Griininger, 1965; Golubic, 1967).

1cm


(lo), has bundled and ramified filaments which grow upright and radially. Sediment is trapped between the bundles producing up to 80cm high obtusely conical stromatolites in Shark Bay (Fig. Id). This species was first described by Harvey (1857) as Microcoleus corymbosus “near high water mark at Key West” with the remark “half sunk in the mud”-probably the first

Calothrix crustacea Thuret (ll), a rivulariacean common in marine intertidal (or periodically air-exposed) environments develops in later successional stages of

~~~ y \ : “ k established algal mats (Walter et al., 1973). It forms a fortifying support to the structure or a smooth leathery coating from which wind-blown sand particles are deflected. In this sense Calothrix protects the mat from

In calcareous freshwater springs green lithified crusts are formed by radiating branched thalli of the chaeto- phoracean green alga Gongrosira incrustans (Reinsch) Schmidle (12; Golubic and Fischer, 1976), which are ex- ternally similar to crusts of Oocardium stratum (Golubic and Marcenko, 1965; Golubic , 1967, p. 123).

While the radial and upright filament arrangement coincides with the direction of sediment accretion and is convenient for “escape”, the prostrate growth of filament networks provides a better stabilizing agent of loose sediments. Several adaptations evolved as variations on this growth pattern. Simple loose fabrics of interwoven branched filaments serve as a base for specific calcification in subaerial cave algae Geitleria calcarea Friedmann (13) and Scy tonema julianum (Kutzing) Meneghini (Friedmann, 1955; Bourrelly and Dupuy, 1973). In microenvironments within structural cavities of coral skeletons, a similar calcification pattern was described in the endolithic green alga Ostreobium quekettii Bomet et Flahault (Schroeder, 1972).

Calcified networks of Scytonema myochrous (Dillwyn) Agardh (14) combine advantages of both horizontal and vertical growth (Monty, 1967) fortifying the structure in a trabecular fashion which may obscure the stromatolitic lamination. Wellexpressed lamination was described in the related Scytonema mirabile (Dillwyn) Bomet forma zonata by Geitler (1932) while Thunmark

(1926) observed an annual rhythmicity in layers of S. crustaceum Agardh. Several poorly known marine species of Scytonema form stromatolites by particle entrapment in Shark Bay (Logan et al., 1974) and the Bahamas (Monty, 1967,1972).

I

lo

ion of a marine Recent stromatolite. 1 1 a ,

! .I . 1 ,: -,, ’.. 1 :I ; f 0

m

: . :,

s

122 S. GOLUBIC

icm. A similar combination of horizontal and vertical growth in stromatolite formation evolved in the red alga Rhodothamniella floridulu (Dillwyn) J. Feldmann (Hommeril and Rioult, 1965) and in the siphonalean green alga Caulerpa fastigiata Montagne (15) (Golubic, 1973a). Vuucheria geminata (Vaucher) De Candolle (16) is a versatile type of freshwater stromatolite builder. Its

non-septate tubular filaments form hydrodynamically shaped calcified cushions in travertine deposits (Emig, 1971 ; Wallner, 1934a; Griininger, 1965).

MICROBIAL DESTRUCTION OF STROMATOLITES

Like any large-scale ecosystem, e.g. a tropical forest, an established mat has an increased capacity to maintain constant and optimized conditions and to support a diversified interstitial microbiota not directly involved in stromatolite construction. The overall direction of successional changes within an algal mat leads from predominant photosynthesis, producing organic matter, to predominant decomposition, degrading that organic matter into inorganic compounds. Little is known about the biology and geochemistry of microbial degradation of algal mats which may proceed either aerobically or anaerobically and probably involves several successional steps in both cases.

Aerobic decomposition of algal mats in the marine intertidal zone of the Persian Gulf is nearly complete leaving a soil-like sediment frequently stained reddish by oxidized iron compounds. Under such natural conditions, where the rates of aerobic decomposition follow closely the rates of primary production, only thin algal mats are maintained and the mat material is destroyed as fast as it is formed (Fig. 2 4 .

Anaerobic decomposition within the same coastal environment is found in algal mats of stagnating pools and slowly draining tidal channels. I t is often incomplete leaving a peat-like organic-rich sediment normally stained black by ferrous sulfide (Fig. 2c). Algal peats with partially preserved cellular structures and pigmentation have been found after 8000 years of burial beneath the Abu Dhabi sabkha (Golubic and Park, in prep.).

Under the conditions of comparable rates of primary production and sediment entrapment, it is the ratio of aerobic (complete) to anaerobic (incomplete) decomposition which largely determines the overall (organic + inorganic) rate of sediment accumulation. Along flat protected coasts of the Persian Gulf at Abu Dhabi, pools and depressions which tend toward stag- nation and anaerobiosis have faster sediment-filling rates than the well- drained elevations. The ultimate consequence is a levelling of the entire landscape and formation of sabkha plains (Golubic and Park, 1973).

Anaerobic conditions, as evidenced by the presence of FeS, free H2S and


of anaerobic purple sulphur bacteria, normally underlie the areas of intensive respiration, and result from depletion of oxygen in the immediate environment. Further aerobic decomposition depends then on an efficient aeration which is able to replenish the used oxygen. In well-drained intertidal mats anaerobiosis is a local and temporary condition, but it becomes permanent in stagnant pools. Oxygen deprivation within the euphotic zone favors dominance of anaerobic bacterial photosynthesis (purple-pink layer) although many cyanophytes show preference for low oxygen tensions (Stewart and Pearson, 1970; Golubic, 1973a, p. 458).

Changes in microenvironmental chemistry caused by decomposition of algal mats also influence their mineral content. Release of ammonia by anaerobic bacterial decomposition, under high pH, within marine algal mats, may cause CaCO, precipitation (Greenfield, 1963; Dalrymple, 1965). How- ever, pH is usually lowered as a general consequence of decomposition processes in both marine and freshwater environments, a condition under which available carbonate is taken into solution. Localized carbonate dissolution beneath algal crusts in lakes removes the base and may cause a collapse of the entire structure and thus limit the growth of stromatolites (Golubic, 1969a). This effect is reduced when high carbonate deposition rates (relative to organic production), or an exchange of interstitial waters dilutes the res- piratory products. Such conditions are exemplified in submerged bioherms of the Green Lake, New York (Dean and Eggleston, in prep.) and calcareous nodules of the Rhine river and Conestoga Creek (Golubic, 1973a; Golubic and Fischer, 1976). A new generation of the Entophysalis mat grown over hard lithified stromatolites of Shark Bay passes through the same stages of production and subsequent decomposition of organic matter (Fig. le) . A partial dissolution effect on the sediment accumulated beneath the growing mat is demonstrated by secondary softening and disappearance of the now buried hard crust. This mechanism may be at least in part responsible for secondary obliteration of stromatolite lamination (S. Golubic and S.M. Awramik, unpublished). Ground-water when recharged with dissolved carbonate may deposit it elsewhere as soon as the partial pressure of surrounding COz is lowered. Cave sinters within travertine deposits in freshwater (Golubic, 1969a) and underground aragonite crust formation in marine lagoonal environments (Shinn, 1970), probably have this origin (Golubic, 1973a, pp. 447, 457).

17 Finally, there is a specialized microflora of boring endolithic algae and fungi which actively degrades carbonate substrates under illuminated surface conditions. A dark coating over lithified stromatolitic heads of Shark Bay, also called the “film mat” (Hoffman, Ch. 6.1) is composed of the boring cyanophyte Hormathonema uio- laceo-nigrum Ercegovid (17) which does not contribute

to stromatolite construction, but is actively destroying them (Fig. l e , f ) . -

124 S. GOLUBIC

(Golubic et al., 1975; Golubic and Awramik, in prep.). The distinction between constructive, neutral and destructive elements in the stromatolitic microbiota needs to be refined in order to assess their relative contributions.

DISTRIBUTION OF ALGAL MATS

The distribution of stromatolite-building microorganisms is largely determined by ecological factors. The same or very similar taxa may have a cosmopolitan distribution (Golubic and Awramik, 1974). However, subtle differences in climate and local ecology are frequently responsible for significant changes in microbial composition, their interaction within a community,

Fig. 1. Sequence of algal mat and stromatolite types at Carbla Point, Hamelin Pool, Shark Bay, W. Australia. MHT, mean high tide; MLT, mean low tide; loose sediment and beach dunes are coarsely dotted. a = subtidal columnar and branched stromatolites, often internally lithified. Macroscopic algae (e.g. Acetabularia) grow on top of these structures. b = prostrate modification of a. The mat morphology and algal community are the same. c = finely laminated mats and domes dominated by a Schizothrix sp.; d = obtusely conical, mostly unlithified stromatolites of Gprdnerula corym bosa. The internal microstructure reflects a radiating arrangement of algal filaments with zones showing rhythmic growth. e = initially unlithified, and then (probably periodically) lithified stromatolitic heads built by the mamillate mat of Entophysalis m q o r Ercegovid. Only the soft parts are actively growing. Lithified crusts in the upper range of this type (black top of the heads in e and f ) are colonized by a film-mat of epilithic and endolithic (boring) cyanophytes: Entophysalis granulosa and Hormathonema violaceo-nigrum. f = tufted and reticulate mat, with a cbmposition that ranges from an absolute dominance of Lyngbya aestuarii to a vertically differentiated community dominated by Schizothrix splendida. This mat covers the flats between lithified, film-mat-covered stromatolitic heads. g = pink smooth mat dominated by Microcoleus tenerrimus covers flats in the uppermost range of algal mat distribution, between cracked and partially eroded stromatolitic domes (S. Golubic and S.M. Awramik, unpublished).


and in the resulting stromatolite morphology. The zonal distribution of algal mats observed in two marine peritidal environments, of the Persian Gulf and Shark Bay, are selected here to exemplify biological and sedimentological diversification along an ecological gradient of increased exposure to air (Figs. 1 and 2 ) .

The subtidal stromatolites of Shark Bay (Fig. l a , b ) are built by algal mats composed of a variety of eucaryotic as well as procaryotic microorganisms. In the intertidal range, however, algal mats and stromatolites are formed by less diverse, exclusively procaryotic microbial communities. Algal mat types are arranged in zones roughly parallel to the coastline. The relief of stromatolitic structures can modify the distribution of local ecological conditions and cause a more irregular or mozaical distribution of algal mat types (Fig. 1 above, plan view). The mamillate mat of Entophysalis major Ercegovid comprises the most important stromatolite-building mat in the intertidal range (Fig. l e ) . In protected embayments this mat is unlithified and flat.

In the lagoons of Abu Dhabi, Persian Gulf, unlithified laminated biscuits of Phormidium hendersonii are the only subtidal stromatolitic structures (Fig. 2a). The mamillate mat (Fig. 2b) is here the most important stabilizer

Fig. 2. Sequence of algal mat types across the intertidal zone NE of Abu Dhabi, Persian Gulf. MHT, mean high tide; MLT, mean low tide; sediment beneath the mats and loose sediment are coarsely dotted. a = subtidal gelatinous, finely laminated biscuits of Phor- midium hendersonii. b = mamillate mat of Entophysalis major. c = flat “low” mat, often polygonally cracked, represents a vertically differentiated microbial community characterized by Microcoleus chthonoplastes. d = pinnacle “high” mat, also a composite, multi- layered community characterized by Schizothrh spiendida. e = blister mat, an ecological modification of c and d. f = wrinkle mat, a residual modification of d adjacent t o the “sabkha” plain (S. Golubic and R. Park, unpublished).

of loose sediment in the lower intertidal range. Both types are composed of monospecific microbial populations. Algal mat types in the upper intertidal range (Fig. 2 c - f ) are structural and functional modifications of two composite, vertically differentiated microbial communities.

Water-logged depressions and tidal pools and channels are covered by a flat “low” mat (Fig. 2c), which becomes polygonally cracked after prolonged

126 S. GOLUBIC

exposure to air. It is a vertically differentiated, layered community with: (1) a brown surface layer of Lyngbya aestuarii; ( 2 ) a blue-green layer of Micro- coleus chthonoplastes; ( 3 ) a salmon-pink layer of pigmented filamentous bacteria; and (4) a purple-pink layer of purple sulphur bacteria. The base of the mat is anoxic, normally stained black by FeS.

Welldrained elevations are occupied by a pinnacle “high” mat (Fig. 2 d ) . This is also a multilayer community, with the blue-green layer occupied by Schizothrix splendida. A similar community forms the tufted and reticulate mat of Shark Bay (Fig. I f ) .

The blister mat in the upper intertidal range (Fig. 2e) is composed of modified “low” and “high” mat communities re-arranged at a smaller scale. Water-logged depressions between the blisters differentiate into the cclow” mat, and the welldrained blister surfaces into the “high” mat community. Pinnacle formation is suppressed due to desiccation. The wrinkle mat at the top of the sequence (Fig. 2 f ) is a simplified residual modification of the “high” mat community.

The morphology of a Recent stromatolite is a final product of complex biological (structural and functional) and environmental interactions. It may be correlated with the distribution of microbial populations and communities, or with environmental conditions such as energy exposure, water supply and drainage, rates of sedimentation and mineral precipitation.


Chapter 4.2

TAXONOMY OF EXTANT STROMATOLITE-BUILDING CYANOPHYTES

Stjepko Golubic

INTRODUCTION

Algal mats and Recent stromatolites are products of microbial communities composed of cyanophytes, eucaryotic algae, photosynthetic bacteria (thio- rhodo-, athiorhodo-, and chlorobacteria), and various heterotrophic bacteria. Cyanophytes or blue-green algae are the most significant single group engaged in the formation of stromatolites. Historically, they represent the earliest water-photolyzing and 0,-releasing photosynthetic organisms, and they dominated Precambrian environments for billions of years. Ecologically, they occupy today a wide array of niches under extreme environmental conditions where other organisms are less successful or cannot exist at all. Their diversity in terms of ecological niches they occupy, and in terms of the resultant stromatolite morphology, is parallelled by a diversification of taxa (Walter et al., 1973).

Despite difficulties in the application of the species concept to a largely asexually reproducing group of procaryotes, the concept of ecological niche differentiation by species-equivalent taxonomic units (one species = one niche: Hutchinson, 1968) does apply to this group and it underlies the systematic division used here. Classical systematics of the Cyanophyta are largely consistent with this ecological approach to taxonomy (Golubic, 1961b, 1969b), and the outline presented here follows this tradition. The system used is based on the scheme of the French phycologists Thuret (1875), Bornet and Flahault (1886- 1888), and Gomont (1892), as continued and modified later by Tilden (1910) in the U.S.A., Geitler (1932, 1942) in Austria, Fremy (1930, 1934) in France, Elenkin (1936, 1938, 1949) in the U.S.S.R., and Fritsch (1945) in England. More recent contributions by Kosin- skaya (1948), Desikachary (1959), Umezaki (1961), Starmach (1966), and Bourrelly (1970) have also been consulted. These works are recommended for further taxonomic study; unfortunately few are available in English. The classification of Drouet and Daily (1956) and Drouet (1968, 1973) does not withstand critical scrutiny (Stanier et al., 1971; Desikachary, 1973) and is

128

TABLE I

Systematic outline of extant cyanophytes

Nostocaceae I Nodular ia , Anabaena, Nostoc, Cylindrospermum, Aphanizomenon. Scy tonemataceae , Microchaete , Scytonema , Tolypothr ix . Rivular iaceae

Calo thr ix , Dichothr ix . Gardnerula , Rivular ia .

S. GOLUBIC

am C

4 S

?JXE* F

amaao T

0 H

Y B

.ass :der

Family I Genera . . . . . . . . . . . . . . . . . . . . . 1 <oneae

I i roococcales

Chroococcaceae Synechococcus, J o h a n n e s b a p t i s t i a , Gloeothece, Aphanothece Synechocyst is , Merismopedia, Gomphosphaeria, Coelosphaerium Chroococcus. Eucapsis , Gloeocapsa, Aphanocapsa, Microcys t i s .

Chlorogloea, Entophysa l i s , Cyanostylon, Hornathonema, S o l e n t i a . Entophysal idaceae

l a m a ,es iphonales Chamaesiphonaceae I Chamaesiphon, Clast idium, St ichosiphon, Cyanophanon.

leu?ocapsales De rmocarpaceae 1 Chroococcidiopsis , Dermocarpa. Pleurocapsaceae I Siphononemataceae

Myxosarcina, Xenococcus, P leurocapsa , Hydrococcus, Hyel la .

I Siphononema* goneae

Symbols on the left refer to ordinal characters: C = coccoid unicells, reproduction by fission; S 7 sessile unicells, reproduction by exospores; F = filamentous cell arrangements, reproduction by endospores; T = cell organization in trichomes; H = heterocyst formation; B = true branching and pit-like connections.


not recommended. General information on cyanophytes can be found in two recently published books by Carr and Whitton (1973) and by Fogg et al. (1973), respectively.

The brief and necessarily incomplete presentation that follows intends to relate the stromatolite-building microorganisms to the principal lines of cyanophyte generic diversity. It should provide an overall orientation that will facilitate and encourage the use of the specialized literature recommended above for determination of genera and species. Many cyanophyte genera and species are interconnected by transitional forms. These transitions reflect continuous speciation common among procaryotes, but cause difficulties in the separation of taxa. However, taxa may be characterized on the basis of morphological variation of entire populations, and by the clustering of data (Hutchinson, 1968). Modern taxonomic methods, including evaluation of physiological and biochemical properties of organisms, or of their genetic make-up have to date been applied only to a few cyanophytes (Stanier et al., 1971), and at this time cannot be used in determination of taxa in natural populations.

The arrangement of taxa as shown in Table I and in the illustrations of the present paper do not imply phylogenetic relationships, but rather morphological affinities. The outline used seeks maximum clarity while presenting a gradual increase in degree of thallus differentiation and complexity. Categories used here have all been previously applied, although the arrangements vary from author to author. For example, the filamentous forms without ability to differentiate heterocysts and akinetes are identified here as a separate order Oscillatoriales (Elenkin, 1949) rather than as a family within Nostocales (Geitler, 1932).

The phylum (division) Cyanophyta can be conveniently subdivided into two classes: Coccogoneae, or coccoid., unicellular cyanophytes; and Hormo- goneae, or filamentous, multicellular cyanophytes. Both coccoid and filamentous cyanophytes often produce extracellular gelatinous matter of various consistency, that may provide a more or less cohesive matrix for colony formation. When soft and amorphous the gelatinous matter will be referred to as mucilage (which often accompanies gliding movement); when firm and cohesive it will be called gel. Cells of coccoid cyanophytes are often surrounded by siligle or multiple gelatinous envelopes. Although coccoid forms can give rise to filament-like cell series when dividing predominantly in one direction, the true filamentous forms are characterized by closer association of cells in chains called trichomes. Trichomes are often surrounded by simple or layered gelatinous sheaths, and the trichome and sheath together are referred to as filament. Coccoid cyanophytes reproduce by simple cell division (fission) and by exo- and endospores, whereas the principal mode of reproduction in filamentous cyanophytes is by fragmentation of trichomes in short, motile portions called hormogonia (Table I).

130 S. GOLUBIC

CLASS: COCCOGONEAE

Orders or coccoid cyanophytes can be distinguished largely by their patterns of growth, by mode of reproduction, and by differentiation of special reproductive cells. Members of the order Chroococcales comprise single or colonial unicells characterized by regular (with respect to division planes and their alternations) cell division. They form no special reproductive spores. The order Chamaesiphonales contains unicellular or colonial attached forms that reproduce by exospores. The order Pleurocapsales comprises unicellular or colonial forms, with a pseudoparenchymatous or filament-like arrangement of cells, and less regular (less consistent with respect to division planes) cell division. They reproduce by endospores.

Order: Chroococcales (Fig. 1)

The Chroococcales encompass simple and colonial coccoid forms which secrete quantities of extracellular gelatinous matter of varying consistency. Normally, simple cell division provides for the growth and spreading of colonies. When the frequency of cell division is high relative to the rate of cell enlargement, large numbers of very small cells called nannocytes develop. The opposite tendency results in formation of larger cells which may function as resting spores. However, no morphological change or specialized differentiation accompanies the various modes of reproduction in Chroococcales.

Within the Chroococcales the family distinction is based on the symmetry of the colonies. Growth and gel production resulting in radially symmetrical colonies (R ) characterize the family Chroococcaceae, while unidirectional growth of colonies (U) is characteristic of members of the family Entophy- salidaceae. Within the Chroococcaceae diversity of form is expressed mainly by the combination of two characters: the planes of cell division resulting in one-, two- or three-dimensional cell arrangements (ID, 20, or 3D), and the manner of production of gelatinous matter. Cell division in one plane is usually accompanied by cell elongation in the direction of division so that most cells in a population are elliptical or rod-shaped. Genera within the ID-group can be distinguished by the amount and quality of the produced gelatinous matter (none, sparse, layered, diffluent) as follows: Synechococcus, Johannesbaptis- tia (Cyanothrix), Gloeothece and Aphanothece. The 2D-group dividing in two planes also contains aseries of genera which differ in their gel production. Synechocystis, which lacks gel, does not form colonies. Merismopedia produces a small amount of gel which holds flat or irregularly curved plate- shaped colonies together. In Gomphosphaeria the cells divide in two alternate planes following the curvature of a spherical surface, and produce gelatinous stalks in the third plane, toward the center. The stalks branch dichotomously, revealing the pattern of division. In Coelosphaerium the gel is less structured

Aphano thece

Gloeo thece '....-, '

Johannesbap t i s t i a

Synechococcus I Synec

T4icrocys t i s

Coelosphaerium Aphano caps a

roococcus

( S y n e ~ h o c ~ s t i s )

10 pm I - CHROOCOCCALES

Fig. 1. Morphological affinities of coccoid cyanophytes: Chroococcales. Genera similar with respect to gel production are placed at the same level. The sequence from below upward includes genera with no, sparse, layered and diffluent gelatinous envelopes. 1 D, 20 , + and 3 0 = one-, two-, and three-dimensional cell-division patterns; R = radially symmetrical colonies (Chroococcaceae); U = uni- 2 directional colonies with rows of cells (Entophysalidaceae).

132 S. GOLUBIC

and the stalks indistinct. The genera which divide in three planes (30) are indistinguishable from the 20-group if no gel is produced (Synechocystis). The three-dimensionality , however, is evident in the cubical cell arrangement of Eucapsis held together by a small amount of firm and layered gel. Increased gel production in Gloeocapsa shifts this arrangement into a deformed prismatic shape (Golubic, 1965). The sequence of cell divisions, revealed by layered gelatinous envelopes in Eucapsis and Gloeocapsa becomes indistinguishable in Aphanocapsa and Microcystis with their unstructured gelatinous envelopes. While in most Chroococcales cell division proceeds concurrently with volume increase, and the daughter cells “pinch off” from a constricting separation ring (pinching-type), in the genera Johannesbaptistia (10) and Chroococcus (30) the cell division precedes cell enlargement, and the daughter cells remain separated by a flat plane (cleavage-type).

Most members of the family Chroococcaceae are found in marine algal mats (Aphanocapsa, Chroococcus, Gomphosphaeria), in thermal springs (Sy- nechococcus), on wet cliffs, and in freshwater (Gloeothece, Gloeocapsa, etc.).

In the family Entophysalidaceae an asymmetric, predominantly unidirectional gel production determines the shape of the colony irrespective of the planes of cell division. In the genus Chlorogloea the cells are packed closely, within tight e‘nvelopes, and show an indistinct arrangement in rows. Uni- directional shift of cells due to a unidirectional and basipetal gel production becomes more evident in Entophysalis. The thickness and arrangement of envelopes in this genus is comparable with that of Gloeocapsa. Extremely unidirectional production of gel in Cyanostylon and Hormathonema results in displacement of cells on the tips of gelatinous stalks. In Cyanostylon the growth is upward, in Hormathonema downward, penetrating carbonate substrates.

Several members of the family Entophysalidaceae build marine algal mats and recent stromatolites (Entophysalis, Cyanostylon). Some members of this family are boring microorganisms capable of destruction of carbonate stromatoli tic structures (Horma th onema, Solen tia).

Order: Chamaesiphonales (Fig. 2)

The order Chamaesiphonales contains simple and colonial, attached forms which reproduce by specialized exospores. Exospore formation follows asymmetric, external “pinching off” (budding) of conidia-like exospores from a larger cell, the exosporangium. The genus Chamaesiphon, representative of the family, includes solitary and colonial forms. Less common genera Clas- tidium and Stichosiphon are mostly solitary.

The most complex genus of this group, Cyanophanon, and the colonial members of the genus Chamaesiphon form compact crusts and coatings over the rocks in freshwater. They are sometimes encrusted with iron but rarely with CaC03. Other members are epiphytic on algae and aquatic mosses (Chamaesiphon).

Cyanophanon

Democarpa

Chamaesiphon

Chroococcidiopsis

1 Yyxosarcina Siphononema St ichosiphon

C l a s t i d i u m

\ \ J F

CHAMAESIPHONALES PLEUROCAPSALE S

Fig. 2. Morphological affinities of coccoid cyanophytes: Chamaesiphonales and Pleurocapsales. S = sessile unicells, reproduction by exospores (ex, Chamaesiphonales); F = filament-like cell arrangements, reproduction by endospores (en , Pleurocapsales).

134 S . GOLUBIC

Order: Pleurocapsales (Fig. 2 )

Coccoid cyanophytes, ranging from simple unicells to complex colonial forms with filament-like arrangement of cells, that reproduce by endospores, are placed in the order Pleurocapsales. They are characterized by a wide variation in cell sizes and shapes, and by a less consistent cell division pattern with respect to the planes of division. Endospore formation proceeds by multiple successive division within a specialized larger cell, the endosporangium (not homologous with bacterial endospore formation). Endospores are released by bursting of the endosporangium.

In the family Dermocarpaceae (which is often placed in the order Chamae- siphonales on the basis of sessile unicellular habit), the least differentiation of functions occurs in the genus Chroococcidiopsis: all cells may subdivide into 2, 4, 8 or up to 64 endospores. Within the genus Dermocarpa, a differentiation into a stalk cell and an endosporangium takes place.

In the family Pleurocapsaceae the cells are arranged in more or less distinct rows, as a result of cell divisions predominantly in one plane. Ramifications are caused by occasional change in the plane of cell division or by “slipping” of individual cells out of the row. Cell divisions in two and three planes results in pseudoparenchymatous packing in three-dimensional cell aggregates in Myxosarcina, and in formation of crusts in Xenococcus. A more pronounced filamentous cell arrangement and an increase in morphological variation is found in Hydrococcus, Pleurocapsa and Hyella. The cells form closely packed vertical rows, or uni- and multiserial filamentous arrangements with lateral and dichotomous ramifications. The terminal cell of a filament-forming series becomes increasingly differentiated into a larger morphologically distinctive end-cell. Endospores form only in few specialized cells. Hyella has several species which penetrate carbonate substrates.

In the family Siphononemataceae, the filaments form from a series of gloeocapsoid cell aggregates.

Members of the order Pleurocapsales form crustose coatings over rocks in freshwater (Pleurocapsa, Hydrococcus, Siphononema). Other members are marine epilithic and endolithic algae (Hyella), or are epiphytes on larger algae (Dermocarpa, Xenococcus).

CLASS: HORMOGONEAE

Distinction between the orders of filamentous cyanophytes is based largely on the degree df cellular differentiation and/or on branching patterns. Simpler members of the class with little cellular differentiation along the trichomes comprise the order Oscillatoriales. They are also called homocystous filamentous cyanophytes. In the more complex members of the orders Nostocales and Stigonematales, some cells along the trichome differentiate into heterocysts,

BIOLOGY OF STROMATOLITES

Hydrocoleum S c h i z o t h r i x Micro co l e u s

Phormidium Lvngbva

I

Trichodesmium \-’ I I O s c i l l a t o r i a S p i r u l i n a Pseudana iaena

135

4 OSC I 11 ATOR I A1 ES

Fig. 3. Morphological affinities of filamentous cyanophytes: Oscillatoriales. Genera similar with respect to gel production are placed at the same level. The sequence from below upward includes genera without gelatinous sheath, with sheath surrounding single trichomes only, and with sheaths surrounding 1, 2 o r more trichomes. T = cell organization in trichomes.

which are cells specialized for enzymatic fixation of atmospheric nitrogen. These may accordingly be called heterocystous filamentous cyanophytes.

Order: Oscillatoriales (Fig. 3 )

The diversity of the Oscillatoriales can be accommodated within a single family, the Oscillatoriaceae. The cells divide in one plane of division forming unbranched trichomes; all the cells along a trichome retain the ability to divide. The only morphologically distinct cell is the mature end cell, which develops from an undifferentiated terminal cell of a recently fragmented trichome (hormogone). Therefore, a population usually contains all stages in this process of differentiation.

136 S. GOLUBIC

Generic distinction within the Oscillatoriaceae is based on the diversity in formation of sheaths. Oscillatoria, Trichodesmium, Spirulina and Pseud- anabaena do not produce distinct sheaths. However, loose unstructured and diffluent mucilage is sometimes found associated with the trichomes. The genus Phormidium is characterized by a thin, firm, sometimes diffluent sheath, while the sheath of Lyngbya is always thick, clearly visible and sometimes layered. Both these genera contain only one trichome per filament. The genera Schizothrix, Hydrocoleum and Microcoleus form (in addition to single trichomes) filaments with two or more trichomes within a common gelatinous (often layered and uneven) sheath. The presence of an end-cell wall thickening or calyptra distinguishes Hydrocoleum from Schizothrix. In other genera this characteristic is used to separate groups of species within a genus. Bundles of numerous actively motile trichomes characterize Microcoleus.

Members of the order Oscillatoriales frequently form stromatolitic structures in both freshwater and marine environments (Phormidium, Lyngbya, Schizothrix, Hydrocoleum, Microcoleus). Some members form soft algal mats in marine embayments and thermal springs (Oscillatoria, Spirulina). Several members are planktonic in tropical oceans (Trichodesmium) or in freshwater lakes.

Order: Nostocales (Fig. 4)

Diversity within the Nostocales can be observed along a series of increasing differentiation of cells within trichomes. Like the Oscillatoriales the cells divide only in one plane resulting in the formation of uniserial unbranched trichomes. Unlike the Oscillatoriales, the Nostocales have specialized cells: heterocysts, and less frequently akinetes (resting spores). Within the family Nostocaceae all vegetative cells retain the capacity to divide, while members of the families Scytonemataceae and Rivulariaceae possess specifically located “meristematic” zones (zones of intensive cell division) along the trichome. Distribution of these “meristematic” zones as well as heterocysts determine the type and position of trichome breakage and thus the formation of false branching. The heterocyst “anchors” the trichome within the sheath, while the “meristematic” zone causes localized growth pressure. As a consequence, the trichome ruptures and penetrates the sheath forming a false branch. In the Scytonemataceae, the “meristematic” zone is located subapically close to the distal end of a growing trichome, while the location of “meristematic” zones in the Rivulariaceae is basal or intercalary, but always away from the distal end of a growing trichome. The result is the characteristically tapered whip-like trichome end.

In the Nostocaceae both heterocysts and resting spores are formed, and genera are characterized by the location of these differentiated cells. Hetero- cysts are terminal in Anabaenopsis and Cylindrospermum, intercalary in Nodularia, Anabaena, Nostoc and Aphanizomenon. Soft mucilage surrounds

Rivularia Gloeotrichia

Nostoc - Aphanizomenon

100 prn -

Hicrochae te

M2

E 0

NOSTOCALES CI

Fig. 4. Morphological affinities of filamentous cyanophytes: Nostocales. H = heterocyst formation; T = trichomes without differ- 2 entiation in distinct “meristematic” zones (Nostocaceae); M I = “meristematic” zones basal (or intercalary), trichome ends tapered (Rivulariaceae); M2 = “meristematic” zones subapical (Scytonemataceae).

138 S. GOLUBIC

thc trichomes in most genera; only Nostoc produces large amounts of firm gel.

Most Nostocaceae are planktonic and many of them form blooms in freshwater lakes (Anabaena, Cylindrospermum, Aphanizomenon). Few are benthic in freshwaters, or terrestrial (Nostoc) . They sometimes trap sediment but do not form distinct stromatolites; others are minor components of marine algal mats (Nodularia).

The Scytonemataceae have both terminal and intercalary heterocysts, but spore formation is extremely rare. Firm gelatinous sheaths often layered and stained brown occur in all members. In Scytonema intercalary heterocysts with two pores anchor the trichome at several points. Growth at secondary “meristems” which develop between two heterocysts causes bending, formation of loops and finally breaking of the trichome. The two loose ends, each now with a distal “meristematic” zone, continue to grow as false branches. As a consequence, Scytonerna grows horizontal filaments with upright false branches. In Tolypothrix intercalary formation of a one-pore heterocyst determines that the position at which the trichome breaks is immediately on the side away from the pore. A “meristematic” zone which originates here penetrates the sheath forming a single false branch. The growth of Tolypothrix is generally upright in a form resembling a paint-brush.

In the family Rivulariaceae single false-branching at a basal one-pore heterocyst is the rule. Genera of the Rivulariaceae are distinguished by patterns of false branching and by production of gelatinous matter which shapes the colony. Calothrix is unbranched or sparsely branched, while Dichothrix forms cushions of densely branched paint-brush type filaments; neither produces common gelatinous matter between the filaments. Dichotomously branched bundles of filaments inside a common gel is characteristic of the genus Gardnerula. Riuularia and Gloeotrichia have branched filaments completely embedded in semiglobose or globose masses of common gel.

The family Rivulariaceae has several important stromatolite-building species in freshwater (Calothrix, Riuularia) and marine environments (Calothrix, Dichothrix, Garnerula). Other forms are epiphytic or planktonic in freshwater (Gloeotrichia).

Order: Stigonematales (Fig. 5 )

The order Stigonematales represents the highest degree of complexity and morphological differentiation on the procaryotic level of organization. The cells within trichomes have pit-like protoplasmic connections. In addition to transverse cell divisions within trichomes and false branching, longitudinal and oblique cell divisions occur producing true branching and multiseriate filaments (composed of two or more rows of cells). A true branch originates from a cell with three points of connection (pits) with surrounding cells. Both intercalary (two pore) and terminal (one pore) heterocysts occur.


100 urn - C’

Brachy trichia ‘

STIGONEMATALES

Fig. 5. Morphological affinities of filamentous cyanophytes: Stigonematales. B = true branching, pit-like cell connections; @ = true branching starting from trichome loops (Mastigocladaceae); @ = true branching starting from straight trichome portions (Nosto- chopsidaceae, Stigonemataceae).

In the family Mastigocladaceae the branching originates from “humped” or loop-forming trichome portions. The cell at the loop curvature bulges and divides sideways starting a branch (Y -ramification). Mastigocladus produces a variety of growth forms without common gel, forming soft algal mats in thermal springs. Brachy trichia in marine environments forms semiglobose thalli embedded in common gel with peripheral branches tapered into whip- like ends resembling Riuularia. Some forms of Brachytrichia are endolithic within carbonate substrates.

In the family Nostochopsidaceae the branching originates laterally by simple bulging of intercalary cells (T-ramification). The branches are either long and continuous, or short, terminating with a heterocyst. In the extreme case the heterocyst forms from the first cell of a branch and lies laterally on the trichome. Nostochopsis forms thalli embe.dded in common gelatinous masses resembling Nostoc. Mastigocoleus is a marine endolithic organism which penetrates carbonate substrates.

The family Stigonemataceae includes a series of genera with increasing differentiation of thallus and true branching. Hapalosiphon has prostrate, torulose main filaments and upright cylindrical branches. Heterocysts are intercalary. More pronounced heterotrichous thalli characterize Fischerella, with its torulose and multiseriate prostrate filaments and lateral as well as intercalary heterocysts, and upright uniseriate branches with exclusively

140 S. GOLUBIC

intercalary heterocysts. In Stigonema, however, all branches become increasingly multiseriate, and the distinction between prostrate and erect filaments is lost. Multiseriate filaments originate from initially uniseriate trichomes by extensive true branching. Differentiation starts from the actively growing tips and proceeds down the branch, ultimately forming crowded gloeocapsoid aggregations. All branches, however, can be traced through the pit-like connections between the cells. The similarity with coccoid cyanophytes Gloeo- capsa and Siphononema is therefore only superficial. With respect to growth pattern, cell differentiation and organization of the thallus, Stigonema represents the most complex procaryote, comparable to some eucaryotic rhodo- phytes.

Most Stigonemataceae form soft algal mats in freshwater environments, and some are characteristic of bogs (Hapalosiphon). Some species of Stigonema form algal mats on wet cliffs.

Genera with highly differentiated thalli, but without heterocysts are difficult to classify, and some of them may be heterogenous. Species of Plecto- nema, for example, without differentiated “meristematic” zones, have most in common with the Oscillatoriaceae, whereas those species with “meristematic” zones resemble members of the Scytonemataceae. Homoeothrix and Ammatoidea with tapered trichomes are best classified within the family Rivulariaceae, and Geitleria with pit-like cell connections and true ramifications, within the Stigonemataceae.

ACKNOWLEDGEMENTS

The research reported in Chapters 4.1 and 4.2 was supported by Grants GB-25271, GA-31168 and GA-43391 by the National Science Foundation, U.S.A., and by travel grants by the National Academy of Sciences, Foun- dation for Microbiology, American Philosophical Society and National Geo- graphic Society, and by grants GA-1170,13489, PRF-3009-A2 and API-99 to D.J.J. Kinsman, who collaborated in the field research. Manuscripts were critically read by Drs Stanley M. Awramik, Thomas D. Brock, William Doemel, Lynn Margulis, M.D. Muir, Malcolm R. Walter, David J.J. Kinsman, Dorothy Z. Oehler, H. Marchant and Caroline Tropper. I appreciate their help, and assistance in the field by Drs Phillip E. Playford, S.M. Awramik, Victoria Koehler and Mike Gatrall.


Chapter 4.3

ENVIRONMENTAL MICROBIOLOGY OF LIVING STROMATOLITES

Thomas D. Brock

INTRODUCTION

This brief review will consider the environmental limits and tolerances of the blue-green algae (cyanophytes) and associated bacteria, in relation to the limits and tolerances of other algal groups. The information has been derived from a number of ecological studies on blue-green algae which have already been summarized (Brock, 1967a; Brock, 1969b; Brock, 1970; Brock, 1973). The present review will also include some new information not covered in the reviews just cited.

Ultimately, an organism will develop in nature only if all of the environmental conditions necessary for its growth and function are appropriate, and if it can grow at faster rates than other competing organisms. Blue-green algae develop under a wide variety of conditions, and because of this it is commonly assumed that they are tolerant to most environmental extremes. In actuality, for a number of environmental factors, blue-green algae are no more resistant than other algae, and for certain factors they are even less tolerant. There is only one type of habitat known on earth today where bluegreen algae are quite clearly the exclusive oxygen-evolving photosynthetic organisms, and that is in thermal springs of neutral to alkaline pH (Brock, 1967a, 196913, 1970; Castenholz, 1969). In other habitats, even those in which blue-green algae are widely held to be successful, they are never exclusive (or if they are exclusive, it has not been shown that other algae will always be excluded) and often they are not dominant.

Table I presents a brief summary of the environmental limits of the blue- green algae, as compared to other organisms. For some environmental factors, our knowledge is quite limited, and thus the data are only tentative. The uncertain data are indicated by question marks in Table I. In the following brief discussion, I would like to comment on each of the environmental factors. From a consideration of the data in Table I, it may be possible to indicate why it is that bluegreen algae form mats in certain habitats in which eucaryotic algae are absent. Note that not all blue-green algae can grow at the environmental extremes; only a few species will be so tolerant.

142 T.D. BROCK

TABLE I

Environmental limits for growth of the blue-green algae and other microorganisms

Environmental factor

Bluegreen Heterotrophic Eucaryotic algae bacteria algae

Temperature

Upper limit

Lower limit

P H

Upper limit Lower limit

Salinity

Upper limit

Water potential (matric)

Lower limit

Light intensity

Lower limit

70-73'C

freezing

> 10.5 (?) 4-5

- 50 bars (tentative)

2000 lux (tentative)

Anaerobic growth Yes

> 99OC

freezing

> 10.5 (?) < 1.0

- 50 to - 100 bars

56'C (40'C at neutral pH due to competition)

freezing

> 10.5 (?) < 1.0

- 50 bars (tentative)

probably similar

probably no

Data from Brock (1967a, 196913, 1970, 1973), Smith and Brock (1973), Tansey and Brock (1972), and unpublished. Environmental tolerances may be more extreme than the limits for growth. TEMPERATURE

The upper temperature limit for blue-green algal growth has been well established from observations in thermal springs, and there is no reason to believe that a similar upper limit would not exist in other habitats. The contrast here between blue-green and eucaryotic algae is quite striking. Although the upper temperature limit for eucaryotic algae is listed as 56"C, in reality eucaryotic algae are rarely found in nature at temperatures above 40°C. This seems to be because they are unable to compete with blue-green algae for growth in most habitats at temperatures above 40°C. Indeed, the only kinds of habitat in which eucaryotic algae seem to be present significantly at temperatures above 40°C are acidic habitats (pH <4-5) where blue-green algae are completely unable to grow (see below), and eucaryotic algae do not face any competition. It should be emphasized that only a very few species


of blue-green algae are able to grow at high temperature. At temperatures above 60-64"C, only a single blue-green alga, the unicellular Synechococcus lividus, exists. The upper temperature limits for a number of other blue- green algae have been given by Castenholz (1969) and are listed here: Masti- gocladus laminosus (filamentous, with heterocysts) 60-64"C, Oscillatoria okenii (filamentous, no heterocysts) 6OoC, Phormidium laminosum (filamentous, no heterocysts) 57-60°C, Synechococcus minervae (unicellular) 6OoC, Spirulina sp. (filamentous, no heterocysts) 55-60°C, Aphanocapsa thermalis (unicellular) 55'C, Oscillatoria terebriformis (filamentous, no heterocysts) 53OC, Calothrix sp. (filamentous, heterocysts) 52 - 54OC, Pleurocapsa sp. (unicellular to short-filament, colonial) 52- 54OC, Symploca thermalis (filamentous, branched, no heterocysts) 45- 47"C, Synechocystis aquatilis (unicellular) 45-50°C. It can be seen that at temperatures below 60°C, there exists a variety of blue-green algae, of various morphological types. It should also be pointed out that blue-green algae are unable to tolerate, at least in the moist state, temperatures above 70-73'C, and are rapidly killed.

A distinction should be made between ability to grow at a particular temperature, and ability to survive this temperature without growing; In general, organisms are able to tolerate temperatures somewhat higher than those at which they can grow, although the upper growth temperature and the upper temperature for survival are usually very close. Thus, blue-green algae will be killed at temperatures much above 70-73OC, and eucaryotic algae will (in most cases) be killed at temperatures above 40'C. Even a few hours heating may be all that is required to kill an organism. Thus, it seems likely that brief solar heating to temperatures above 4OoC (which can easily occur) will eliminate eucaryotic algae from many habitats which might otherwise be favorable for growth. Thus, solar heating on tidal flats and other shallow- water environments probably is sufficient in many parts of the world to select for blue-green algae. There seems to be no differential sensitivity of blue-green and eucaryotic algae to low temperature, although in many frigid habitats the favored algae seem to be diatoms or other eucaryotes. Mosser and Brock (1976) studied an algal mat developing in the outflow stream from a permanently ice-covered lake in the Beartooth Mountains. The temperature never rose above l 0 C and the sole alga was the coccoid eucaryote Protococcus sp.

The upper pH limit for algae has not yet been determined. Indeed, high pH environments are not common on earth, although many aquatic systems may show rises in pH to greater than 10 in mid-day as a result of high photosynthetic rates leading to a shift in the carbonate equilibrium. There is, however, good evidence that blue-green algae are unable to colonize low pH

144 T.D. BROCK

environments, the lower limit being around 4-5 (Brock, 1973). This seems to be a real evolutionary barrier for blue-green algae, since it is likely that acidic habitats have been available on earth since the atmosphere became oxidizing in the Precambrian, so that there has been sufficient time for evolutionary adaptation to take place. When it is realized that eucaryotic algae not only grow in low pH environments but often grow profusely, it seems clear that acid is a definite limiting factor for blue-green algal growth. Interestingly, mats are formed by eucaryotic algae at low pH (Lynn and Brock, 1969; Doemel and Brock, 1971), although I have never seen laminated ones.

SALINITY AND WATER POTENTIAL

Space does not permit a detailed discussion of the concept of water potential as it relates to growth of living organisms. However, this is an extremely important concept for the understanding of the formation of blue-green algal mats in hypersaline habitats, and the reader is urged to consult the book by Griffin (1972) for a detailed treatment of this subject. Briefly, the availability of moisture to an organism is expressed in terms of water potential, which in essence expresses the suction which an organism must place upon the water in a system in order to take up water. Water potential is expressed in pressure units of negative bars. There are two ways in which availability of moisture to an organism can be restricted: by osmotic and matric means. Osmotic water potential is controlled by the ions dissolved in the water, such as salts and sugars. Salinity is often used to express this effect, but water potential is more accurate, and is preferable in the present context. A solution saturated with sodium chloride (such as a hypersaline lake) has a water potential of about - 300 bars. (Such a solution would have an equilibrium water vapor pressure equivalent to a relative humidity of about 8076.) There are probably blue-green algae which can grow at osmotic potentials equivalent to saturated NaCl (Hof and Fremy, 1933), although the eucaryotic alga Dunaliella salina is almost always dominant in such habitats, and is probably in general more successful (Kirkpatrick, 1934; see also Brock, 1975a). Matric water potential is controlled by adsorption phenomena. In natural systems, clays and other minerals with high absorptive power have the strongest affinity for water, and hold it tightly. As water is added to a dry soil, initially most of the water adsorbs to mineral surfaces and is unavailable, but as more and more water is added, the layers of water on soil minerals become thicker and adsorption forces holding it are weaker. The adsorptive power of a soil (or other surface) for water is expressed as water potential. In most soil systems, microbial growth ceases at matric water potentials around - 50 to - 100 bars, although a few fungi are able to grow at water potentials as low as - 600 bars. For most fungi, response is about


the same whether water potential is varied osmotically or matrically, although this is not necessarily true for other organisms. When living in soil, the alga Cyanidium caldarium (a eucaryote) has a lower limit for photosynthetic activity of about - 30 bars (Smith and Brock, 1973), and the dese j crust alga Microcoleus sp. has a lower limit at -46 bars (see Brock, 1975b). In the case of C. caldarium, both osmotic and matric variation in water potential were studied, and it was found that the alga was more resistant to lowered water potential osmotically than matrically. Lichens, which often live essentially on atmospheric moisture, are able to photosynthesize at about -200 bars (matric), considerably lower than other organisms, but in this case we know (Brock, 1975c) that the fungus partner in the symbiosis partially protects the alga from effects of reduced water potential. The alga Dunaliella salina, which lives in saturated sodium-chloride brines in places such as salterns and Great Salt Lake, is able to grow and function at osmotic water potentials as low as - 300 bars. It has not been possible to study the effect of matric water potential on Dunaliella salina, since the alga required sodium ions for growth. We do know, however, that if the total moisture available in the system is reduced, Dunaliella quickly ceases to function, suggesting that it is quite sensitive to drying phenomena.

The reader may appreciate the importance of the above discussion on the effects of water potential on algae, when it is considered that a prime habitat of stromatolite-forming blue-green algae seems to be shallow, intertidal, hypersaline lagoons in warm climates. In such habitats, drying almost certainly occurs, at least during parts of the tidal cycle, and severe water stresses must be placed on the organisms living there. From the little that we now know about the effects of water potential on algae, it can be suggested that growth of algal mats should be considerably restricted in habitats which are dry for more than a brief period. Much work needs to be done on the water potentials of growing algal mats, in order to define the lower limits of water potential for the organisms which have been successful. Although studies on desert crust algae and other soil algae show that blue-green algae are able to tolerate drying for long periods of time, there is no good evidence that they can grow under conditions of low water potential. Indeed, it is likely that they cannot grow if water potential remains too low for too long a period of time. The degree of aerial exposure during the tidal cycle is probably the critical factor controlling the extent of mat formation.

LIGHT INTENSITY

Little work has been done on the lower limit of light at which blue-green (or any other) algae develop, although this is quite essential in understanding how photosynthetic organisms are to function in thick, self-shaded mats. Our own work in this area (unpublished) has dealt with the light attenuation

146 T.D. BROCK

through blue-green algal mats from thermal springs. In such mats, the blue- green layer is generally no thicker than 3mm (Bauld and Brock, 1973; Doemel and Brock, 1974), and beneath this layer the only photosynthetic organisms are photosynthetic bacteria, which can develop at considerably lower light intensities than the algae. We have measured the attenuation of light through these mats and have found that at the 3-mm level, light has been attenuated by 96-97%. This means that if full sunlight is 70,000 lux (a normal summer intensity in clear air), the light intensity at which blue-green algae would stop photosynthesizing would be about 2100-2800 lux. We showed by use of microscopic autoradiography that photosynthesis by the blue-green algal cells at the bottom of the algal layer was virtually zero (Brock and Brock, 1969b), showing that light is indeed low. In the absence of boring animals and other agencies which might disturb the structure of a mat, it is reasonable to assume that similar mat thicknesses would lead to cessation of photosynthesis in other habitats. Thus, most of the lower portions of very thick algal mats must consist either of dead or dormant cells. The self-shading which develops in the lower portions of the mat converts the system from one which is growing and photosynthesizing to one which is dying and decomposing. Thus, built into the light-driven system of the algal mat are the seeds of its own destruction.

The depth in lakes or the sea reached by light sufficient for net photosynthesis varies from place to place, depending on water turbidity and density of phytoplankton population. According to Sverdrup et al. (1942), the depth at which light is attenuated to about 2500 lux is 100 m in the clearest ocean water, and less than 50m in average ocean water. In lakes, because of much greater turbidity, the 2500-lux depth will in general be even shallower (Hutchinson, 1957).

ANAEROBIC GROWTH

It seems reasonably well established that blue-green algae can grow under conditions which are essentially anaerobic (Stewart and Pearson, 1970), although since blue-green algae form O2 as a result of photosynthesis, it is not clear that a cell is ever completely anaerobic. As recently shown by Weller et al. (1975), the stromatolite-forming blue-green alga Phormidium sp. from Yellowstone Park, not only requires anaerobic conditions for photosynthesis, but is markedly stimulated by low levels of sulfide or other reducing agents. We suggested (Weller et al., 1975) that one reason this alga might form compact nodes and conical structures is that oxygen penetration would be restricted, and the presence of oxygen-consuming bacteria would render the mats essentially anaerobic. This Phormidium has been grown in pure culture for many transfers under anaerobic conditions, and actually seems to develop better under these conditions than in air. Further work on the


anaerobic growth and oxygen sensitivity of blue-green algae would be of great importance. Although I have not seen a satisfactory study, I have the distinct impression that eucaryotic algae do not grow as well anaerobically as do blue-green algae. Because the depths of an algal mat should be anaerobic (our preliminary measurements indicate that this is so), it is quite important to know if bluegreen but not eucaryotic algae can grow anaerobically, since this might explain in part why blue-green algae form compact mats so much more readily than do eucaryotic algae.

PRESERVATION

One final point relates to the whole question of preservation of algal mats. Clearly, this is of prime interest to the geologist, because only a preserved mat has the potentiality of becoming lithified. As noted above, self-shading in the lower portions of a mat converts the system from a growing to a decomposing one. In some of our thermal algal mats we have essentially complete decomposition of organic matter within less than a year. If growth is inhibited by darkening a mat, all material disappears. Under what conditions might decomposition be inhibited? We don’t really know, but can suggest some possible clues. First, it should be emphasized that anaerobic conditions do not, of themselves, lead to inhibition of decomposition. Anaerobic food chains leading to complete conversion of organic matter to methane are well known (Water Pollution Control Federation, 1968), and should function in the types of habitats in which algal mats are growing. There must be some other reason why decomposition processes cease. It seems to me most likely that reduced water potential leads to inhibition of decomposition, and as we have noted, there are lower limits of water potential below which microbial life is not possible. As far as I know, low water potential is the only environmental factor which can completely inhibit microbial activity under natural conditions. Reduced water potential is most likely to arise due to drying arising as a result of changes in sea level or changes in tidal channels due to storm erosion, placing algal mats high and dry. This is clearly an area where considerable work is needed.

CONCLUSION

I would like to conclude this brief paper with a plea for much more extensive ecological study of blue-green algal mats in natural situations. It is important to know not only tlie environmental conditions present within mats, but it is also necessary to know if the organisms are actually growing under the measured environmental conditions, or are merely surviving. The techniques for measuring algal activity in nature are available (see Chapter

148 T.D. BROCK

2.3), and can readily be applied to new situations. Further, it is necessary to study the functions of algal mats throughout the seasons of the year, since temperature, light, salinity, and other parameters can vary markedly, and observation during one season will not provide a total picture. Even a few days of unusual warming or unusual drying may be enough to kill off a population and completely change the community structure. This is especially important in marine habitats, because of the marked seasonal changes due to tidal variations, as well as because of the unpredictable but often drastic influence of storms. Algal mats are living, dynamic systems, and can only be understood if they are studied as such.

ACKNOWLEDGEMENTS

Research of the author in this field has been supported by the National Science Foundation, U.S.A. (GB-35046).

4. BIOLOGY OF STROMATOLITES -

Chapter 4 .4

EVOLUTIONARY PROCESSES IN THE FORMATION OF STROMATOLITES

S.M. Awramik, L. Margulis and E.S. Barghoorn

THE STROMATOLITE HABIT

The algal mat as precursor

To understand the evolution of stromatolite-building organisms it is necessary to understand the processes operating in the formation of microbial mats, the precursors of stromatolites. The mat is the actively accreting surface or zone in which the organisms metabolize, grow, and reproduce. Stromatolites are organosedimentary structures produced by sediment trapping, binding, and/or precipitation resulting from metabolic activity and growth of organisms, primarily blue-green algae. These structures may suffer subsequent alteration by degradative biological processes, erosion, diagenesis, compaction and tectonism.

Since this volume contains much of the background information on biological and sedimentary processes leading to stromatolite formation and preservation, we shall not elaborate further here; we refer the reader to Chapters by Golubic (4.1) and Brock (4.3) on the biology and ecology of extant stromatolite-building organisms and Monty (5.1) on fabric.

Stromatolites must be recognized as products of complex microbial communities, generally composed of several to many species of microorganisms interacting with a large number of environmental factors (Golubic, Ch. 4.1; Walter et al., 1973). The study of the change of stromatolitic structure as a function of geological time is not at all the same sort of problem as for example the study through time of silicious diatom tests or gastropod shells. Whereas direct inferences concerning the evolutionary trends within putative monophyletic groups can be made for the latter, such evolutionary inferences are inapplicable to stromatolites. The study of the evolution of stromatolite-building organisms finds an analogy in terms of evolutionary processes to the evolution of forest-forming or reef-forming organisms. Selection for the stromatolite-building habit has occurred in several independent groups of interacting organisms at different times and places.

There are two approaches to reconstruction of the microbial communities

150 S.M. AWRAMIK ET AL.

that actually build stromatolites : the ontological and the micropaleonto- logical. The former utilizes direct field observation of the ecology of the microbial communities and has been amply reviewed (Monty, 1965a; Golubic, 1973a, and Ch. 4.1). These studies have shown that organisms directly and regularly associated with algal mats and stromatolites may not necessarily be stromatolite-building organisms, but in many cases are endolithic and destructive. The problems of interpretation of extant algal mats and stromatolites are multiplied many-fold when extrapolations to ancient microbial communities are made. The potential for preservation of organic matter in microbial mats varies from nearly zero, as is the case where sediments are trapped and bound by the organisms without precipitation of mineral matter, to great, where chemical precipitation accompanies the sedimentary processes. This is a further reason that attempts to derive evolutionary information concerning the organisms composing the communities may be premature.

Evolutionary relationships among organisms associated with mat-building habits

Mat-building organisms are notably restricted to only a few major groups; they have appeared in distantly related groups that share only non-stromatolite-building ancestors. In Table I of Golubic (Ch. 4.1) the genera of mat- and stromatolite-associated extant organisms are listed. They are assigned to higher taxa within the framework of the expansion of Whittaker’s five king- dom classification (see Margulis, 1974a). Here, the names of these organisms are placed on a phylogenetic diagram indicating the relative time of appearance (Fig. 1): e.g., the most advanced are those that appear most recently. The criteria for relatedness and “advancement” from the metabolic point of view have been discussed extensively (Margulis, 1970, 1971). The ability to form mats and stromatolites is extremely ancient and is not limited to photosynthetic microorganisms.

ORIGIN OF STROMATOLITES

Mat surfaces: modified benthos

Why did the stromatolite habit evolve and persist so early in the history of life? The oldest traces of life (Barghoorn and Schopf, 1963; Schopf and Barghoorn, 1967; Engel et al., 1968) known from the Precambrian fossil record are the unicellular blue-green algal and microbial forms of the Swaziland Sequence older than 3,360 million years (Van Niekerk and Burger, 1969) and the Bulawayan stromatolites (ca. 3,000 million years; Bond et al., 1973).


BLUE GREEN

Manganese nodule ,,'- forrninq boclerio ,I'

MONERANS (prokoryofes)

KEY: PHYLA

Gmem *hKn burM Stmnwtd~tes / * w r P /

<> Ploslid cdors

G Green R Red Y Yellow

_. .-

\

Fig. 1. Phylogeny of stromatolite-building organisms. (See Golubic (Ch. 4 .1) for description of these genera. See Margulis (1970) for the endosymbiotic view of the origin of plastids upon which this phylogeny is based.)

Microbial photoautotrophs occur in both benthic or planktonic habitats. The benthic habitat seems to be the more primitive for well-recorded groups of organisms. Today, in the modern seas there is only one genus (Trichodes- mium) of blue-green algae with members that are planktonic throughout their life cycles (S. Golubic, personal communication, 1974); a similar situation may have existed in the Precambrian. I t is possible then that benthic blue-green algae evolved before planktonic forms.


The stromatolitic habit may be considered a modified benthic habit: life at or near the sediment-fluid interface. This habit apparently was favored by natural selection in the Precambrian whether or not the first autotrophs were benthic or planktonic. What are the relative advantages of benthic and planktonic habits? A planktonic, photoautotrophic microorganism must live near the surface of the water in order to photosynthesize. If short wavelength ultraviolet radiation were reaching the hydrosphere in the Early Pre- cambrian, in the absence of other UV-shielding mechanisms, those microorganisms that could not maintain a sufficient column of water overhead to filter out the radiation would be disadvantaged. Alternatively, the production of ozone by the photodissociation of water to H2 and 0, followed by the formation of ozone in the upper atmosphere, may have shielded the biosphere in the Early and Middle Precambrian. Organisms forced to sink below the photic zone for long periods of time would likewise be at a disadvantage. There might have been strong selection pressures for organisms able to maintain neutral buoyancy within the photic zone or able to change density to either sink or rise depending on conditions.

In the presence of turbulence, benthic organisms would require a means of attachment or adherence to the substrate. Possible solutions to this problem include: the production of a holdfast or other sort of adhesive allowing the organism to “glue” itself to the bottom; adaptations allowing the organism to live within sediment; and adaptations allowing organisms to live within an environment favoring mineral precipitation. In the absence of turbulence, such microorganisms require densities greater than that of the fluid medium. Like planktonic photoautotrophs, benthic photosynthesizers must live within the photic zone.

Benthic microorganisms living within the sediment or in active regions of mineral precipitation are faced with the additional problem of remaining at or near the sediment-mineral-fluid interface. Evolution of gliding motility in Precambrian blue-green algae may have been related to emergence from the sediment. In Recent algal mats blue-green algae have solved this problem through natural selection by growing or moving at rates that are as fast or faster than the rate of sedimentation or mineral precipitation.

The stromatolitic habit may have been of greater selective advantage over far larger areas of the earth in Early and Middle Precambrian than it is today. Benthic microorganisms living within the sediment might have been shielded from radiation prior to the evolution of the ability to form pigmented sheaths in which the genetic potential for intensity of pigmentation could be released by the environment. Adaptations to life in the protected sediment may even have preceded the evolution of certain photosensitive enzymatic DNA repair mechanisms. Indeed, the intertidal t o supratidal region was occupied early in the history of stromatolites as evidenced by the 2,300 m.y. old laminated desiccation cracked stromatolites in the Transvaal Dolomite of South Africa (Truswell and Eriksson, 1972).


With stabilization of continental masses at the end of the Archean, stromatolites became widely distributed in Proterozoic basins and shallow waters. We can deduce that the microbial mat habit was of wide selective advantage and geochemically highly amenable to preservation at many localities on the Precambrian earth.

The sub tidal environment

Subtidal stromatolite-building organisms are known from the Recent (Bermuda: Gebelein, 1969; the Bahamas: Neumann et al., 1970; Hamelin Pool, Australia: P.E. Playford, personal communication, 1973). However, these communities apparently are rarer than their intertidal counterparts and tend to be diverse and composed of highly specialized species. The manganese-nodule stromatolites built by oxidizing heterotrophic rod-shaped and chlamydia bacteria are also subtidal. Because of the diversity of the organisms making up these communities it is likely that the subtidal environment was the locus of the evolution of stromatolite-generating microorganisms. The more stable subtidal environment was probably more conducive to growth and development than the intertidal environment during the Early and Middle Precambrian.

The intertidal invasion

The microorganisms which built intertidal stromatolites would have had to evolve additional mechanisms to cope with a periodically exposed (subaerial) environment. In addition to shorter wave lengths, high-intensity visible light may have influenced the distribution of cyanophytes and thereby stromatolites in the Precambrian. High-intensity visible light, in the presence of oxygen, decomposes chlorophyll and carotenoids and produces potentially lethal epoxide compounds in the photosynthetic apparatus of cells (Krinsky, 1966; Leff and Krinsky, 1967). In fact, ultraviolet radiation may have been less a limiting factor in the distribution of cyanophytes in the Precambrian than visible light; blue-green algae are notoriously resistant to ultraviolet radiation. The formation of pigments within the sheaths of blue- green algae was presumably an adaptation favored through natural selection to cope with intense solar radiation, both ultraviolet and visible light. Kylin (1927, 1937) called this material “scytonemine”*. Indeed, observations on algal mats have shown that during summer months, the sheaths become stained with a dark brown to black pigment on the upper surfaces while

* In a recent repeat of Kylin’s extraction from Lyngbya algal-mat sheath material, a strong short UV absorption feature was found. This feature turned out to be due to the presence of nitrate and nitrite salts in the medium. In subsequent experiments ions of these simple salts were shown to protect Lyngbya from lethal doses of UV (Rambler et al., 1975).

TABLE I

Microorganisms reported to be associated with fossil stromatolites

Kingdoma Phyluma Subphyluma Class Phylumb No. of Reference' Comments (as reported) genera

Monera unknown Eubacteria

Cyanophyta aerobic photosynthesizers

Cyanoph yta

Protistad Chlorophyta Mitotica

Chrysophyta

Dinophyta

Rhodophyta

Rhodophyta

Fungi Zygomycota (?)

unknown unknown 4 , 5 , 9 impossible to decide on the basis of morphology to which of the dozen or so eubacterial phyla these belong

Coccogoneae Cyanophyta 7 1 , 2 , 3,5, 6, 8, 9, 10, - 11,12

8 , 9 , 10, 11 Nostocaceae, and Rivulariaceae represented

Hormogoneae Cyanophyta 18 1, 2, 3, 5, 6 Families Oscillatoriaceae,

Chlorococcales Chlorophyta 1 3, 5, 9, 10 debatable morphological

- Chrysophyta 1 8 debatable morphological

- Pyrrhophyta 2 5 Dinoflagellate-like fossil

criteria (Awramik et al., 1972)

criteria (Margulis, 1974b)

Zosterosphaera tripunctata Schopf; if so, probably planktonic Conchocelis-like filament from [%1

isolated algal tetrad; dubious & criteria for taxonomic identification (see Margulis, 197413) 2 Cambrian stromatolites k

Bangiales Rhodophyta 2

Florideales Rhodophyta 1

7

9

Eumycophyta 1 195 i

Eomycetopsis, may be cyano- 5 phyte sheath; oldest unques- m tioned fungi from the Ordovi- cian (Elias, 1966; see Tiffney and Barghoorn, 1974)

a According to Whittaker (1969); for modified Whittaker classification see Margulis (1974a). U r 0 0 4

-I v1 H v 0

5 0

Phylum as reported in the literature for the microorganisms. ' Numbers refer t o the following major references: (1) Tyler and Barghoorn, 1954; (2) Barghoorn and Tyler, 1965; (3) Barghoorn

and Schopf, 1965; (4) Schopf et al., 1965; (5) Schopf, 1968; (6) Hofman and Jackson, 1969; (7) De Meijer, 1969; (8) Licari,

The claims of Middle and Late Precambrian eukaryotes are of utmost importance for the interpretation of the history of life. Their reported first appearance in stromatolitic cherts may mean that the algal mat was the site of major evolutionary events in the Precambrian. The possibility that there is no well-documented eukaryote fossil record prior t o the Ediacaran fauna of the Vendian has been raised (Margulis, 1974b). On the basis of post mortem cell alteration and preliminary artificial silicification studies, caution in the interpretation of the microstructures preserved in the record has been advocated (principally through the work of Golubic; see also Awramik et al., 1972; Barghoorn, 1974). Thus these identifications are tentative, as our comments

1971; (9) Schopf and Blacic, 1971; (10) Licari and Cloud, 1972; (11) Nagy, 1974; (12) Hofmann, 1974.

indicate. L! m


during winter the sheaths are pale brown to colorless (Golubic and Awramik, 1974). In addition to solar radiation, intertidal microorganisms must also be able to tolerate periods of desiccation. It is not known how cyanophytes are protected from the effects of water loss but it is possibly by metabolic shunts that favor the relative production of nonpolar lipid compounds; all intertidal mat-building blue-green algae produce gels which encapsulate and protect them during periods of desiccation.

THE STROMATOLITE RECORD

Microorganisms in ancient stromatolites (Table I)

,The oldest known stromatolites are from the Huntsman Limestone (Bulawayan Group) in Rhodesia (Macgregor, 1941), with an age probably in excess of 3,000 m.y. (Bond et al., 1973). Microfossils have been reported from the Bulawayan graphitic rocks (Oberlies and Prashnowsky, 1968) but we believe that they are not syngenetic with the rocks. Schopf et al. (1971) did not find any microfossils in their investigation of the Rhodesian sediments but concluded that filamentous blue-green algae had evolved by Bulawayan time based on the presence of stromatolitic laminations, gross morphology and stable carbon-isotopic composition. However, unicellular algae (Precambrian Belcher Island stromatolites : Hofmann, 1974; Eocene Green River stromatolites: Bradley, 1929a) and bacteria (Monty, 1973a; Doemel and Brock, 1974) also can produce laminated stromatolites. Recent- ly Walter et al. (1972) suggested that the Bulawayan and other Archean stromatolites may be of bacterial rather than algal origin. Clearly, the Bulawayan stromatolites do not offer direct evidence identifying the microorganisms responsible for their formation, but they are indicative of the evolution of the stromatolitic habit early in the history of life.

Nagy (1974) in a preliminary report described some well-preserved microfossils, morphologically comparable to blue-green algae, from stromatolites of the Transvaal Dolomite in South Africa (approximately 2,300 m.y. old). She assigned these micro fossils to the Families Chroococcaceae (Cocco- goneae), Nostocaceae and Rivulariaceae (Hormogoneae; see Golubic, Ch. 4.2). Similar fossils also from the Transvaal Dolomite were described by MacGregor et al(1974).

In a study of fossiliferous sediments that may bridge the gap between the 3,360 m.y. old Fig Tree-Onverwacht and the 2 b.y. old Gunflint rocks, MacGregor et al. (1974) recently reported an assemblage of columnar intertidal and subtidal stromatolites from the Transvaal Dolomite (Boetsap River section). This is probably the most diverse pre-Riphean assemblage of stromatolites described to date; scanning electron microscopic studies revealed filamentous fossil microorganisms, perhaps oscillatoriaceans and rivulariaceans.


However, the oldest well-studied fossil locality that contains both stromatolites and well-preserved microfossils is the 1,600-2,000 m.y. old Gunflint Iron Formation, Lake Superior region, Canada (Tyler and Barghoorn, 1954; Barghoorn and Tyler, 1965). The Gunflint contains a wide diversity of stromatolite forms (Hofmann, 1969b). From these stromatolites, eighteen morphotypes* of microfossils have been observed, but not all were responsible for stromatolite formation (Awramik and Barghoorn, 1975). Most, but significantly not all, of the microorganisms have a presumed blue- green algal affinity referable to both the Coccogoneae and the Hormogoneae (Licari and Cloud, 1968). No representatives are known of the Rivulariaceae or of the more morphologically complex blue-greens, the Stigonematales. Unlike other microfossiliferous chert& the Gunflint stromatolites contain at least ten morphotypes of uncertain affinities.

Stromatolitic cherts older than 1,800 m.y. on the Belcher Islands, Hudson Bay, Canada, contain another diverse assemblage of microfossils (Hofmann and Jackson, 1969; Hofmann, 1974). Nine morphotypes have been described including eubacteria and some coccoid and possibly oscillatoriacean blue- green algae. Recently Hofmann (1974) described some algae from the Belcher Island assemblage that are morphologically indistinguishable from Entophysalis (Coccogoneae), thus recording the first appearance of the Entophysalidaceae which are today common mat-building cyanophytes.

The Paradise Creek Formation in Queensland, Australia (approximately 1,600 m.y.) also has five or six types of coccoid and oscillatorian-like algae preserved in stromatolites (Licari and Cloud, 1972).

Licari (1971) described nine different stromatolite-building morphotypes from the Beck Springs Dolomite, California, U.S.A. (1,200--1,400 m.y. old). These include four “genera” which have supposed morphological similarities with extant eukaryotic algae and pratists (chlorophytes and possible chryso- phytes). The cyanophytes, not very diverse in this assemblage, are represented by only four species, both filamentous and coccoid. If the identification of eukaryotes in the Beck Springs is correct (Licari, 1971), this is the earliest evidence of eukaryotes in the fossil record and the earliest evidence for the participation of eukaryotic algae in the stromatolite-building biocoenose. (See Margulis, 1974b, for disagreement with this interpretation.) Other reports of putative eukaryotes with organelles have appeared (Hofmann and Jackson, 1969; Barghoorn, 1971; Edhorn, 1973; Muir, 1974). However, pseudoorganelles can be produced in several ways in lysing prokaryotes (Awramik et al., 1972), rendering such interpretations inconclusive.

The approximately 900 m.y. old Bitter Springs Formation of central Australia contains the most diverse stromatolitic microflora described (Barghoorn and Schopf, 1965; Schopf, 1968; Schopf and Blacic, 1971).

* Morphotype: fossil form distinctive enough to be considered representative of the morphology; this may or may not correspond to fossil species.


Fifty different species of microorganisms have been described, 38 of which have blue-green algal affinities. Cyanophytic diversity in extant mats commonly ranges from one or two to about eight dominant species (Golubic, 1973a; Walter et al., 1973); however, as many as about 65 species have been found- in some Recent mats from the Persian Gulf (Golubic, unpublished).

The Phanerozoic record of stromatolite-building microorganisms is meagre by comparison with the Precambrian record. De Meijer (1969) described unicellular and filamentous blue-green algae in addition to a filamentous red alga from Lower Cambrian stromatolites in Spain. In addition, during the Early Cambrian, calcareous algae (such as codiaceans and Giruanella) began to take part in the construction of stromatolites (Johnson, 1966; also see Ch. 4.1 by Golubic). In Carboniferous stromatolites De Meijer found members of the Scytonemataceae, Oscillatoriaceae, and Chroococcaceae (all families of Cyanophyta) and some members of the protist phyla Rhodophyta, Xanthophyta and Chlorophyta. Bradley (1929a) described spherical, unicellular organisms about 100 pm in diameter (Chlorellopsis coloniata Reis) from the Eocene Green River stromatolites of Wyoming. Because of their large size, he compared the cells to the extant alga Chlorococcum infusionurn (Shrank) Meneghini (Chlorophyta).

Macrostructure of stromatolites

Macrostructure or gross morphology of stromatolites usually refers to features measuring from centimeters to meters. Very little useful information on the evolution of stromatolite-building microorganisms has been obtained from the study of the macrostructure. Not very much is known about the biological constraints on stromatolite morphology. 'In general we can not relate size, styles of branching and other macrostructural details to a specific community or to a specific environment. However, Walter et al. (1973) in studies of Coroong (S. Australia) stromatolites found that the shapes of three types of stromatolites to a large extent were determined by the dominant species of organisms. The environment acted both by selection of the organism and also by directly shaping these Recent stromatolites. Certain features of stromatolite morphology are temporally restricted and thus may be related to evolution of the biosphere, hydrosphere, and atmosphere. The following generalizations seem valid : stromatolites in the Precam- brian are generally larger than those found in the Phanerozoic (Monty, 197313); Precambrian and Early Paleozoic columnar stromatolites are commonly narrow,.tall, erect, and branch while those from younger rocks are commonly broader and unbranched (Walter, 1972a) ; columnar stromatolites increase in diversity from the pre-Riphean to Middle-Late Riphean and decrease in diversity during the Late Riphean, Vendian, and Cambrian (Awramik, 1971b). The significance of these observations for the evolution of stromatolite-building organisms is not known. They may eventually be useful for stratigraphic and environmental inferences.


Microstructure of stromatolites

Microstructure refers to features less than a few centimeters in dimension and includes the relief along a single growth surface, grain to grain relationships, nature of branching, and the distribution of organic and mineral matter (Gebelein, 1974). Far more than macrostructure, microstructure appears to be biologically controlled (Monty, 1967; Golubic, 1973a; Hoffman et al., 1972; Gebelein, 1974). The interpretations of distinctive microstructure are difficult and the determination of the algae that were responsible for stromatolite construction is still not possible. Gebelein con- tends, however, that the microstructure is a function of the dominant taxa of blue-green algae in the community. For instance, a community dominated by oscillatoriaceans is thought to produce an algal sediment with fine horizontal or domal laminae which are penetrated by a vertical filament pattern.

Some regular changes in the microstructure of stromatolites with time have been observed. Komar et al. (1965a) observed that the ratio of the thicknesses of the dark laminae over those of the light laminae in subgroups of Conophyton increases from the Early Riphean to Late Riphean, although H.J. Hofman (personal communication, 1974) feels that the laminae have been too altered diagenetically to justify this conclusion. Assuming that the dark laminae represent original organic matter, Walter (1972a) suggests that the observed increases in thickness may reflect the evolution of algae with thicker filaments. A vermiform microstructure first appears in Vendian stromatolites and is strongly developed in some Cambrian forms (Walter, 1972a). Walter (1972a) considers that this may reflect the evolution and wide participation in stromatolite construction of coarse filamentous cyanophytes.

No relation has yet been demonstrated between the macrostructure of a stromatolite and its microstructure. Some macrostructural features may be entirely independent of the specific microbiota and controlled solely by environmental factors. It seems evident to us that the final form a stromatolite attains depends on an integration of biological and environmental factors.

EVOLUTION OF MICROORGANISMS IN STROMATOLITES

The evolutionary history of stromatolite-building organisms is one of generally increasing complexity in community structure through time. Stromatolites, like reefs, have had many different organisms participating in their construction (Newell, 1971). Blue-green algae and bacteria, however, appear to be the dominant forms in stromatolites of all ages while eukaryotic organisms occur on occasion.

The filamentous habit in cyanophytes and bacteria may have greatly


facilitated the building of stromatolites in the Precambrian. Filaments possess greater motility and faster growth rates than coccoid blue-green algae and these abilities enable the organisms to successfully compensate for sediment deposition. With the evolution of the filamentous habit, algal mats probably would have had greater capacity to disperse. The oldest filamentous forms known in the geologic record are the fossil morphotypes assigned to the Nostocaceae and the Rivulariaceae from the Transvaal Dolomite, about 2,300 m.y. old (MacGregor et al., 1974; Nagy, 1974). This assignment is in doubt since tapering of terminal cells, a characteristic of Rivulariaceae, has also been observed because of post mortem loss of turgescence in Oscilla- toriaceae and Nostocaceae (Awramik et al., 1972). It is not certain when the first filamentous forms appeared because sedimentary rocks older than 2,300 m.y. are not well known (Cloud, 197313).

The tectonic instability with accompanying rapid rates of sedimentation which characterizes the Archean (older than 2,600 m.y.) greenstone belts may have provided both an unstable environmental setting for stromatolite formation (although mats may have been widespread) and selective pressure for organisms whose motility and growth eventually kept pace with or exceeded sedimentation (Awramik, 197313). Preservation might have been enhanced by mineral precipitation although Cloud (1973b) suggests that a relatively high partial pressure of C02 in the Archean may have limited carbonate precipitation. However, the presence of the two known occurrences of Archean stromatolites (the Bulawayan and Steep Rock Lake structures) indicate that carbonate precipitation occurred at least on a local level. The reason why no Archean silicified stromatolites have been discovered is not known.

Approximately 2,300 m.y. ago stromatolites, both carbonate and siliceous forms, became widespread. The explanation of the widespread appearance of Proterozoic stromatolites is unknown. It may have involved any one of the following factors or a combination of them: evolution of the filamentous habit, adaptive radiation of aerobic photosynthesis due to the origin of photosystem 11, evolution of other blue-green algal metabolic pathways (such as for nitrogen fixation or for certain mucilages), crustal stability, hydrospheric and atmosphere development permitting widespread precipitation of carbonates and the diversification of particular stromatolite- build- ing groups of filamentous and/or coccoid blue-green algae. The evolution of new forms probably proceeded primarily through sequential mutation followed by patterns of stabilizing and disruptive selection against individuals at the extreme ends of frequency distributions (Lerner, 1968). Although gene transfer via transducing cyanophages may have led to highly fit recombinant organisms, extrapolations from current data about extant forms suggests such prokaryotic recombination mechanisms are relatively rare. The fact that regular sexual reproduction based on meiosis was completely absent in all prokaryotes does not imply that evolutionary rates were


necessarily slower. Rates of evolution are not limited by mutation rates; mutations that themselves alter the mutation rate are known (Cox and Yanofsky, 1969). In morphologically simple prokaryotes variation is more easily detectable a t the physiological level. The appearance of new genotypes with some selective advantage over others in the population may have radi- cally altered the community structure of the biocoenose in which it appeared and eventually modified the morphology of the stromatolites.

SUMMARY AND CONCLUSIONS

Analysis of the evolutionary trends in stromatolite-building organisms is comparable to the analysis of evolutionary trends of reef-building organisms. The stromatolitic habit is notably polyphyletic (Fig. 1). It involves communities of microorganisms interacting with the physical environment. Unfortunately, a major problem in recognizing evolutionary events in blue- green algae is the constraint imposed by the limitation of information preserved in the morphology of filamentous and coccoid forms. Statements that blue-greeen algal evolution is slow and conservative are misleading; they actually refer to the change in fossil morphotypes rather than the evolution of the microorganisms. It is likely that the morphological conservatism of blue-green algae optimally adapted to the stromatolite-building habit is highly evolved in the sense that stability itself is under genetic control. Blue- green algae may have been subject to rapid rates of evolution of new highly fit genotypes; but limited morphology severely limits the phenotypic information that can be deduced by studies of fossil microorganisms.

Nevertheless, an increasing complexity and diversity of blue-green algal fossils can be traced through the Precambrian. Apparently evolution proceeded from coccoids to filaments and then to coccoids and filaments of greater complexity, giving rise to the major families. This is quite evident from comparison of the blue-green algal microfossils of the Gunflint Iron Formation with those of the Bitter Springs Formation. The Bitter Springs assemblage has a decidedly more modern aspect (Schopf, 1970; Schopf and Blacic, 1971). Significantly, the Gunflint stromatolites contain many morphotypes of uncertain affinities. Morphological “experimentation” may have taken place during Gunflint time giving rise to many bizarre, now extinct, genera. It is certainly possible that the Gunflint microbiota includes representatives of higher taxa of extinct prokaryotic photoautotrophic organisms. Although preservation is poor, preliminary data from Paradise Creek stromatolites (Licari and Cloud, 1972) indicate that already, by about 1,600 m.y. ago, the microbiota was morphologically comparable to modern blue-green algae. The Beck Springs stromatolites (ca. 1,300-1,400 m.y. old) have a few problematica but not as many as the Gunflint. By Bitter Springs time (ca. 800-1,000 m.y. old), the algal mat biocoenose is morphologically extremely similar to that of Recent algal mats.


It is within the blue-green algal template that other organisms (including some bacteria and diverse eukaryotic algae) have contributed to the construction of stromatolites. Most of these have independently evolved the capacity to build stromatolites. In view of this polyphyletic origin, evolutionary information derivable directly from stromatolites and their microbiota is intrinsically limited.

ACKNOWLEDGEMENTS

The authors are grateful to Prof. S. Golubic for his unpublished wisdom concerning algal mats, his continuing cooperation and his encouragement. Dr. H. Hofmann critically read the manuscript. We thank NSF GA-37140, NSF GA-1382 and NSA NGL-22-007-069 (to Dr. Barghoom) and NASA NGR-004-025 (to Dr. Margulis) and the Boston University Graduate School for support.


Chapter 4.5

BIOCHEMICAL MARKERS IN STROMATOLITES

David M. McKirdy

INTRODUCTION

Stromatolites still await systematic and detailed organic geochemical study. Apart from the pioneering work of Hoering (1962,1964, 1967) and a later investigation of the oldest known stromatolites by Schopf et al. (1971), carbonate rocks containing laminated organosedimentary structures built by microscopic algae and bacteria have been largely overlooked by geochemists in their search for chemical and isotopic evidence of the origin and development of biological activity on the earth’s surface (Maxwell et al., 1971; McKirdy, 1974). Analyses of the organic matter in black fossiliferous cherts from several Precambrian stromatolitic beds (Or6 et al., 1965; Schopf et al., 1968; Smith et al., 1970) were prompted by the antiquity and remarkable preservation of the fossil assemblages, rather than by their occurrence in stromatolites per se. A marked and not altogether explicable sampling bias in favour of argillaceous and siliceous sediments has meant that most of the reported finds of chemical (or molecular) fossils, i.e. hydrocarbons, fatty acids, porphyrins, amino acids, carbohydrates and stable carbon-isotopic fractionation of apparent biological origin, have been made in shales and cherts, some older than 3,000 m.y. Consequently, little is yet known about the fossil organic matter which represents the remains of benthonic mat- building communities in ancient carbonate-depositing environments.

Stromatolites are among the most common and long-ranging fossils in the geological record. Their macroscopic size and widespread geographic distribution¶ their usual occurrence in fine-grained, highly indurated carbonate or chert, and the existence of “living” analogues in a variety of Recent environments, together make stromatolites extremely attractive materials for organic geochemical investigation. This chapter gathers together biochemical and geochemical data pertinent to our understanding of the composition, derivation and function of the organic content of fossil stromatolites. New analyses of some Australian Palaeozoic and Precambrian stromatolites are reported and, finally, the potential of organic geochemistry to contribute to future stromatolite research is assessed.

164 D.M. MCKIRDY

A L K A N E S

normal * sterane

I S 0 /LAvvvv

R=H,CH,,$H, etc anteiso 7

7-methyl

8-methyl

isoprenoid &d.d,vb

hopone

@ R = H,CzH5. is0 C3H7 etc.

FATTY ACIDS

Saturated Unsaturated

normal ~ r + a w v v C O , H C,, I -CO,H

ISO k C 0 , H C, , , -COZH

onteiso y - C O z H a-Cie CO,H

isoprenoid W C 0 , H Y -CIe -COLH

P I G M E N T S

R 7 Re R I = CHzCHzC02H

Rz-7= CH,, CzH5 etc R, = CH:CH, or CH,CH3

corotenoid

AMINO ACIDS

R = H , C H 3 .CH2C02H etc 7

H- C - C 0 2 H I

NH,

Fig. 1. Carbon skeletons of some potential biochemical markers in stromatolites.

An impressive number and variety of stable organic compounds, interpreted as chemical fossils by virtue of their biologically indicative molecular configuration (Fig. 1) or isotopic composition, have now been extracted from sedimentary rocks of all ages (Maxwell et al., 1971). When found in sufficient concentration in a stromatolite such compounds are likely to have been derived (at least in part) from the organisms that participated in its formation (but see below). Stromatolites are 'commonly abiophoric (i.e. lacking preserved microfossils: Hofmann, 1973, p. 350), in which case indigenous chemical fossils might prove to be the only remaining clues to the


identity of the original mat-builders. Of course, a laminated sedimentary structure can be stromatolite-like in appearance and yet abiogenic. Any organic matter it contained would then be allochthonous and largely incidental to its origin.

Despite recent advances in stromatolitology, some fundamental questions remain unresolved or, as is particularly true of abiophoric stromatolite occurrences, can be only tentatively answered (cf. Hofmann, 1973). Are all fossil stromatolites biogenic? If biogenic, were oxygen-producing cyanophytes mainly responsible? Were the cyanophytes filamentous or coccoid? How can the contribution of eukaryotic algae to a stromatolitic biota be recognized? Did ancient bacteria (and perhaps even fungi) ever build stromatolites? How does one distinguish such stromatolites from those built by algae? To establish in what way, and to what extent, organic geochemical data may shed light on these problems, we shall consider the various classes of chemical fossil and the potential significance of each for the palaeobiochemistry and palaeoecology of stromatolites.

ORGANIC MATTER IN ALGAL MATS AND STROMATOLITES

The microstructure of most algal laminates and stromatolites, whether Recent or ancient, comprises an alternation of light (sediment-rich) and dark (organic-rich) laminae arising from periodic fluctuation in the relative rates of mat growth and sediment deposition. The sediment-rich layers in Recent stromatolites contain less than 5% organic matter (Gebelein and Hoffman, 1973). The organic matter in the dark laminae is derived mainly from dead blue-green algal cells, filaments and empty gelatinous sheaths, although eukaryotic (green and red) algae and bacteria may also have been present and contributed detritus.

Disseminated organic matter can influence both the grain size and the mineralogy of stromatolitic sediments. Many of the Australian stromatolites described by Walter (1972a) “are composed of carbonate with a grain size of less than 30pm, and frequently less than 15pm. Furthermore, the pigmented laminae are nearly always finer grained than contiguous pale laminae” (op. cit., p. 97). Similarly, Schopf and Blacic (1971, p. 929) noted a correlation of finer-grained quartz with regions of relatively high organic content in laminated cherts from the Bitter Springs Formation. In both instances organic matter has presumably inhibited recrystallization. Walter (1972a, p. 59) also describes a situation in the Bitter Springs Formation at Jay Creek where several stromatolite biostromes contain columns which are black (organic-rich) and consist mainly of calcite, in striking contrast to the pale grey predominantly dolomitic interspaces. In this case, organic matter coating carbonate grains in the stromatolite columns appears to have helped prevent their post-lithification dolomitization. The completely opposite

166 D.M. MCKIRDY

TABLE I

Total organic carbon content (TOC) of algal mats and other Recent sediments

Location TOC ("/.I

Reference

Harbor Island, Texas

living mat first mud layer

2 second mud layer E

7 Baffin Bay, Texas

* Kleberg Point Lagoon,

well-developed laminae disrupted oxidized laminae

Gulf of Batabano, Cuba

2 shallow marine carbonate

; Saanich Inlet, British Columbia ' carbonate-poor, reducing muds c 6 Choctawhatchee Bay, Florida

B

c

estuarine muds

Parker and Lea (1965)

32 1.1 0 .e4

Behrens and Frishman (1971)

2-5 1-1.7

Hunt (1967)

1 .4 (mean)

Brown et al. (1973)

3.4 (mean)

Palacas et al. (1972)

3.6 (mean)

effect is also possible, in which organic matter acts as a source of Mg for dolomite formation (Friedman et al., 1973; Gebelein and Hoffman, 1973).

The amount of organic matter preserved in a stromatolite upon lithification represents that portion of the primary algal and bacterial production which has survived decomposition by the aerobic and anaerobic bacterial (and sometimes fungal) heterotrophs originally present in the lower zones of the living mat (Golubic, Ch. 4.1). A measure of this residual organic matter is provided by the total organic carbon (TOC) value of the stromatolitic sediment (Tables I and 11).

On the limited data available, Recent laminated algal sediments appear to have TOC values comparable to those for other Recent sediments (Table I). No organic carbon values have yet been reported for Recent non-stratiform stromatolites such as the club-shaped forms at Shark Bay, Western Australia.

The concentration of organic carbon in fossil stromatolites (Table 11) is generally 1-2 orders of magnitude lower than in Recent algal laminates. A Precambrian algal dolomitic limestone containing 1.1% TOC from the Purcell Supergroup, Alberta (Hodgson e t al., 1968) is exceptionally organic-rich. Such high TOC values may reflect a lower energy, less emergent or more poorly drained depositional environment, and hence a more stable anoxic regime below the actively growing mat. TOC values for Precambrian


TABLE I1

Total organic carbon content (TOC) of stromatolites and other ancient carbonates

Location TOC Reference

Cavan Limestone, N.S.W. (Devonian) 0.06-0.42, McKirdy (this paper) Wilkawillina Limestone, S.A. 0.01 McKirdy (this paper) (Cambrian) Australian stromatolites 0.01-0.05 McKirdy (this paper)

.d d (Precambrian) e, o Purcell Supergroup, Alberta 1.1 Hodgson et al. (1968)

(Precambrian) 2 Transvaal Dolomite Series, S.Africa 0.1 Hodgson et al. (1968) ;j (Precambrian)

Bulawayan Group, Rhodesia 0.5 Hoering (1964) (Precambrian)

-

Russian Platform carbonates (mean values)

Mesozoic and Cainozoic 0.47 8 2 Palaeozoic 0.26 5 Late Proterozoic 0 .06

Ronov and Migdisov (1971) 3

0 8 Early and Middle Proterozoic 0.01

stromatolitic chert - e.g. Gunflint Iron Formation, 0.03-0.07% (Kvenvolden, 1972); and Skillogalee Dolomite, 0.08-0.21% (Table IV and unpublished results) -fall within the same range as those for stromatolitic carbonate. Mean TOC values for carbonates of different ages from the Russian Platform are included in Table 11. By comparison, stromatolitic carbonates are not unusually rich in organic matter.

The organic content of a sediment may be divided analytically into two major fractions. The material insoluble in organic solvents (kerogen) usually constitutes the bulk of the total organic matter. The solvent-extractable portion (bitumen or geolipids) is an extremely complex mixture of many different compounds. It is in this fraction that chemical fossils have been most commonly sought and found. The chloroform extracts of six fossil (Cam- brian to Recent) algae analysed by Das and Smith (1968) amounted to 0.3-0.6% of the sample weight. By contrast, exhaustive extraction of the stromatolitic carbonates listed in Table IV with benzene/methanol gave (with one exception viz. sample 8, 132ppm) geolipid yields of 2-20ppm. These yields vary with the diagenetic grade of the sediments (Fig. 3). A sample of fossiliferous laminated black chert from the stromatolitic Bitter Springs Formation contained 27 ppm extractable organic matter (Schopf, 1968), whereas the corresponding yield from the Precambrian stromatolitic cherts analysed by Smith et al. (1970) was less than 0.1 ppm.

168 D.M. MCKIRDY

SY NGENEITY

Most analytical techniques currently employed in organic geochemistry are designed to ensure that the compounds isolated from a rock are indigenous to it and are not field or laboratory artifacts. However, where the chemical fossils are present in very low concentration, as for example in many Precambrian sediments, it is difficult to be certain that they (or their biochemical precursors) were actually deposited with the sediment and did not enter the rock some time after its lithification (Hoering, 1967; Smith et al., 1970 ; Kvenvolden, 1972 ; McKirdy, 1974). This uncertainty unfortunately hampers the interpretation of much of the organic geochemical information so far obtained on ancient stromatolitic sediments, especially cherts.

Aliphatic hydrocarbons

The stromatolitic cherts analysed, all Precambrian in age, are from the Gunflint Iron Formation (Or6 et al., 1965; Smith et al., 1970), the Bitter Springs Formation (Smith et al., 1970) and the Paradise Creek Formation (Smith et al., 1970). The concentrations of total alkanes found were variable but generally low (0.005 -5 ppm). Porosity and permeability measurements on ancient cherts (Smith et al., 1970; Sanyal et al., 1971) demonstrate that such sediments are in fact sufficiently porous and permeable to have admitted significant quantities of younger organic compounds from migrating formation fluids under pressure while buried at depth, or later by capillary action when exposed at the surface. Smith et al. (1970) found most of the extractable organic matter to be located along microfractures within the chert matrix and concluded that it was of post-depositional and probably comparatively recent origin. Only trace amounts (several parts per billion) of the alkanes were considered likely to be truly indigenous.

Stromatolitic carbonates (Table IV) yield alkanes in concentrations which are low (1-5ppm) but are less variable, and on the average higher, than those for stromatolitic cherts. Moreover, their porosity (0.5-4.4%) and nitrogen permeability (3.1 to 1.2 lo-' millidarcy) are at least as low (Table IV) as those of the cherts studied by Smith et al. (1970). But are the hydrocarbons isolated from the carbonates any more likely to be syngenetic? An answer to this question can be attempted by turning to the composition of the coexistent insoluble organic matter (kerogen) in the carbonates (Table IV).

The atomic hydrogen to carbon ratio (H/C) of kerogen provides a sensitive measure of its diagenetic (or incipient metamorphic) rank and, by inference, that of the-host sediment (Tissot et al., 1974; McKirdy et al., 1975). Kerogen, being insoluble and immobile, is almost certainly syngenetic with the sediment in which it is found. Hence, where the H/C value of the kerogen is known, the degree of thermal maturation any syngenetic


hydrocarbons have undergone may be assessed. Both the yield and the distribution pattern of the alkanes in a sedimentary rock change systematically with increasing diagenesis (Brooks and Smith, 1967; Albrecht and Ourisson, 1969) and in a manner dependent on the rock type and the nature of the paretn organic material (Powell and McKirdy, 1973a). The degree to which these two parameters (yield and distribution of alkanes) coincide with what might be predicted in view of the rank of the kerogen and the lithofacies of the host sediment may thus in turn indicate whether the alkanes are coeval with the kerogen or have since migrated into the rock (cf. McKirdy, 1974, p. 122). Figs. 3 and 4, plots of geolipid yield and alkane yield, respectively, against kerogen H/C, show that the stromatolitic carbonates listed in Table IV conform with the general diagenetic trend for marine carbonates. This is consistent with a syngenetic origin for the stromatolitic alkanes.

Fatty acids

The significance of the fatty acids found by Smith et al. (1970) in stromatolitic cherts of Precambrian age is just as uncertain as that of the dkanes they yielded (see previous section). Both free and bound fatty acids were sought, and again (with the exception of the free fatty acids from the Bitter Springs and Paradise Creek cherts) the highest concentrations were obtained from the surfaces of the chip size fraction. More importantly, however, labile, unsaturated acids were found in unexpectedly high concentration in all extracts, some of which also contained longchain acids in the range Czl to CZ4. The latter are features more indicative of recent contamination than of a Precambrian microbial source.

Fatty acids interact with carbonate mineral surfaces by way of processes such as chemisorption and form various complexes, including natural calcium soaps, i.e. adipocere (Suess, 1970). Although yet to be specifically investigated in Recent algal mats this fatty acid-carbonate association is probably important in determining the ultimate fate of the fatty acids. Adipocere could be expected to form rapidly following the release of fatty acids from decaying algal cells, and thereafter resist further bacterial decomposition (cf. Berner, 1968). Burlingame and Simoneit (1968) suggest that such mineral-bound fatty acids are more likely to be syngenetic with the sediment in which they are found than afe free or interstitially trapped geolipids.

Amino acids

Amino acids have been recovered by ammonium acetate leach (free acids) and HC1 hydrolysis (combined acids) from many ancient sediments, including stromatolitic limestone of the Bulawayan System (Oberlies and Prashnowsky, 1968) and laminated fossiliferous cherts from the Gunflint

170 D.M. MCKIRDY

Iron Formation and Bitter Springs Formation (Schopf e t al., 1968). Serine and threonine, both extremely labile amino acids, were found in all three rocks. On stability grounds alone, these two compounds are most unlikely to have survived from the Precambrian. This places in doubt the syngenetic origin of the remaining acids. Stereochemical examination of the amino acids extracted from samples of Gunflint chert (Abelson and Hare, 1969) and Fig Tree chert (Kvenvolden et al., 1969) showed that they comprised only the L-isomers. In the absence of any stabilization by the mineral or kerogen matrix, amino acids in sediments older than about 2 m.y. should be racemic mixtures of the L- and D-isomers. Clay minerals and kerogen may inhibit racemization (Akiyama and Johns, 1972), but the calcareous fraction of Recent marine sediments (Hare, 1972) and the silica of ancient cherts appears to exercise no such stabilizing effect (Abelson and Hare, 1969). A younger biological source for part, if not all, of the extractable amino acids in the above stromatolitic sediments thus appears likely.

ALGAL AND BACTERIAL MARKERS

Aliphatic hydrocarbons

Potential markers The lipid fraction of algal and bacterial microorganisms includes hydro-

carbons, fatty acids (as esters), fatty alcohols (as glycosides), pigments, sterols, etc. (Kates, 1964; Nichols, 1973), which together comprise a quantitatively variable (0.3-85% of cellular dry weight, mostly < 30%: data from Abelson, 1967), but geochemically highly significant, group of compounds. Of these compounds the hydrocarbons are the least labile and hence the most likely to survive sedimentation and burial with their carbon skeleton more or less intact. Hydrocarbons constitute 0.006-0.12% (dry cell weight) of blue-green algae (Or6 et al., 1967; Han et al., 1968a; Winters et al., 1969). Three algal mats studied by Or6 et al. (1967) yielded 2.7-5.8% lipid and only 0.001-0.01% hydrocarbons. By comparison, the ubiquitous sulphate- reducing bacterium Desulfouibrio desulfuricans contains 5-976 lipid of which 25% is hydrocarbon (Davis, 1968).

The composition of the aliphatic hydrocarbons present in various extant blue-green (and green) algae and Recent algal mats is summarized in Table 111. The C1, n-alkane (and/or alkene) is almost invariably conspicuous and Czl to Cz9 n-alkanes are less common than homologues of shorter chain length. Nevertheless, certain algae (blue-green, green and red) (Clark and Blumer, 1967; Gelpi et al., 1970; see also Table 111), bacteria (Albro and Huston, 1964; Davis, 1968; Albro and Dittmer, 1970) and fungi (Weete, 1972) do contain long-chain n-alkanes (or n-alkenes), among which, moreover, the oddcarbon-numbered homologues tend to be dominant. Another

E r 0

0 0 cc

m r3

TABLE I l l Y 0

Hydrocarbons and fatty acids of geochemical significance in algal mats, mat-building cyanophytes, and extant representatives of algal genera implicated in Precambrian stromatolitic microhiotas

Algae and algal mats Hydrocarbons’ %

ai br br 2 15:O 16:O 17:O 17:l 17:2 18:0 18:O 18:O 18:l 19:0 19:l pris20:O phyt 21:O 212 22:O 23:O 23:l 25:l 26:l 27:l 272 Others R

m Algal mats, Texas Gulf Coast

Harbor Island2 not determined Port Aransas 2 2 3 6 4 3 2 1 11 1 8 Oyster Bay3 < 1 2 12 31 7 10 6** 7 2.5 5 2.5 4 2 . 2 6 Laguna Atacosa 3 29 43 5 2 3.75 1 1.5 < 1 < 1 Padre Island 2 24 22 1 16 21** 1 < 1 1 < 1 < 1 1 9

Mat-building cyanophytes

Lyngbya aestuarii4 2 6 35 1 38* 16

Genera implicated in Precambrian stromatolitic microbiotas

Chlorella e.g. C. pyrenoidosa 19 77 Anacystis e.g. A. cyanea’ 87 13’

Spirulina e.g. S. platensis 10 20 70 < 1

Microcoleus chthonoplastes 2 4 92 2

A. montana’ 12

Lyngbya e.g. L. lagerharmii 1 4 8 6 1 8 Oscillatoria e.g. 0. wifliamsii 9 < 1 91 < 1 < 1 Myxosarcina e.g. M. chroococcoides not determined

9 8 15 4 35 3

(continued on p. 172)

TABLE Ill (eonrmued)

Lyngbya e g. L logerhoimii < 1

Oscillolono e g. 0 . willioma8i < 1

Myxosoreina e g. M. ehrooeoecondes

Algae and algal mats Fatty ae~ds ' References

I I i ai i ai 12:O 12:O 1 3 0 14-0 14:O 1 4 0 14: l 15:O 15:O 15:O 16:O 16-0 16-1 1 6 2 17:O 17:l 18-0 l 8 : l 18:2 18:3 Others

Alral mats. Texas Gulf Coast

Harbor ~sland' ' 58 11 1 2 3 6 < 1 Parker and Leo (1965)

Port Aranras not determined Winters ct al. (1969)

Oyster ~ a y ' 1 < 1 3 4 1 3 24 4 23 4 6 3 4 4 16 Orde t al. (1967) Laguna Ataeosa 2 < l 4 3 2 4 6 4 17 3 10 1 3 2 3 7 27 Old e t a ] . (1967) Padrp Mand 2 2 5 6 7 3 1 2 2 7 9 3 9 3 2 19 O r d e t nl. (1967)

Mat-building cyanophytes

Microcobus chlhonaplasbs 2 5 31 13 4 14 5 18 Parker e t al. (1961) Winters e t al. (1969)

Lyngbyo oestuarii' 26 15 3 8 17 Gelpi et al. (1970) Schnetder et al. (1970)

Genera implicated in Rceambrtan stmmstolitne mierobiotar

Chlonllo e g. C. pyrenoidoso 2 21 2 6 17 19 26 Gelpi et a1. (1970) Sehneider et al. (1970)

Anoeyrlls e.g. A eyoneo" 23 20 2 7 1 0 5 Gelpi et al. (1970) Schncader et a1 (1970)

A monlano5 21 3 25 18 19 Gelpi et al. (1970) Sehneider ct d. (1970)

Spirultno e.g. S phfemis 45 6 2 10 16 13 Gelpi et al. (1970) Schneider et al. (1970) Parker e l d. (1967) Wnnters et al. (1969) Parker et al. (1967) Winters et al. (1969) Nichols and Woad 11968)

' Straightihain (i.e. normal) isomers except where otherwise indicated, viz.. I = k o ; ai = onteiso; br = branched; pris = pristane; phyf = ~ h ~ t a n e . Numbers before colon = numbers of carbon atoms in molecule; numbem after colon = number of double bands. Concentrations given as relative percent of total hydrocarbon or fatty-acid fraction.

' Mainly Lyngbyo eonfervoides

' Mainly Microeobw chlhonoplrrrtes. Lyn,gbya oestum'i and, in l e e r m o u n t . Sch~ro lh r i r coleieolo

' Morphologically similpr to Polueolyngbya B o g h o o r n h a , Bitter Springs Formation (Sehopf, 1968).

Morpholog#edy similar t o Polwoonocystk vulgaris. Bitter Springs Formation (Schopf. 1968).

Mixture of 7.. and 8-methylheptadeeane

"includes some 17:l hydrocarbon.


possible ultimate source of high molecular-weight n-alkanes are the long-chain (c26 and C28) polyhydroxy alcohols present as glycosides in the heterocysts of bluegreen algae (Nichols, 1973). It is worth noting here that high molecular weight (> CzO) compounds comprise a much greater proportion of the total aliphatic hydrocarbons in non-photosynthetic bacteria (> 50%) than in photosynthetic bacteria (< 15%) (Han et al., 196813). Non-photosynthetic bacteria are the last viable organisms to populate an algal mat before burial finally removes it from the zone of microbiological activity.

Various aliphatic hydrocarbons appear to be specific to certain taxa of extant microorganisms. These hydrocarbons may well constitute a geochemical basis for recognizing the contribution of a particular biota to stromatolite communities. Examples of such chemotaxonomic markers are 7-methyl and 8-methyl heptadecane (Fig. 1) in cyanophytes (Gelpi et al., 1970; Han and Calvin, 1970); and pentacyclic triterpanes of the hopane type (Fig. 1) in prokaryotic algae and bacteria (Kimble et al., 1974). Steroidal alcohols, which readily revert to steranes (Fig. 1) during early diagenesis, are indicative of eukaryotic algae and fungi (although see e.g. De Souza and Nes, 1968; Reitz and Hamilton, 1968; Schubert et al., 1968; Gelpi et al., 1970; and Nadal, 1971, for reports of their occurrence in prokaryotes).

The isoprenoid (branched) alkanes pristane (C19) and phytane (C20) (Fig. 1) are two widely occurring biochemical markers usually considered to be products of the diagenesis of the phytyl (C2,,) side-chain of chlorophyll. Pristane (and rarely phytane) may also occur as discrete hydrocarbons in certain algae (Clark and Blumer, 1967; Or6 et al., 1967; Han et al., 1968b), bacteria (Or6 et al., 1967; Han et al., 1968b) and algal mats (Table 111). The precise route to their diagenetic formation in sediments has not yet been fully elucidated, but it appears that their relative concentration depends largely on the degree of oxidation of the parent organic matter during the early stages of chlorophyll decay. In short, the generation of pristane is favoured under oxidizing (aerobic) conditions, whereas a reducing (anaerobic) environment preferentially gives rise to phytane (Brooks et al., 1969; Powell and McKirdy, 1973b).

Stromatolitic cherts The normal alkanes isolated from Precambrian stromatolitic cherts by

Smith et al. (1970) typically display a smooth unimodal distribution in the range c16 to c36, with a maximum concentration at CZ2. The alkane patterns obtained from the Gunflint chert by Or6 et al. (1965), however, tended to be bimodal with maxima at n-C18 or -CI9 and n-Cz2. Similar bimodal distributions have been reported from other Precambrian sediments (McKirdy, 1974, p. 112).

A slight dominance of odd- over even-carbon-numbered homologues is evident in the n-alkanes larger than c 2 6 from the above cherts. A predominance of odd-carbon-numbered n-alkanes is characteristic of many sediments to

174 D.M. MCKIRDY

which higher plants of terrestrial origin have contributed detritus (Powell and McKirdy, 1973a). Its occurrence in the geolipid fraction of Precambrian sediments is unexpected. Even if such an odd-over-even predominance was originally present in the n-alkanes of a precursor Precambrian stromatolitic microbiota (see previous section), it would probably have disappeared during burial diagenesis (Brooks and Smith, 1967).

In view of their doubtful indigeneity, any interpretation of the alkanes isolated from stromatolitic cherts is necessarily equivocal. The alkane distributions (in the absence of data on the diagenetic grade of the cherts) could be equally well explained in terms of either a microbial or a contaminative higher plant source, although the relict odd-over-even predominance strongly favours the latter. Even the low values (< 1) for the ratio of the concentrations of the two isoprenoid alkanes, pristane and phytane, are ambiguous. LOW pristane to phytane ratios have been found both in marine crude oils derived from aquatic microorganisms (Powell and McKirdy, 197313) and in low-rank coals of non-marine origin (Brooks e t al., 1969).

Stromatolitic carbonates Distributions of n-alkanes obtained from limestones (unlike those from

shales) display little variation with advancing diagenesis (Powell and McKirdy, 1973a) and this seems to hold true for stromatolitic carbonates (Table IV). The stromatolitic n-alkanes range from CI5 to C29 with a maximum between c18 and Czs , and show no noticeable preference for odd- nor evencarbon-numbered molecules. However, certain subtle differences are evident in the total alkane patterns (Fig. 2), and these are largely attributable to diagenetic maturation, as is discussed below.

The least altered stromatolite in Table IV, Acaciella australica (sample 8, kerogen H/C = 0.82), gives an alkane pattern (range n-ClS to n-Cz3, maximum n-Cl8 ) considered typical of organic matter derived from blue-green algae (e.g., Winters et al., 1969; Gelpi et al., 1970). The straight-chain hydrocarbons (alkanes and alkenes) in modern mat-building cyanophytes, Recent algal mats, and other algae similar to th%e implicated in certain Precambrian stromatolitic microbiotas (Table 111), generally fall in the same range. In the stromatolites of higher rank (kerogen H/C = 0.7-0.2), the n-alkane envelope is somewhat broader (CI6 to Czs) and its maximum is located at a higher carbon number (between C19 and CZs). That is to say, maturation of the organic matter in those stromatolites above a rank equivalent to H/C = 0.8-0.7 has resulted in the generation of longer-chain (CZl to Czs) n-alkanes from the kerogen (cf. Albrecht and Ourisson, 1969; Powell and McKirdy, 1973a). There is a corresponding increase in the proportion of alkanes in the geolipid extract (Fig. 4) but an accompanying diminution of the total geolipid yield (Fig. 3).

The diagenetically more mature stromatolites also differ from Acaciella australica in having a much higher proportion of branched and cyclic isomers


- 7 SKILLOGALEE DOLOMITE

7 80!CO60 I

175

C A V A N LIMESTONE c S!,"hforrn

IC

n

BITTER SPRINGS FORMATION

n-C1* I Acac!e/la australtca

n

A looo 150' 200'' 250'C -

Column Tempcrolurs

TOOGANlNlE FORMATION

1000 i5O0 ZOOo P50°C - Colunn rcmperolure Column Temprolun

Fig. 2. Gas chromatograms of saturated hydrocarbons (alkanes) isolated from Australian stromatolites.

176 D.M. MCKIRDY

I I I

0 100 200 300 400 500 600 700 800

Geolipid e x i r a c t - rng/g C o,g

Fig. 3. Influence of diagenesis on geolipid yield from stromatolites and other marine carbonates. Data from Powell and McKirdy (1973a), Powell et al. (1975), McKirdy and Powell (1976) , and McKirdy (unpublished results).

in their total alkane fraction. (In Fig. 2 the branched/cyclic alkanes comprise the baseline hump and the peaks between the n-alkanes.) This increase in relative concentration of branched/cyclic alkanes is particularly apparent in Conophyton f. (sample 14, kerogen H/C = 0.65) and ?Baiculia f. (sample 13, kerogen H/C = 0.23). It may be a result of catalytic cracking of hydrocarbons during late diagenesis. Such cracking reactions are facilitated by clay minerals and proceed via carbonium ion intermediates which can undergo alkyl isomerization and 0-splitting (Eisma and Jurg, 1969) and lead eventually to the production of branched and cyclic alkanes (McKirdy, 1971). An alternative explanation is that bacteria populating the dark interior of the original mat preferentially metabolized the algal n-alkanes (Bailey et al., 1973).

The stromatolite containing the most highly altered organic matter is Baiculia burru (sample 9, kerogen H/C = 0.10). Its alkane distribution (range n-CI5 to n-C,,, maximum ~I-C,~) is similar to that of Acaciella australica and presumably reflects the thermal cracking of the longer-chain alkanes generated from the kerogen at an earlier stage of diagenesis into hydrocarbons of lower molecular weight.

Pristane and phytane were identified in the saturated hydrocarbons

E 0

Preliminary analysis of some Australian Precambrian and Palaeozoic stromatolites (see also Appendix) P 0

Sample TOC Geolipid extract Alkanes n-Alkanes Pristane Kerogen Porosity Permeability 0 4 0 extract (% ash-free basis) (%) (millidarcy) q

TABLE IV

(56) (ppm) (mg/gC)ppm % range max. Phytane C H N S O* ash H/C O/C (%) V H

1 0.24 11 4.7 3.3 29.8 C,,-C,, C,, - 81.2 3.7 0.9 1.7 12.5 8.4 0.540.12 2.3 0.0005 0.12 2 2 0.42 10 2.7 1.4 13.3 C,,<,, C,, 0.5 88.2 1.9 0.3 2.6 7.0 20.2 0.26 0.06 4.4 0.0002 0.00003 3 0.06 8 13.2 1.8 22.2 C,,-C,, C,, - n.d. n.d. n.d. n.d. n.d. n.d. 0.70n.d. 1.2 0.0002 0.0003 4 0.01 7** 69.3**3.7** n.d. n.d. n.d. n.d. 73.03.0 2.1 9.0 12.9 6.4 0.49 0.13 1.0 0.005 0.004 5 0.03 5 18.1 1.4 26.1 C,,-C,, C,, 0.6 77.9 3.1 2.7 6.6 9.7 44.2 0.47 0.09 0.5 0.0004*** $

0 6 0.01 17 173.4 4.7 27.4 C,,-C,, C,, - 81.2 3.4 0.9 n.d. n.d. 15.5 0.49 n.d. 2.2 n.d. n.d. 7 0.02 10 48.6 2.5 25.9 C16-Cz8 C,, 0.6 84.0 1.7 <0.9 5.7 7.7 44.3 0.240.07 2.8 0.006 0.006 8 0.02 132 659.3 1.5 1.2 12,,<,~ C,, 1.1 72.7 5.0 1.5 7.0 13.8 5.7 0.82 0.14 0.8 0.007 0.006 5 9 0.02 11 77.2 2.5 10.9 C1,-C2, C,, 0.8 90.9 0.8 <0.5 3.8 4.0 23.2 0.10 0.03 2.2 0.003 * * * M

10 0.05 6 28.2 3.0 53.8 C16-C26 C,, 0.4 75.3 2.1 1.0 1.6 20.0 8.0 0.33 0.20 n.d. n.d. n.d. m 11 0.09 n.d. n.d. n.d. n.d. n.d. n.d. n.d. 85.2 2.2 0.7 2.3 9.6 53.1 0.31 0.08 n.d. n.d. n.d. 12 0.08 n.d. n.d. n.d. n.d. a d . n.d. n.d. 85.2 2.0 0.7 2.0 10.1 54.0 0.27 0.09 n.d. n.d. n.d. 13 0.03 8 16.6 1.4 16.7 C,,-C,9 C,, 1.3 88.7 1.7 1.3 1.8 6.5 38.3 0.230.06 0.8 0.0008 0.0005 14 0.04 5 12.3 1.9 39.1 C,,-C:,, C,, 1.2 77.24.2 1.8 6.2 10.6 16.4 0.6Y0.10 1.7 0.0007 0.0006 15 0.01 2 15.0 n.d. n.d. n.d. n.d. n.d. 73.8 2.7 <0.6 5.8 17.1 20.2 0.43 0.17 2.1 0.005 n.d.

* By difference; ** includes free S; *** orientation unknown; n.d. = not determined.

178 D.M. MCKIRDY

L I I 6 1 I I 0 10 20 30 40 50 60 70 8 0

Percent alkanes in geolipid e x t r a c t

Fig. 4. Influence of diagenesis on alkane yield from stromatolites and other marine carbonates. Data from Powell and McKirdy (1973a), Powell et al. (1975), McKirdy and Powell (1976), and McKirdy (unpublished results).

isolated from many of the stromatolitic carbonates in Table IV. The low pristane to phytane ratios (0.5-1.3) reflect the existence of reducing conditions (Powell and McKirdy, 197313) in the zone of active organic decomposition below the surface of the original algal mat (cf. Sorensen and Conover, 1962; Golubic, Ch. 4.1).

In such a reducing microenvironment, the unsaturated hydrocarbons (alkenes) which occur in significant concentrations in blue-green algae and Recent algal mats (Table 111) are unlikely to survive for long before under- going hydrogenation to the corresponding alkanes (Blumer, 1965).

Finally, the remarkable similarity of the alkane patterns of Tungussia inna (Wonoka Formation) and ?Tungussia wilkatanna (Skillogalee Dolomite) (Fig. 2) should be noted. Not only are these hydrocarbon distributions nearly identical, but they are also quite distinct from the patterns obtained from other Precambrian carbonates. The exciting possibility of chemotaxonomic correlation within certain stromatolite form genera is suggested. A lower order of congruence is evident between the alkanes from Conophyton f. (Tooganinie Formation) and ?Baicalia f. (?Skillogalee Dolomite) (Fig. 2). Stratigraphic gradations between these two forms are common (M.R. Walter, personal communication, 1974).


Fatty acids

179

Potential markers Contemporary algae contain fatty acids in appreciably higher concen-

tration than aliphatic hydrocarbons (Schneider et al., 1970). However, algal fatty acids possess shorter chain lengths (Clo to C22) than do the hydrocarbons (CIS to C33), and exhibit a higher degree of unsaturation. The C14, C,, and c18 acids occur most frequently, whereas acids with an odd number of carbon atoms in their chain are rare (Table 111). The CzO and CZ2 acids found in eukaryotic algae (Ackman et al., 1968) appear to be absent from the bluegreen algae (Nichols, 1973).

In contrast to the eukaryotic algae, cyanophytes display a considerable interspecific diversity in fatty-acid composition, arising mainly from differences in the relative proportions of the (saturated and unsaturated) acids (Nichols, 1973). According to Kenyon and Stanier (1970), many filamentous blue-green algae differ from certain unicellular blue-green algae in containing high concentrations of poly-unsaturated fatty acids, particularly the

series. If it could be shown that the diagenesis of poly-unsaturated fatty acids followed a different course (e.g., preferential incorporation into an insoluble kerogen-like polymer) to that of saturated and mono-unsaturated acids, then the uniformly low concentration (or complete absence) of poly- unsaturated acids in unicellular blue-green algae might well provide a geochemical means for identifying stromatolites built by non-filamentous prokaryotic algae. However, the finding of appreciable quantities of the c18:2 and c18:3 acids in three other unicellular cyanophytes by Schneider et al. (1970) suggests that a chemotaxonomic distinction between coccoid and filamentous blue-green algae, based on fatty acid composition, may not be universally valid.

The chances of recognizing a bacterial contribution to the organic matter in Recent (and possibly ancient) stromatolites are considerably better. Bacteria commonly contain high concentrations of the unusual singly branched (iso, anteiso) fatty acids, in addition to normal (straight-chain) acids (Kates, 1964; Kaneda, 1967; Parker et al., 1967). Is0 and anteiso acids have been identified in Recent marine sediments (Cooper and Blumer, 1968) and algal mats (Table 111). They might be expected to survive for a time in fossil stromatolites, particularly those with a mild thermal history, but will eventually undergo decarboxylation to the corresponding is0 and anteiso alkanes.

The concentration of fatty acids in benthonic algae diminishes rapidly following post mortem disintegration of cellular membranes. For example, a living algal mat at Harbor Island (Table 111) contained 0.17% free fatty acids, compared with concentrations of 0.002% and 0.003%, respectively, in the first two buried mats immediately underneath it (Parker and Leo, 1965). A similarly rapid decrease was also observed in the relative concentration of

180 D.M. MCKIRDY

unsaturated acids, from 39% of the total acids in the living mat to 7% in the second buried mat. The exact fate of the unsaturated fatty acids is unclear. In the anaerobic enviroment of the mat interior they could be converted to the equivalent saturated acid by hydrogenation of the double bond(s) (Blumer, 1965; Rhead et al., 1971). Under persistently low Eh and and high pH conditions, such as exist in poorly drained algal mats (Golubic, 1973a), it is conceivable that the carboxylic-acid functional groups of saturated fatty acids might also be reduced, resulting in alkanes with the same number of carbon atoms as the original acids (Blumer, 1965; Welte and Waples, 1973). This latter reaction path would explain the otherwise unexpected presence of small amounts of evencarbon-numbered (c16, CIS, CzO) n-alkanes in Recent algal mats (Table 111). In intermittently aerated mats, unsaturated fatty acids are more likely to disappear into the insoluble kerogen fraction by way of oxidative cross-linking reactions (cf. Abelson, 1967, p. 70).

Stromatolites Fatty-acid distributions similar to those in extant microorganisms have

been reported from ancient algal sediments. Das and Smith (1968) identified the n-Clo:o to n-C20:0 acids (n-C16:0 being the most prominent) in five fossil algae (presumably preserved in carbonate), ranging in age from Eocene to Cambrian. Included in their suite of samples was the unicellular stromatolite- forming alga Chlorellopsis coloniata from the Green River Formation (Bradley, 1929a). Predictably, the level of the CI4, c16 and ClS unsaturated acids detected decreased steadily with increasing age so that a trace of the c16 unsaturated acid(s) was all that remained in the Cambrian Giruanella examined. These data cannot yet be interpreted in terms of algal taxonomy.

Recently, organic geochemists have turned their attention to the fatty acids and hydrocarbons which occur in the natural ecosystems and sediments of thermal environments (Jackson and Meinschein, 1973). Both non-biogenic stromatolite-like structures (geyserite) and biogenic (algal and bacterial) siliceous stromatolites are presently forming around hot springs and geysers in Yellowstone National Park (Walter et al., 1972). This particular hot-spring environment was postulated as a Recent analogue for the stromatolitic chert facies of the Gunflint Iron Formation by Walter (1972b). According to Jackson and Meinschein (1973), bacterial-algal mats in streams draining Yellowstone hot springs yield fatty-acid distributions similar to those obtained from ancient cherts. However, it is somewhat salutory to note that these workers also report finding syngenetic fatty acids in non-biogenic geyserite 102-103 years old; the acids have apparently changed little since being trapped in the silica. Nichols (1970, p. 114) quotes the fatty-acid composition of the important thermophilic blue-green algae Mastigocladus laminosus; apart from a minor amount (2.1%) of the c 1 8 : 2 acid, it lacks other poly-unsaturated acids. The tendency to synthesize a high proportion of saturated acids is a feature of organisms growing in warm environments (Abelson, 1967).


Pigm en ts

181

Po ten tial markers Quantitatively, pigments are the next most important part of the lipid

fraction of blue-green algae after the fatty acids and associated acyl lipids (Nichols, 1973). Algal pigments perform such vital functions as mediating photosynthesis, screening out harmful radiation, and preventing photo- oxidation. The major algal photosynthetic pigment is the dihydroporphyrin, chlorophyll a . The related bacteriochlorophylls are tetrahydroporphyrins which differ slightly in their peripheral structure. The iron-containing coenzymes (cytochromes, catalase, peroxidase) of microbial photosynthesizers are also porphyrin-based compounds.

Other key chromophores are the phycobilins (e.g., phycoerythrin, phycocyanin) present with chlorophyll in the photosynthetic lamellae of blue-green algae (Chapman, 1973) and the chloroplasts of red algae; and various carotenoids, some of which are unique to the Cyanophyta (Nichols, 1973). Carotenoids are cyclic or acyclic isoprenoid compounds, usually containing a series of conjugated double bonds.

The degradation of free, extracellular chlorophyll into porphyrins of the kind found in sedimentary rocks and crude oils has been documented by Orr et al. (1958). Chlorophylls (and other pigments) that are deposited in intact or partially degraded microbial cells may instead become “grafted” onto a variety of cellular macromolecules (Oehler et al., 1974). With further maturation these “grafted pigment complexes” may fragment into smaller pigment units; or alternatively condense into humic and fulvic acids, or insoluble kerogen. Blumer and Rudrum (1970) predict the presence of polymeric porphyrin structures in kerogen. Early diagenetic oxidation and polymerization of labile carotenoid pigments would likewise tend to result in their incorporation into a kerogen precursor, perhaps resembling sporopollenin (cf. Brooks and Shaw, 1970).

Stromatolites

Reports of fossil pigments in algally laminated sediments are few. Bac- teriochlorophylls a and c have been identified in organic matter separated from the Yellowstone stromatolites of .Recent age (Walter et al., 1972). Whereas trace amounts of presumably indigenous porphyrin-like compounds were found in sediments of the Swaziland Sequence, none could be detected in the stromatolitic Gunflint chert (Kvenvolden and Hodgson, 1969) despite its rich microfossil content. On the other hand, two Precambrian algal limestones examined by Hodgson et al. (1968) yielded chlorins and metal (Ni, V)- chelated porphyrins in small concentration. The presence of chlorins could indicate contamination from modem plants, soils or younger sediments. Chlorin pigments are very abundant in Recent sediments and soils.

182 D.M. MCKIRDY

Phycobilins are less stable to heat than chlorophyll (Oehler et al., 1974) and so recognition of their geochemical derivatives in ancient stromatolites may be more difficult. Jackson (1973) has nevertheless inferred that original differences in the ratio of phycoerythrin to phycocyanin in fossil algae are still decipherable in geolipid extracts of sediments as old as 3,400 m.y. He interpreted secular variations in the spectroscopic (visible and infrared) properties of soluble humic matter in sediments of various ages and depositional environments, in terms of major evolutionary changes in the pigmentation of the algal ecosystems populating those environments. Among the samples studied were cherts from the Gunflint Iron Formation, Paradise Creek Formation and Bitter Springs Formation, and a dolomite from the Beck Spring Dolomite, all stromatolite-bearing units. These four samples help define quite distinct trends in pigment concentration and/or composition over the 1,900-900-m.y. segment of Precambrian time. Evolutionary developments claimed to be reflected in these changes include: the presence of high concentrations of protective pigments in blue-green algal cells before 2,000 m.y.; a steady decline in algal production of protective chromophores (or a change in pigment type) between 1,900 and 900 m.y., corresponding to the gradual accumulation of atmospheric ozone; and the appearance of eukaryotic algae between 1,600 and 1,300 m.y. ago.

Amino acids

Carbonate-trapping and precipitating algal mats growing on the shelf of a hypersaline pool adjacent to the Gulf of Aqaba contain organic matter rich in amino acids (Friedman et al., 1973), notably aspartic acid (most abundant), alanine, glutumic acid, glycine, leucine, phenylalanine, serine and valine. These and twelve or so other amino acids are the building blocks of algal and bacterial proteins. Because of their lack of long-term geochemical stability the importance of such peptides and their constituent amino acids lies nbt so much in their value as biochemical markers per se but in their ability to modify mineralogical equilibria during early diagenesis, particularly in carbonate-depositing environments.

Peptides and free amino acids, like fatty acids, readily form monomolecu- lar layers on the surface of carbonate minerals (Suess, 1970), thereby slowing inorganic carbonate equilibration reactions, e.g. the inversion of aragonite to calcite (Jackson and Bischoff, 1971). Extracellular organic matter produced by blue-green algae is mainly peptide and its ability to chelate metal ions (e.g. Ca2+, Mg*+) is well known (Fogg and Westlake, 1955). Acidic amino acids such as aspartic acid and glutamic acid can bind Ca2+ ions from solution, and in so doing influence the precipitation of CaC03 (Mitterer, 1968; Mitterer and Carter, 1973). This may account for the calcitized organic matter in an algal laminate from the Middle Devonian Winnipegosis Formation (Shearman and Fuller, 1969). The high Mg2+ concentration in the


sheath mucilage (a polysaccharide-peptide complex: Dunn and Wolk, 1970) of stromatolite-building blue-green algae such as Schizothrix calcicola has been invoked (Gebelein and Hoffman, 1973) to explain the origin of dolomitic lamination in stromatolitic limestones. Similarly, Friedman et al. (1973) reported algally precipitated laminae of alternating aragonite and high-Mg calcite, the latter enclosing abundant dispersed organic matter. The organic matter actually contains one third of the total Mg in the laminates and is presumably the source of the Mg in the calcite.

The possibility that amino acids released from the original algal peptides by hydrolysis may still be present,, albeit in a less accessible form, cannot be entirely ruled out. Amino acids react with carbohydrates to produce dark, complex polymers having properties strikingly similar to natural humic acids (Abelson and Hare, 1971; Hoering, 1973). Humic acids are thought to be intermediates in the generation of kerogen. Once formed both humic acids and kerogen have the capacity to continue taking up free amino acids and peptides by adsorption as well as irreversible reaction (Abelson and Hare, 1970, 1971). Hence it would appear that any truly indigenous algal and bacterial amino acids are likely to be locked away in the kerogen fraction of the sediment.

Thus, although the peptides and amino acids of the precursor algae may not be amenable to preservation as recognizable biochemical markers, the effects of their original presence are often evident in the mineralogy and texture of the lithified stromatolite.

Kerogen

Potential markers Kerogen was once aptly and succinctly described by Degens (1967) as a

heteropolycondensate. Indeed, all of the biochemical markers so far discussed are capable of incorporation during diagenesis into this amorphous syngenetic substance, making it an important but still largely untapped reservoir of chemical fossils. The mucilaginous secretions of blue-green algae are probably a major progenitor of stromatolite kerogen (Shearman and Skipwith, 1965). Sporopollenin, the resistant and chemically inert carotenoid polymer claimed by Brooks and Shaw (1970, 1972) to be present in “many algal and fungal spore exines”, although recently reported in the cell wall of Chlorella and other green algae (Atkinson et al., 1972), has not yet been found in blue-green algae. Its contribution to the insoluble organic matter in fossil stromatolites is probably minimal and limited to those in which eukaryotic algae and/or fungi were part of the mat biota.

The main advantage kerogen has over the geolipid extract as a source of palaeobiochemical information, viz. freedom from post-lithification contamination, stems from its particulate, insoluble and relatively unreactive nature. These same attributes, however, make it difficult to analyse except by

184 D.M. MCKIRDY

relatively drastic degradative techniques involving hydrogenation, oxidation, alkali fusion or pyrolysis.

The products of low-temperature pyrolysis (Hoering, 1967; Giraud, 1970) or mi.ld oxidation (Burlingame and Simoneit, 1969) provide information about the basic structure of the kerogen, e.g. whether it is aliphatic or aro- matic. Thermally immature algal kerogens may be expected to yield hydrocarbons (on pyrolysis) and fatty acids (on oxidation) that are predominantly aliphatic (cf. Brooks and Shaw, 1970). In certain cases it might be possible to recognize more specific clues to the identity of the source material among the degradation products of stromatolitic kerogens. For example, the Cz0 and C22 unsaturated fatty acids characteristic of certain eukaryotic algae should give rise to anomalous amounts of CI9 to CZl alkyl residues; high concentrations of dicarboxylic acids following oxidation would be consistent with a sporopollenin-based kerogen (Brooks and Shaw, 1970). Comparison of the hydrocarbons obtained by low-temperature pyrolysis of the kerogen with those in the geolipid extract of the sediment provides a useful check on the syngeneity of the soluble fraction (cf. Hoering, 1967).

The elemental composition of kerogen broadly reflects the nature of its precursor organic matter and the environment of deposition of its host sediment (Tissot et al., 1974). However, the greater the degree of post-depositional thermal alteration the kerogen has undergone, the less diagnostic of both source and environment is its composition.

Stromatolites Hoering (1964) hydrogenated kerogen from a Precambrian (Bulawayan)

stromatolite with phosphorus and anhydrous hydrogen iodide‘ and analysed the resulting low molecular weight (< C,) hydrocarbons. As Hoering (1967, p. 103) points out, however, the method “is apparently too destructive to help reveal high molecular-weight compounds, diagnostic of a biological origin ” .

The elemental compositions of kerogen isolated from some Australian Palaeozoic and Precambrian stromatolites are given in Table IV. Their dry ash-free carbon content varies from 72.7% to 90.9% corresponding to a progressive increase in diagenetic rank (Fig. 5 : cf. Tissot et d., 1974, fig 2). Fig. 5 shows that algal kerogens vary widely in composition. Nevertheless, the stromatolitic kerogens are distinctive in that, as a group, they tend to be unusually rich in oxygen. This is highly indicative of an oxygen-rich precursor, e.g. algal mucilage. Three samples (4, 10 and 15, Table IV) actually plot outside the shaded compositional field. These may be derived from algal mats which grew in a more desiccated environment and hence produced greater amounts of mucopolysaccharides (and perhaps other compounds such as unsaturated fatty acids). Mucilaginous secretions help prevent loss of water from algal colonies during desiccation (Fogg et al., 1973, p. 320). The envelopes of desiccation-resistant spores from the blue-green alga Anabaena cylindrica contain 41% carbohydrate (Dunn and Wolk, 1970).


I 5 r

185

1 ;" I 0 2 Y = I

d a

V

V

INCREASING DIAGENESIS

\

v v a n

m n

V v a

a A A \ , \ \o'

8 B St romoto l i t e

0 Limestone

a Phorphor l te

v Shole

\ \

L I 1 0 0 10 0 20 0 30 0 40

Kerogen - OtomiC O/C ro t i o

Fig. 5 . Diagenetic evolution of stromatolitic and other algal kerogens. Data from Powell et al. (1975) , McKirdy and Powell (1976), and McKirdy (unpublished results).

Stable carbon isotopes

Potential markers The uptake and fixation of C02 by photosynthetic organisms leads to

fractionation between 12C and I3C such that the resultant organic matter is richer in the lighter isotope than the inorganic carbon source. For contemporary marine algae that source is dissolved CO, , and parameters such as C 0 2 concentration, pH, water temperature and growth rate have been shown to be major controls of the observed carbon isotopic composition (Degens, 1969). All photoautotrophic organic matter incorporated into sediments is stamped with an isotopic biochemical marker which is largely unaffected by subsequent diagenetic processes. In contrast, marine non-photosynthetic bacterial heterotrophs like Desulfouibrio desulfuricans are isotopically only slightly different from their (?algal) carbon source (Kaplan and Rittenberg, 1964). Should the sediment eventually undergo incipient metamorphism the 613C value* of its kerogen may become appreciably altered as a result of

* - 1 x 1000 1 13c/12c sample

13C/12C standard 613c per mil =

The Chicago PDB carbonate is a commonly used standard.

186 D.M. MCKIRDY

z I I METAMORPHOSED SEDIMENTS

1 NON-METAMORPHOSED SEDIMENTS

I 1 STROMATOLITIC CHERT

I 1 STROMATOLITIC CARBONATE

(-1 D E E P - S E A MARINE SEDIMENTS

I] HOT SPRING ALGAL MATS

PARALIC ALGAL MATS

r I LIPIDS OF MARINE PLANKTON ( 2 Z ' C l > a I ] MARINE PLANKTON I I ' C l

n d

+ a

8

LL o w 1- MARINE PLANKTON (22'CI

Irn W A THERMOPHILIC GREEN ALGA I I (rubrtmts COL, 259-500C1

46 THERMOPHILIC GREEN ALGA n lwbslrote o i r , 25'-5O'Cl

46 BLUE -GREEN ALGAE ~ [ ( r u b r l r o t a I 5% C02 In O K , 39-C)

PRECAMBRIAN - R E C E N T MARINE CARBONATES (calcite, dolomite1 W Z 00

3(2

o a In0 ATMOSPHERIC C 0 2 0 SEA WATER (HCO; +CO:-+ dissolved C O P ) r-1 m m

I I I I I I I I I 1 I - 4 5 -40 - 3 5 - 3 0 - 2 5 - 2 0 - I 5 - 10 - 5 0 + 5

s'3cPD, *i A8I3C data, where AS"C = S"C sample - SI3C substrate X M -157

Fig. 6. Stable carbon isotopic composition of contemporary algae, and algal organic matter in sediments (after Hoering, 1962, 1967; Weber, 1967; Degens, 1969; Behrens and Frishman, 1971; Oehler et al., 1972; Perry and Tan, 1972; Calder and Parker, 1973; Seckbach and Kaplan, 1973; McKirdy and Powell, 1974).

preferential rupture of "C- 12C bonds and loss of 12C-enriched, low molecular- weight hydrocarbons (McKirdy and Powell, 1974; cf. Fig. 6). There is good evidence that the carbon isotopic composition of atmospheric C 0 2 has remained relatively constant since the Early Precambrian (Weber, 1967; Perry and Tan, 1972). The 613C values of most ancient marine carbonates (limestone and dolomite) fall in a narrow (3O/,,) range which overlaps that for the inorganic carbon in present-day seawater (Fig. 6). Hence, the 6 13C values of living and unmetamorphosed fossil organic matter may be directly compared and used for interpreting the origin of the latter.

In laboratory cultures of blue-green algae and other photosynthetic microorganisms, the degree of fractionation between cellular carbon and inorganic substrate, A613C, is - 24 to - lo//oo (Abelson and Hoering, 1961; Calder and Parker, 1973; Seckbach and Kaplan, 1973; Fig. 6). The maximum fractionation was achieved by the thermophilic mat-forming green alga Cyanidium caldarium when growing in pure C02 at 45°C. Recent hot spring algal mats tend to be more enriched in 12C (6l3CpDB = - 24 to - llo/oo : Seckbach and Kaplan, 1973) 'than those inhabiting paralic environments (6l3CPDB = - 17 to -lOo/oo: Behrens and Frishman, 1971; Calder and Parker, 1973), although interestingly enough they are no more depleted in 13C than the organic matter (presumably phytoplanktonic detritus) in deep-sea marine sediments (613CmB = - 24 to - 16°/00 : Degens, 1969), as shown in Fig. 6.


Under conditions of restricted water circulation, CO, derived from the metabolic activity of mat-building algae, or from their decay, can be directly incorporated into the associated carbonate, particularly if the algae involved are carbonate precipitators. Then, both the carbonate matrix of the stromatolite and its organic matter will be enriched in 12C.

Stromatolites Hoering (1962, 1967) and Schopf et al. (1971) used carbon isotopic

measurements on coexisting carbonate and reduced organic carbon (Table V) to establish the biogenicity of abiophoric stromatolite-like structures in various Precambrian formations. As Schopf et al. (1971, p. 483) explain for the oldest known stromatolites: “The large difference in 6C13 values between

TABLE V

Carbon isotopic composition of some Precambrian stromatolites (after Hoering, 1962, 1967 ; Schopf et al., 1971)

Location 6 13cPDB

carbonate TOC

Belt Supergroup, Montana + 0.4 - 24.6

Transvaal Dolomite Series, S. Africa - 0 . 3 - 31.0 Bulawayan Group, Rhodesia - 0.1* - 32.1*

* Mean values.

Iron River Formation, Michigan -I- 1.3 - 19.9

coexisting organic and inorganic carbon compounds in the unmetamorphosed Bulawayan sediments is most reasonably interpreted as having resulted from carbon isotopic fractionation during the biological fixation of carbon dioxide by photosynthetic microorganisms”. Thus, carbon isotopic data, in conjunction with stromatolite morphology, demonstrates “the existence of photosynthesis and biological activity [early in] the Precambrian era” (Hoering, 1962, p. 191).

The carbon isotopic composition of kerogen in Precambrian stromatolites is typical of that found in other unmetamorphosed Precambrian (and most Phanerozoic) sedimentary rocks (Oehler et al., 1972; Fig. 6). However, the kerogen contains much less 13C than algal organic matter in Recent sediments. Various reasons for this have been advanced. Seckbach and Kaplan (1973, p. 168) speculated “that low 613C values measured in Precambrian kerogen may in part represent organic matter deposited in algal mats growing at elevated temperatures and under atmospheric conditions where Pco was substantially greater than at present”. Whereas similarities admittedly exist between the Yellowstone hot-spring environment, in which algal and bacterial stromatolites composed of silica are presently growing (Walter et al.,

188 D.M. MCKIRDY

1972), and the stromatolitic facies of the Gunflint and Biwabik iron formations (Walter, 1972b), thermophilic algae cannot be held responsible for the majority of Precambrian stromatolites which are preserved in marine carbonates. A possible alternative explanation is that lipids comprised a higher proportion of the Precambrian algal biomass and gave rise to lipid-rich kerogens. The lipids of modem organisms, including blue-green algae, are isotopically very much lighter than the whole organism (Abelson and Hoering, 1961; Degens, 1969; Fig. 6). Nevertheless, Calder and Parker (1973) maintain that incorporation of the lipid of contemporary blue-green algae into kerogen would not, by itself, produce the same low 6 13C values found in Precambrian organic matter. It seems that the correct explanation for the enigmatic isotopic composition of Precambrian kerogens lies in some as yet unknown combination of environmental variables.

The difference in kerogen isotopic composition between unmetamorphosed Precambrian stromatolitic carbonates (10 samples: 6 l3CpDB range = - 33.5 to - 19.9°/oo, mean = - 28.4°/00) and cherts (11 samples: 6 l3CPDB range = - 37.2 to - 24.8°/00, mean = - 29.0°/00) illustrated in Fig. 6 is probably too small to be of any potential palaeoecological significance, although analyses of more samples are obviously needed to fully explore this possibility. McKirdy and Powell (1974) found the kerogens from two Precambrian evaporitic carbonates, one of which was a stromatolite (sample 9, Table IV), to be some 7-8°/00 heavier than normal marine organic matter of similar diagenetic rank. On the same basis, it may be possible to distinguish peritidal from subtidal fossil stromatolites. The Boetsap River section of the stromatolitic Transvaal Dolomite (Truswell and Eriksson, 1973) would seem to be an ideal Precambrian locality in which to test this hypothesis.

Tan and Hudson (1974) report 613CpDB values as low as - 14.0°/00 for early diagenetic carbonate in algal limestones from the Great Estuarine Series (Jurassic) of Scotland. Restricted water circulation leading to a build-up of organicderived, '*Cenriched bicarbonate in their environment of deposition is implied.

ORGANIC GEOCHEMISTRY - A NEW TOOL IN STROMATOLITOLOGY?

If one thing is obvious from this account of the exploratory application of organic geochemical methods to stromatolites, it is perhaps that there have been more failures than successes. The composition of kerogen, the major portion of the organic matter preserved in stromatolites, is as obscure as ever. The syn'genetic origin of the extractable geolipids is still equivocal. But most important of all, stromatolites and their associated sediments have yet to be seriously considered as a source of biogeochemical information on the evolution of the mat-building biota and the diagenetic fate of algal and bacterial organic matter in carbonate-depositing environments.


Nonetheless, organic geochemistry has already contributed significantly to our understanding of stromatolites as sediments and the organisms that built them. Carbon isotopic measurements confirm the assignment of a photosynthetic (or chemoautotrophic) physiological mode to ancient stromatolitic biota. Where the geolipid fraction of the total organic matter appears to be largely syngenetic with the host rock, as is apparently more often true of carbonates than cherts, the hydrocarbons and fatty acids are consistent with an algal and/or bacterial derivation. The structure and composition of kerogen (e.g. hydrogen to carbon atomic ratio) provides an index of thermal alteration which can be used to monitor other more subtle diagenetic changes in stromatolites, particularly those affecting their microstructure. The interaction of organic matter in the form of fatty acids, amino acids and humic compounds with the surfaces of mineral grains helps explain such important phenomena as lack of recrystallization, preservation of calcite in stromatolite columns, and preferential dolomitization (or calcitization) of organic-rich laminae.

What then are some of the directions that future organic geochemical research might profitably take in the expandhig field of stromatolitology? Continued study of the lipids and other potential chemical fossils present in the organisms of modern algal-mat communities is essential because it is on such biochemical data that interpretation of the molecular fossil record is ultimately based. Particular attention should be paid to compounds that are peculiar to certain classes of microorganisms, viz. coccoid and filamentous blue-green algae, eukaryotic algae, fungi, and photosynthetic and non-photosynthetic bacteria. Non-photosynthetic bacterial decomposers in algal mats are of considerable geochemical importance because they are the last organisms to leave their biochemical imprint (e.g. singly branched fatty acids, high proportion of > Czo compounds in their aliphatic hydrocarbons, preferential metabolism of algal straightchain hydrocarbons) on the residual organic matter. Mats from different environments should be compared to ascertain whether characteristic differences in their “biochemical facies” exist. Several organic geochemical parameters are to a large extent environmentally controlled: pristane to phytane ratio and even-predominance in n-alkanes, by redox potential; and kerogen 613C, by temperature, pH and COz concentration. In view of the ability of carbonate minerals to form complexes with a wide variety of organic compounds, stromatolitic carbonates appear likely to be a more fruitful source of syngenetic chemical fossils than cherts. Finally a concerted attempt, employing a range of analytical techniques (including carbon isotopic analysis), should be made to characterize as many stromatolite kerogens as possible, with the aim of detecting original bio- chemically derived (as distinct from geochemically imposed) differences which may be of evolutionary significance.

Now that biological control of stromatolite microstructure (Gebelein, 1974) and some aspects of gross morphology (Serebryakov and Semikhatov,

190 D.M. MCKIRDY

1974) has been verified, the potential rewards awaiting further study of stromatolites by organic geochemists would seem to far outweigh the present analytical and interpretational difficulties.

ACKNOWLEDGEMENTS

This paper would not have been written without the patient encouragement of Malcolm Walter who freely shared his knowledge of stromatolites during many discussions. Drs K.A. Kvenvolden, K.S.W. Campbell and D.J. McHugh kindly read the draft manuscript and made helpful suggestions for its improvement.

Samples of stromatolites were donated by Drs M.R. Walter and W.V. Preiss. Mr Z. Horvath, Petroleum Technology Laboratory, B.M.R., provided skilled technical assistance. The organic geochemical data on Australian stromatolites was collected as part of the author's Ph.D. project at the Geology Department, Australian National University, Canberra. The support of an Australian Public Service Post-Graduate Scholarship for this work is gratefully acknowledged.

The author publishes with the permission of the Director, Bureau of Mineral Resources, Geology and Geophysics, Canberra, A.C.T.

NOTE ADDED IN PROOF

Two recent papers on the organic geochemistry of algal mats from Baja California provide important new clues to the likely composition and derivation of stromatolitic lipids and kerogen. Cardosa et al. (1975) reported that hopane-type triterpenes and triterpanes are prominent in the lipid fraction, reflecting the predominantly prokaryotic nature of the mat biota. The extractable fatty acids include a homologous series of normal saturated acids in the range C14 to Cm with maxima at c16 and c 2 6 . This finding corrects an impression conveyed by earlier studies (summarized in Table 111) that long-chain (>C,,) fatty acids do not exist in Recent algal mats. Such acids constitute a plausible source for the long-chain n-alkanes isolated from ancient stromatolites. Treatment of the kerogen with chromic acid (Philp and Calvin, 1975) yielded oxidation products suggestive of a highly aliphatic structure. The kerogen appears to consist of a network of cross-linked polymethylene chains which probably formed by the condensation of unsaturated oxygenated lipids, including the phytyl moiety of chlorophyll.

Eichmann and Schidlowski (1975) found the mean difference between the 6 13C values of inorganic and organic carbon in a group of 22 Precambrian stromatolitic carbonates to be about 3 ' / ~ greater than the average for Precambrian carbonates, but they were reluctant to attribute any significance to the discrepancy.

2 APPENDIX 0

Australian stromatolites for which new analytical data are presented in the text r 0

Sample Reference Formation Location Estimated Form Lithology Source 0 No. age (m.y.) 4

n 1 CMC ao 2 CMC 68 3 CMC 37

4 S 105

5 WP 22

6 S 418

7 S 388

8 S 147

9 S 405

10 S 172

11 0 9

12 0 8

13 WP 64

14 MRW 12 (2/9/73)

15 s 95

Flaggy Limestone Member, Cavan Limestone. Munumbidgee Group

Wilkawillina Limestone, Hawker Group

Wonoka Formation, Wilpena Group

“Etina Formation”. Umberatana Group

Tapley Hill Formation. Umberatana Group

Loves Creek Member. Bitter Springs Formation

Skillagalee Dolomite, Burra Group

? Skillogalee Dolomite, Burra Group Tooganinie Formation. Mc Arthur Group

Pillingini Tuff, Fortescue Group

Taemas. N.S.W.

Wilkawillina Gorge, S.A.

Bunyeroo Gorge, S.A. Kulpara, S.A.

crenulate 375 linked cumulate

565

(Early Dev.) stratiform

(Early Cambrian) cumulate

600 Tungussia inna

650 Kulparia kulparensis

Wilson. S. A. > 700 Gymnosolen ramsayi

Williams Bore. N.T.’

Depot Creek. S.A.

Depot Creek, S.A.

Depot Creek, S.A.

800 Acaciella austral ip

Baicalia burra

800 ? Tungussia wilkatanna

cryptalgal laminate

carbonate carbonate carbonate

carbonate

carbonate

carbonate

carbonate

carbonate

carbonate

carbonate chert

w author author

v author

0 M.R. Walter E

W.V. Preiss 0

4 W.V. Preiss

W.V. Preiss

2

% E: E

M.R. Walter

W.V. Preiss

W.V. Preiss author

Mundallio Creek, S.A. cryptalgal laminate chert author unknown

Top Crossing. N.T. 1,600 Conophyton f. carbonate M.R. Walter

Mount Herbert, W.A. 2.200 Alcheringa narrina carbonate M.R. Walter

carbonate W.V. Preiss ? 800 ? Baicalia f.


Chapter 5.1

THE ORIGIN AND DEVELOPMENT OF CRYPTALGAL FABRICS

C.L.V. Monty

INTRODUCTION

The multiplication of studies on carbonate-building non-skeletal algae has led to the discovery of a wide range of algal structures which cannot always be called “stromatolite” if we want to keep the meaning of this word within reasonable limits; that is why, following Aitken’s definition (1967), the term cryptalgd will be used to designate all these biosedimentary structures which originated through the sediment-binding and/or the mineral-precipitating activities of non-skeletal algae, mainly blue-green algae, and of bacteria.

Such cryptalgal structures are quite diversified in the geological column, and the purpose of this chapter is not to account for all of them, but just to illustrate, with examples from the Recent, the way some of their fabrics originate.

The study of cryptalgal fabrics is but one of the multiple approaches to stromatolitic carbonates, an approach which ranks third in the hierarchy of characteristics of these sediments:

the stromatolite bioherm or biostrome

refers to the overall deposit

the stromatolite* refers to the individual cryptalgal structures constituting the deposit

the fabric refers to internal spatial properties of these structures such as the development of a lamination refers to the microscopic characteristics of the internal properties refers to the habit and arrangement of cryptocrystalline units

In a first approach, cryptalgal fabrics may be divided into two main groups:

the m icrostruc tu re

the ultrastructure

* Sensu Kalkowski (1908)

194 C.L.V. MONTY

(1) the layered fabrics and (2) the non-layered fabrics according to the presence or the absence of a conspicuous biosedimentary layering.

LAYERED CRYPTALGAL FABRICS

General features

Layered fabrics characterize three main types of cryptalgal deposits : Stromatolites, i.e., domal, club-shaped, columnar, etc., structures characterized by a generally non-planar lamination ranging from a few microns to a few millimeters, possessing welldefined threedimensional boundaries and which grow attached to the substrate. Oncolites, i.e., unattached subspherical, ovoid, lobate or flattened structures otherwise similar to stromatolites. Cryptalgal laminites, i.e., laterally continuous stratiform deposits characterized by a subcontinuous planar lamination ranging from less than a millimeter to several centimeters.

Two main types of layered fabrics can be recognized: (1) The laminated fabrics are characterized by a prominent microstratifi-

cation resulting from the superposition of well-defined individual layers or laminae of similar or different composition, microstructure, origin and/or color, separated, in most places, by a clear physical discontinuity, or micro- diastem. This subtle notion of “discontinuity” should however be properly appreciated; in many cases the laminated fabric does not result from the stacking of individual laminae, but, for instance, from rhythmic changes affecting such features as the mineralization of an otherwise continuous algal thallus, producing stripes imprinted on a continuous framework; when this can be shown, it may be useful to speak of a striped laminated fabric.

( 2 ) The laminoid fabrics show no laminae, but the presence of well or poorly aligned structural features which tend to outline an overall layering. Among other things, laminoid fabrics may result from: (a) the repetitive development of generally elongate primary or penecontemporaneous cavities or fenestrae (Tebbutt et al. 1965), open or filled with sediment and/or cement - the laminoid fenestral fabric; (b) a non-gravity-driven common orientation of elongate particles parallel to an overall lineation which pervades the cryptalgal body, showing that they have been bound by, or glued to, successive algal mat surfaces - the laminoid boundstone fabric; and (c) an irregular juxtaposition and superposition of algal cushions or mats much smaller than the width of the stromatoid -the lenticular laminoid fabric.

The layering, which is one of the basic features of stromatolites, records a rhythmicity which affects the development of the constitutive algal population and/or one or several environmental parameters. The characteristics of the rhythmicity determines the organisation of the laminations and

FABRIC AND MICROSTRUCTURE 195

.......................

I D

C -Bt

E F

Fig. 1. Illustration of some frequent modes of organization of laminae in stromatolites: A, simple repetitive lamination; B, C, composite repetitive lamination; D, simple alternating lamination; E, composite alternating lamination; F, cyclothemic lamination.

the following modes can be distinguished : (1) Repetitive lamination. Simple: superposition of laminae of similar

nature and configuration, separated by physical discontinuities (Fig. 1A). Composite: superposition of groups of laminae separated by physical discontinuities which determine the prominent lamination; may also result from the superposition of microstromatolites (Fig. lB, C).

( 2 ) Alternating lamination. Simple: alternation of two types of laminae texturally and/or mineralogically different (Fig. 1D). Composite: alternation of two types of layers, one or the two of which contain a second-order lamination (Fig. 1E).

( 3 ) Cyclothemic lamination. Succession of at least three different laminae which always appear in the same order and which can be grouped into genetic sedimentary units (Fig. 1F; Monty and Hardie, Ch. 8.6, Fig. 8).

Besides the mode of organization of the lamination, the microstructure, the configuration, the linkage, the spacing, the relief of individual laminae are other important properties that have been reviewed and classified by Hofmann (1969a, Figs. 8,13) and will not be considered here.

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Origin of layered cryptalgal fabrics

A number of biological, geochemical, physical and sedimentological factors account for the development of lamination in cryptalgal structures. Some of them are explained here to illustrate the diversity of the processes involved.

Phototactic response of the component organism Many motile microorganisms are characterized by daily movements in

accordance with diurnal light variation. Such phototactic movements can induce the development of a primary biological lamination. Interesting examples have been described in colonies of Schizothrix calcicola sensu Drouet (Monty, 1965b, 1967; Gebelein, 1969) now classified as Phormidium hendersonii Howe (Golubic and Focke, 1976). During the day, upward gliding trichomes produce an organic lamina, up to 900pm thick, made of erect bundles of filaments; at night, filaments grow prostrate and at a much slower rate forming a thin dark organic lamina (100 pm thick) (Fig. 2A, B).

This biological lamination may be outlined by sedimentary particles in environments characterized by low, more or less continuous sedimentation rates: during the day, the rapidly growing mat incorporates detrital particles that will remain scattered so that a definitely hyaline layer is formed; at night, on the contrary, as the filaments grow horizontally and much more slowly, the detrital particles can be concentrated in a well-defined corpuscular lamina (Fig. 2C, D). During the next day, the filaments will permeate and bind these grains firmly to grow a new hyaline lamina. The development of such organosedimentary laminated fabric relies on a balance between algal growth and sedimentation rate: if the latter is too low, a purely organic bio- laminated structure is formed; if on the contrary it is too high, a massive boundstone is formed (Fig. 2E).

Similar phototactic responses have been observed by Walter et al. (Ch. 6.2) in the case of siliceous flat-topped stromatolites built by Phormidium tenue var. granuliferum Copeland in Yellowstone National Park; laminae are however much thinner here. A reverse situation has furthermore been found by

Fig. 2. Laminated domes and mats built by Phormidium hendersonii (formerly Schizo- thrix), Windward Lagoon, Andros Island, Bahamas. A. General view of biolamination in a sediment-poor colony; thin dark films, built at night,

alternate with wider light laminae formed by erect phototactic filaments during the day. Scattered detrital particles appear black. Plane polarized light. Scale bar = 1 mm.

B. Detail of biolamination. Plane polarized light. Scale bar = 100pm. C. Section of a dome in which the balance between algal growth and sedimentation rate

originated a well-laminated fabric due to concentration of detrital particles in biolaminae built at night (see text). View of impregnated slab. Scale bar = 400pm.

D. Close up of Phormidium filaments binding detrital particles. Plane polarized light. Scale bar = 100 pm.

E. Section of a dome swamped by intensive sedimentation. Basic biolamination is no longer outlined but is obliterated by oversedimentation, and a massive fabric tends to develop. Plane polarized light. Scale bar = 500pm.

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Doemel and Brock (1974) in stromatolites built by the photosynthetic filamentous bacterium Chloroflexus: as this bacterium responds to low light and darkness rather than to high light intensities, it migrates upward during the night, building a lamina about 100 pm thick, and grows horizontally during the day.

Alternating growth pattern of the component organism Some algae have a thallus combining both horizontal and vertical growth.

This is for instance the case for Scytonema (Monty, 1965a, 1967; Golubic, Ch. 4.1, fig. 14). A laminated pattern can thus result from the alternation of phases showing horizontally growing filaments, with phases characterized by vertical bundles of filaments (Fig. 3).

Fig. 3. Vertical section in mat built by calcified filaments of Scytonema, subterrestrial, eastern Andros Island, Bahamas. Biolamination results from alternation of horizontally growing filaments with layers of erect bundles. Plane polarized light. Scale bar = 500pm.


Periodic differentiation of an algal assemblage leading to an alternation of dominant algae Not infrequently, algal mats contain a pool of organisms which do not

bloom at the same time or in the same conditions. This enlarges the adaptive range of the community and permits its perpetuation; by the same token, the periodic differentiation of one or another dominant organism may produce a laminated fabric in the mat.

Interesting illustrations of such complex behavior can be found in the laminated mats built by the Scy tonema-Schizothrix calcicola association on eastern Andros, Bahamas (Monty, 1965a, 1967); these algae bloom in different conditions, as Scy tonema is a freshwater terrestrial alga whereas Schizo- thrix calcicola is more successful in flooded conditions and withstands saline waters.

In vertical section, these mats and domes invariably show an alternation of whitish calcareous layers and brownish organic ones. This constant fabric is however genetically very complex and may result from at least three different behavior patterns of the Scy tonema-Schizothrix community, hence of three different mat structures:

(1) A first type of mat (Monty, 1967, type 3) (3-4cm thick, 20-50cm wide) is composed of calcified layers from 5mm thick at the surface com- pressing to 1 mm thick at depth, built mainly by calcareous tubes of Scyto- nema; this alternates with hyaline layers (30-300 pm) made by dense growths of interwoven filaments of Schizothrix (Fig. 4E). These mats were found growing on rocks and on the mud of sinkholes bordering the Fresh Creek (Andros Island). Their genesis can be traced as follows: during the dry season, Scytonema grows at the surface of the mat either erect and vigorously if substratal moisture is important or sluggishly and oblique during drier intervals. At this stage, Schizothrix concentrates beneath the surface, finding in between the bundles of Scytonema the moisture which enables it to survive. A calcified layer is then built; it results from the calcification of the sheaths of Scytonema and from the abundant precipitation of calcite in the intervening mucilages and the overgrowing felts of Schizothrix.

When the mats are temporarily soaked by saline waters (* 20-30°/,) during a storm or an equinoctial high tide, Schizothrh rises to the surface of the mat and blooms; it does not calcify in such waters, however, and a loose organic layer is formed. After such an accident the salinity is progressively reduced by ground water and rain, and the mat progressively returns to subaerial drier growth; Schizothrh recedes while Scy tonema starts producing a new calcified layer. The organic layers compact rapidly at depth where they are reduced to brownish subhorizontal wavy streaks.

(2) A second type of mat (Monty, 1967, type 2), built by the same community, shows the reverse situation, in which calcified layers are built by Schizothrix whereas the hyaline layers are built by Scytonema. In vertical section, it demonstrates an alternation of composite, micritic laminae

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(40-400 pm thick) crowded with Schizothrix, alternating with irregular hyaline layers (20-800 pm thick) built by erect bundles of Scytonema (Fig. 4B). In August 1964, such structures were observed actively blooming in sinkholes along the south shore of Fresh Creek in 1 or 2 cm of brackish water (10-13°/oo) : here, the tops of submerged older algal polygons were covered with greenish brown mats, 2-4 cm thick, composed of a pile of wrinkled films of Schizothrix (about 10pm thick) between which many gas bubbles and eventually colonies of unicells were entrapped. The films tended to be grouped into sets of 4-8, separated by thicker hyaline laminae (Fig. 4A). Each film was made of horizontal,tangled felts of Schizothrix loaded with micron-sized calcite crystals (Fig. 4F). The uppermost films contained motile reproductive units (hormogones) of Scytonerna, as well as very short filaments and isolated cells of other algae (Fig. 4D, F). On the other hand, the underlying films and interspaces were inhabited by fully mature algae including, in the wider interspaces, long greenish to colorless erect filaments of Scytonerna. This microstructure reflects the following genesis: when the dormant mat is soaked or flooded by slightly saline waters, Schizothrix finds there ideal conditions, and blooms. By active division and phototropism it starts building subdaily films which pile on top of each other while calcium carbonate is precipitated in the mucilaginous felts. Metabolic gas bubbles accumulate between the films and distort the microstructure a little. The associated algae (Scytonerna and some Lyngbya) , characterized by a generally much slower growth rate, respond to burial under several hundred microns of calcite and mucilage by sending hormogones upward. These, as well as young unicellar stages, settle in and between the films where they

Fig. 4. Laminated mats built by Scytonema and Schizothrix, subterrestrial, eastern Andros Island, Bahamas. A, B, C, D, F, mats controlled by Schizothrix, E, mat controlled by Scytonema. A. Rapid succession of thin, calcite-rich films of blooming Schizothrix (see F) forming a

set overlying a hyaline layer crowded with Scytonema. Plane polarized light. Scale bar = 50pm.

B. Detail of resulting lamination: calcitic laminae formed by compacted sets of films of Schizothrix alternate with hyaline layers of internally growing non-calcified filaments of Scytonema. Plane polarized light. Scale bar = 100pm.

C. General view of resulting lamination. Some pellets are formed in situ by calcification of colonies of unicells and diatoms, some are left by burrowing organisms. Plane polarized light. Scale bar = 200pm.

D. Growing hormogones sent by Scytonema into the Schizothrix surficial bloom. Plane polarized light. Scale bar = 25 pm.

E. Section in mat controlled by Scy tonema : calcareous layers, made of erect calcified filaments of Scytonema, alternate with hyaline discontinuous ones built by Schizothrix. Plane polarized light. Scale bar = 500 pm.

F. Smear of a surficial film of blooming Schizothrix: (see A) eu- and anhedral low-Mg calcite crystals are precipitated in the felt of filaments. In the upper right corner, there is a young stage of the blue-green Johannesbaptistia pellucida (Dickie) Taylor et Drouet. Plane polarized light. Scale bar = 20 pm.

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soon divide; thus progressively more mature stages are found from the surface downward. Blooms of Schizothrix are, however, not continuous, as seen in subtidal counterparts (Monty, 1967, p. 91); they may stop for some time during which the surface becomes colonized by “competing” algae; after a while, Schizothrix growth resumes and a new set of calcified films is formed. In the meantime, the young Scytonema which have settled in between the films, or rather the sets of films, keep on growing and expand these spaces into hyaline layers; as observed by the author (see also Schonleber, 1936), when Scytonema grows under water or separated from the air interface it may not calcify (see also below). During growth, the superposed films of Schizothrix progressively compact into thicker micritic composite layers, as the metabolic gases escape, while the intervening mucilages are consumed. A laminated mat composed of calcified and hyaline organic layers is formed (Fig. 4C).

(3) A third type of mat, very common in the freshwater lakes of Andros (Monty, 1972), can be built during another behavior of the community. These mats show an alternation of calcified and hyaline layers built by Scytonema and Schizothrix (Fig. 5 ) . A cross-section of the living mat, examined at the end of the seasonal flood when algal structures emerge, shows the living portion, up to 3 cm thick, resting on a subfossil “basement” made of alternating light tan calcareous laminae and brownish organic ones. The living portion can be split into two parts. There is an upper spongy calcareous lamina up to 2 cm thick, made of calcified long vertical bundles of brownish Scytonema which may project above the surface, and which is zoned by horizontal calcified film of Schizothrix; the overall original fabric is thus reticulate. The lower part shows a loose green layer up to 1.5 cm thick with erect non-calcified filaments of Scytonema intermingled with films and draperies of Schizothrix delineating large vacuoles. Field observations show that the calcified lamina forms when, as the freshwater level drops, the mat emerges or is about to. In such conditions, Scytonema actually starts its freshwater subaerial growth phase and builds long filaments encased in a calcified sheath; as the mat is still fully soaked with water Schizothrix can equally grow and concurrently build its calcified films, so that a reticulate pattern is formed; spaces between the successive films can get wider and wider as the Scytonema filaments separating them grow. Beneath this surficial layer, life conditions for both algae are still good as shown by the deeply pigmented green filaments and the abundance of gas vacuoles; but conditions are not suitable for calcite to develop. (In this somewhat closed environment choked by mucilage and the overlying layer, any carbonate precipitated during the day may dissolve at night when the cells are only respiring.) This *layer remains mainly organic, except for scattered flocs of calcite. At the end of the growth period, both laminae collapse and compact severely (at least ten times) and the habitual laminated fabric of thicker calcareous layers separated by thinner organic ones is formed. In this case, as


Fig. 5 . Section in a dome built concurrently by Scytonema and Schizothrix, freshwater marshes, eastern Andros Island, Bahamas. Gray to white calcareous layers show a reticulate microfabric resulting from overlap of erect calcified bundles of Scytonema, perpendicular to growth surface, with calcified films of Schizothrix parallel to this surface. Dark hyaline layers are built by internally or subaqueously growing phases of Scytonema. Note severe compaction of laminae inward. Polished impregnated slab. Scale bar = 1 cm.

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in the first one, fragments of calcareous sheaths can still be found, but the original reticulate microfabric is completely obscured.

These three cases show how subtle variations in moisture and salinity can contra1 differentiation of an algal consortium to finally originate a similar laminated fabric, although through different pathways.

Internal growth Geologists are prone to interpret the succession of stromatolitic laminae in

stratigraphical terms, i.e., a lamina is younger than the one it overlies. The last two examples of the previous paragraph have shown that this is not always true and that algal layers can grow inside the mat. Many other examples have been found on Andros Island (personal observations).

Generally, in algal mats submitted to intense insolation, the surficial layer becomes strongly pigmented to shield the cells from radiation, and/or very

Fig. 6 . Rivularia haematites, Lunzer Untersee, Austria. Development of striped laminated fabrics by periodic calcification. A. General view of a small colony showing dark concentric calcified stripes. Plane

polarized light. Scale bar = 500pm. B. Detail of striped fabric; elongate, tubular, calcite crystals, oriented on the filaments,

form conspicuous bands appearing dark. Interference contrast. Scale bar = 50 pm. C, D. Detail of calcified zones; C, view parallel to the filaments of elongated crystals sur-

rounding the filaments; D, oblique view. Cross polarized light. Scale bars = 100 pm.


compact to prevent excessive evaporation. In this case it is not rare that, in zones where the substratal humidity concentrates, hormogones or dormant cells start actively to grow, originating the development of interstratified organic layers and a thickening of the mat from the inside (Monty, 1965b).

Periodic calcification of dominant alga Periodic calcification of blue-green algal mats relies on patterns and pro-

cesses which are still poorly understood today. If in some structures, one type of crystal habit seems to be preferentially deposited, in others a variety of calcite crystals appear to be simultaneously precipitated in the mat (see for instance Monty and Hardie, Ch. 8.6, Fig. 4). In calcareous freshwater streams and lakes, a collection of blue-green algae are building laminated crusts or cushions resulting from the periodical deposition of calcite crystals organized into distinct concentric layers; this is the case for several species of

Fig. 7. Development of striped or laminated fabrics by periodic calcification. Section of a laminated fluviatile calcareous crust built by Phormidium incrustatum (Ruisseau du bois d’Haumont, Belgium). Lamination results from alternation between densely calcified zones and loosely calcified ones (see detail on Fig. 8D). Aligned fenestrae ( f ) are tubes of chironomids. Plane polarized light. Scale bar = 200 pm.

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Schizothrix, Phormidium, Calothrix, Rivularia, etc. (e.g., Fritsch, 1929,1949, 1950; Geitler, 1932; Kann, 1933, figs. 7, 8, 11, 1941b). Two examples will be briefly illustrated here.

The first concerns R iuularia haematites De Candolle, which builds strongly calcified dome-shaped colonies up to 3 cm across and to 3 cm thick (Fritsch, 1950) in the splash zone of calcareous lakes. A cross-section in the colony reveals its typical striped fabric (Fig. 6A, B): the thallus, made of radially growing filaments, shows concentric bands of interlocking elongate crystals encasing the algal sheaths and alternating with non-calcified zones (Fig. 6C, D; see also Wallner, 1935). There are normally 3-4 calcified zones in colonies (Kann, 1933) but Geitler (1932) has counted up to 33 zones in old ones. The reasons of this periodic precipitation are not yet clearly known.

The second example refers to crusts built by the Phormidium incrustatum (Nageli) Gomont community in calcareous rivers (Fritsch, 1929 et seq.). In vertical section (Fig. 7), these crusts, up to 10cm thick, show a prominent lamination resulting from alternating densely calcified and loosely calcified filamentous layers (Fig. 8D). Palynological studies on Belgian crusts (in the Ruisseau du bois d'Haumont; Geurts, 1976) permitted dating of the respective layers and showed that the lamination was seasonal. Between February and May, when the water is cold (ca. 4"C), Phormidium blooms and grows closely adpressed, radial filaments to form coherent calcareous layers up to 5 mm thick; the component filaments are heavily incrusted with small equant or rhombic crystals of low-Mg calcite. Then, as the water gets warmer (up to 16"C), between May and January, growth decreases and the filaments aggregate into loose and scattered calcified bundles separated by wide cavities.

~-

Fig. 8. Microfabric variations in laminated crusts of Phormidium incrustatum. A. Detail of lamination showing densely calcified flabellate tufts of Phormidium overlain

by a loosely calcified layer, rich in detrital particles; films of Schizothrix form scattered concentric dark zones in the Phormidium lamina (second-order lamination). Here, calcification involves the precipitation of interlocking, elongate crystals on the filaments of Phormidium (see B, C) as in Riuularia (Hoyoux river, Belgium). Plane polarized light. Scale bar = 500 pm.

B. Oblique section through crust illustrated in A, showing the interlocking crystals encasing the Phormidium filaments in two calcified layers sandwiching a poorly to non- calcified one (which appears black). Moulds of the filaments appear as black dots in the lower part of the photograph whereas longitudinal sections are seen in the upper one. Cross polarized light. Scale bar = 100 pm.

C. Calcified tuft of Phormidium, illustrated in A, showing the elongate oriented crystals. Cross polarized light. Scale bar = 500pm.

D. Detail of lamination of the Phormidium head illustrated on Fig. 7. From base to top: end of heavily calcified spring layer, passing to fenestrate loosely calcified summer layer, which grades into another densely calcified growth. Although growth appears continuous here, many cases exist where clear-cut discontinuities separate the different seasonal layers. Calcification results here from precipitation of small anhedral calcite crystals clustering around and between the filaments. Plane polarized light. Scale bar = 200 pm.

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Accordingly, a doublet, formed by a strongly calcified layer followed by a loosely calcified one, is deposited each year and a radial filament microstructure is generated in the component laminae (Fig. 8D).

Precipitation of crystals with different habits Seasonal variation in the algal coenose may originate the alternate precipi-

tation of very different types of calcitic crystals. This is for instance the case for particular crusts built by Phormidium incrustatum and Schizothrix sp. in the Hoyoux river, Belgium. Here, the calcification of the Phormidium layer during each spring appears to be much stronger than in the case reported above: instead of being surrounded by a microcrystalline sheath the filaments are encased in tubular interlocking elongate crystals (Fig. 8B, C) as was the case for Riuularia. Because of the flabellate, radial, growth of the filaments (Fig. 8A) these crystals tend to form juxtaposed or intertwined composite fans. Such crusts may be periodically invaded by blooms of Schizothrix (minor invasions are shown on Fig. 8A) associated with the precipitation of small (less than 10pm) anhedral calcite crystals. Consequently, a vertical section of the crust will show a prominent lamination resulting from the alternation of microcrystalline laminae with layers of radiating elongate crystals which originates a crystalline laminated fabric (Fig. 9).

Primary crystalline fabrics may be even more complex as is shown by crusts up to 20 mm thick built, in the same river, by a community including small oscillatoriaceans, abundant coccoids, and bacteria. The basic fabric, that may be called crystalline laminoid fabric (Fig. lOA), results from the alternation of rather porous layers, made of coarse-bladed crystals (up to 80 pm long and 30 pm wide) growing in all directions, with compact microcrystalline laminae (Fig. 1OC). Continuous observations revealed that clusters of large platy crystals, interspersed with mucilaginous matter and aggregates of unicells, were initially growing in horizontal subcontinuous cavities roofed by the living algal film, rich in coccoids, and floored by the older parts of the crust over which the organic film draped loosely (Fig. 10B). Such crystal aggregates probably formed by bacterial action on algal mucilage (W. Krum- bein, pers. comm., 1975). As precipitation proceeds, the clusters of crystals become larger and more numerous to finally interlock and form a crystalline lamina such as illustrated on Fig. 1OC. These coarse-grained layers alternate with microcrystalline laminae deposited periodically in association with Schizothrix as in the Phormidium crusts (p. 207). Furthermore, erratically or along given horizons, the coarse crystalline laminae are invaded by microcrystalline equant crystals which are deposited on the blades and tend to fill all the interstices, obscuring the coarse crystalline fabric; these microcrystals are here also a'ssociated with small oscillatoriaceans which colonize and overgrow the surfaces of the large platy crystals. All these interactions account for the resulting poor lamination of the crust (Fig. 10A).


Fig. 9. Development of crystalline laminated fabric in a crust built by alternation of Phormidium layers, showing fan-shaped tufts of elongate crystals (as in Fig. 8C) with microcrystalline laminae (M.L. ) associated with Schizothrix (Hoyoux river, Belgium). Cross polarized light. Scale bar = 100pm.

Periodic enrichment in various chemicals Independently of the situation found in manganese nodules, cryptalgal-

laminated or striped fabrics may result from periodic enrichments in various mineral compounds of, for instance, iron and manganese. This is the case for the crusts reported by Kann (1940) from the shores of the Kellersee. These crusts are built by an intermingled continuous growth of Phormidium

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incrustaturn and the red alga Chantransia incrustans (Hansgirg); the lamination results from alternating light and dark bands, the latter representing periods of important enrichment of the mat in iron hydroxide (which may be related to the iron-rich springs surrounding the lake).

Similar bandings have also been reported by Naumann (1924,1925) from globular colonies of Nostoc colonizing the surficial waters of volcanic African lakes; in the larger colonies, these bandings are repeatedly found every two millimeters and result from a heavy impregnation of the mucilaginous algal sheaths by ferric hydroxides. Finally, periodic enrichment in iron and manganese have also been reported by Howe (1931) from crusts and biscuits built by blue-green algal unicells in a creek in West Virginia, U.S.A.

Periodic influx of detrital particles Many blue-green algae, mainly oscillatoriaceans like Schizothrix, Phormid-

iurn, Microcoleus and Lyngbya are motile and are positively phototactic. They usually build patchy or subcontinuous mats and, because of their motility, they can glide up through a layer of sediment recently deposited on them to re-establish a new surface mat, leaving empty sheaths, mucilaginous matter and dead cells behind. This way laminated fabrics can result from concurrent algal growth and sediment deposition. This has been illustrated previously in the case of discrete colonies (p. 197; see also Gebelein and Hoffman, 1968; Gebelein, 1969).

Algal mats growing in shallow subtidal waters (Dalrymple, 1965) and particularly on intertidal and supratidal flats are faced with various periodical influxes of detrital particles, ranging from daily to seasonal and yearly (e.g., Ginsburg et al., 1954; Hommeril and Rioult, 1965; Kendall and Skipwith, 1968; Davies, 1970). The thickness, texture and composition of the various constitutive corpuscular layers will vary as a function of the hydrodynamical agents which deposited them (e.g., tides, storm waves and hurricanes) as these agents have different frequencies and capacities, act during shorter or longer periods of time, and may drain particles from different regions. Furthermore, the superposition of these events will eventually produce several orders of laminae grouped into cycles (Dalrymple, 1965, fig. 40; Davies, 1970, fig. 11).

Fig. 10. Development of crystalline 1arninoid.fabric in crusts built by unicells and oscillatoriaceans (Hoyoux river, Belgium). A. General view of part of the crust showing coarsely crystalline layers interspersed with

microcrystalline laminae. Plane polarized light. Scale bar = 500 ym. B. Incipient coarsely crystalline layer. Elongate calcitic blades are being precipitated in

the mucilaginous interspaces crowded with unicells, delimited upward by an arched film of unicells and oscillatoriaceans (dark streak) and downward by the surface of the solid crust. Plane polarized light. Scale bar = 100 ym.

C. Detail of resulting fabric: contact between a coarsely crystalline layer, made of clusters of calcitic blades such as those forming in B, and a microcrystalline layer associated with Schizothrix. Cross polarized light. Scale bar = 100ym.

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Two main types of cryptalgal laminites can be separated as a function ot their lithotope and of climatic conditions:

(1) On intertidal and supratidal flats of arid regions algae generally grow uncalcified leathery films or mats; they may be interspersed with detrital sands, .silts and eventually muds which may mold some of the filaments; when buried under sediment, these mats shrink and collapse to form organic- rich layers or streaks alternating with sediment-rich layers (Fig. 11).

Fig. 11. Cryptalgal laminite; laminated fabric results from alternating thicker, sediment- rich, layers and organic-rich dark laminae. Intertidal zone, Abu Dhabi, Persian Gulf. Courtesy of B. Purser.

(2) On supratidal flats from maritime tropical regions (e.g., Florida, Bahamas), submitted to seasonal flooding by freshwater during the rainy season, the algal mat filaments may calcify as they grow, and very distinctive cryptalgal laminites are formed (Monty and Hardie, Ch. 8.6); the resulting deposit shows among other things, a prominent lamination resulting from the


alternation of micritic layers of calcareous filaments, that formed during the rainy season, with layers of pellet packstone of wackestone deposited during hurricanes or severe storms (cf. Monty and Hardie, Ch. 8.6, Fig. 9).

However, lamination of cryptalgal laminites is not always so clearcut and regularly bound to a small number of well-defined cosmic events. In many cases, it appears to be somewhat erratic and results from irregular variations in the organic matter/sediment ratio of thk successive laminae. Indeed, intertidal algal mats do not grow at the same rate throughout the year; they may show periods of rapid growth and periods when it slows down depending on the growth rhythm of the component algae or the onset of particular local conditions (e.g., dryer periods, higher salinities). Similarly, the amount of detritus shifted on to the mats per unit time varies cyclicly (in response to normal tides, neap tides, spring tides and so on) or erratically (e.g., in response to accidental storms or stronger winds). In these conditions, subcontinuously growing algal mats will not show such a distinct lamination as illustrated on Fig. 11, but a succession of zones with different organic matter/ sediment ratios or with different detrital grain sizes. If cementation is rapid, a laminated fabric will result from the differentiation of laminae with different packing or size of their constitutive grains (Fig. 12).

Fig. 12. Laminated fabric resulting from cyclic changes in granularity and packing of bound grains. Lower intertidal zone, Shark Bay, Australia. Plane polarized light. Scale bar = 500 pm.

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There are, finally, cases where the production of laminae is completely erratic and episodic. An example is the mats built by Schizothrix, Scytonema, and Microcoleus on the intertidal flats of tidal swamps, eastern Andros (Monty, 1965b, 1967). Beside normal regular corpuscular laminae, these mats show many discontinuous, highly variable layers of detrital particles which result from very local and accidental factors. Adjacent mats, or even parts of mats, may be building hyaline algal layers in one place and binding detrital grains in another. Continuous observations revealed some of the responsible accidental factors: during high tides, the flats are covered by 10-20 cm of water harbouring puffers, voracious fishes common in mangrove channels and ponds. When disturbed by predator or prey, they rapidly swim out of their resting place, stirring up local clouds of sediments that will be redeposited on the surrounding mats to form a local detrital lamina. Another example is provided by burrowing annelids which build conical mounds on the flats, amid the algal structures. Muddy and silty waters expelled from active mounds flow down the cones and deposit their load over the adjacent mats.

Periodic inorganic cementation Periodic cementation due to change in (micro)-environmental, geochemical

or biochemical conditions may result in an alternation of lithified and non- or poorly lithified laminae and hence in the development, or the re-inforce- ment, of a laminated or laminoid fabric. This is true of well-lithified structures growing in the lower intertidal zone at Carbla Point (Shark Bay, Australia), where calcareous particles, mainly ooids and bioclasts, are bound by filaments of Schizothrh (Fig. 13A). Specimens studied show a laminoid fabric resulting from the alternation of cemented layers of packstone about 0.5mm thick, with uncemented grainstone layers up to 3 mm thick (Fig. 13B, C). Cemented layers show generally a much closer packing of grains than the non-cemented ones where microstructure is looser. Detrital grains are embedded into a micritic aragonitic matrix (Fig. 13D, E) rich in organic matter, at least in the

Fig. 13. Laminated fabric resulting from periodic cementation. A. General view of a poorly laminated dome built by Schizothrix sp. (so-called “smooth

mat”) lower intertidal zone, Carbla Point, Shark Bay, Australia. Impregnated sample. Scale bar = 1 cm.

B. Detail of lamination: contact between a non-calcified grainstone lamina (lower half) and a strongly cemented packstone lamina (upper half ). Plane polarized light. Scale bar = 200 pm.

C. General view of lamination. Beside differential cementation, a laminated fabric also results from differential packing which is usually much closer in the dark cemented laminae. Plane polarized light. Scale bar = 2 mm.

D. Detail of an initial stage of cementation of grains by fibrous aragonite. Note relationships between packing and the development of the fibrous aragonitic rim. Cross polarized light. Scale bar = 100 pm.

E. General view of cemented lamina with high porosity. Cross polarized light. Scale bar = 200 pm.

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upper layers, as shown by histological stains. Locally the cemented layers may dichotomize and surround small fenestrae. Grains are abundantly bored and micritized in contrast with those in the non-cemented layers. Cement- ation of grains comprises the deposition of fine rims of acicular aragonite around the ooids (Fig. 13D) and the filling of the interstices by microcrystalline aragonite (Fig. 13E). Meniscus cements are also found. Another occurrence of periodical cementation is found in the organosedimentary laminated crusts built by Scytonema along western Andros, Bahamas (Monty and Hardie, Ch. 8.6); here, the cement is a high-Mg calcite.

D iagenesis Variations in the chemical composition of (parts of) laminae, such as the

presence of higher-Mg calcites or higher concentrations in magnesium in given algal sheaths, may not be conspicuous in fresh specimens unless these zones coincide with lithified laminae, as is often the case along western Andros (Shinn et al., 1969; Gebelein, 1975). During diagenesis deeper reactions can cause either the differentiation of dolomitic laminae alternating with calcitic ones, as indicated by Gebelein and Hoffman (1969), or the systematic dolomitization of parts of laminae as illustrated on Fig. 14.

Alignment o f particles When the surface of an algal mat is smooth, i.e. when the constitutive fila-

ments are small and have a regular tangled mode, elongate free-falling particles being deposited on the sloping flanks of the structure will be submitted to gravity forces as well as to the “retention” forces of the sticky surface of the mat; as a result of the combination of these two forces, larger particles will orient themselves with their long axis parallel to the slope, i.e. to the growth surface of the mat. These particles will soon be bound in that position by overgrowing filaments, and incorporated into the structure. The continuation of this process will lead to the development of an overall lineation and hence of a laminoid boundstone fabric (Fig. 15).

A lignmen t o f fenestrae A fenestra has been defined by Tebbutt e t al. (1965, p.4) as a “primary or

penecontemporaneous gap in rock framework, larger than grain-supported interstices (. . .) The distinguishing characteristic of fenestrae is that the spaces have no apparent support in the framework of primary grains composing the sediment”. Fenestrae are quite frequent in cryptalgal structures

Fig. 14. Diagenetic development of laminated fabrics. Detail of lamination showing a microcrystalline algal layer, with filament molds, confined between two dolomitic laminae. Systematic gradient in the density of dolomitic rhombs (greater just above the algal layer, then diminishing progressively) probably reflects clines in the initial concentration in magnesium of the structure. Ordovician stromatolite. Arbuckle Mountains, U.S.A. Cross polarized light. Scale bar = 1 mm.

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Fig. 15. Outlining of laminoid fabric by commonly oriented elongate particles. Steep flank of a lower intertidal dome built by Schizothrix sp., Carbla Point, Shark Bay, Australia. Plane polarized light. Scale bar = 500pm. Ink line on left emphasizes lamination.

and, among other things, may result from: (1) Detachment of the surficial algal mat layers from underlying ones as a

result of: (a) Generation of gas bubbles during photosynthesis or bacterial decom-

position. The rising bubbles either distort the grain packing and/or accumulate under the surface algal film producing pustules. As predicted by Tebbutt et al. (1965) these bubbles generally extend laterally along physical discontinuities between laminae and originate an alignment of elongate cavities.

(b) Shrinkage and drying of the mat. When the mats (described on p. 199) are submitted to intense seasonal desiccation, not only do the calcified layers shrink considerably, but widespread detachment from underlying layers may occur along the organic partings (compacted former hyaline layers) ; aligned elongated cavities appear which generally have smooth roofs and floors.

Somewhat more complex situations combining separation of laminae aft9r desiccation and internal algal growth have been observed in laminated algal domes (10-15 cm across and up to 2 cm thick) from the intertidal flats of tidal marshes, eastern Andros (Monty, 1965a). When prolonged periods of low water expose the mats to partial desiccation, the surficial layers, rich in water and mucilage, shrivel above the underlying more compact ones (Fig. 16). Lenticular cavities with flat floors and convex roofs progressively develop along a depositional surface and range from 400 to 600 pm in height


Fig. 16. Formation of fenestrae by detachment of surficial algal layers and internal growth in laminated mats built by Scy tonema, tidal marshes, eastern Andros Island, Bahamas.

and 0.5 mm to several millimeters in length. When the mat is soaked again, these new habitats are rapidly colonized by slender filamentous oscillatoriaceans growing erect filaments from floor to roof. Vertical pressure resulting from the growth of these algae distorts the enclosing corpuscular layers and enlarges the organic cavities. As incipient micritic cementation occurs in the overlying corpuscular lamina, these internal algal phases will be preserved as aligned fenestrae separated by layers of particulate carbonate (the organic matter in the fenestrae will be oxidized or consumed). It should be noted that all intermediates exist between aligned discrete cavities and situations where internal growth extends subcontinuously throughout the whole mat originating a new lamina as seen on p. 204-205.

(c) Active growth resulting in rapid lateral expansion. If the available space is limited, the blooming mat will dome over the older parts of the structure and become wrinkled, circumscribing hollow cavities underneath.

Well-developed laminoid fenestral fabrics are found in subtidal heads and columns off Carbla Point (Shark Bay, Australia) (Fig. 17A). They will be described here in some detail as they provide an endpoint of the processes discussed above. The major fenestrae, which are about 5 mm wide and up to 30 mm long, are arranged in zones parallel to the surface of the mat, although many have irregular shapes. They are open, or partially or completely filled with sediment. They are surrounded by dense microcrystalline and cemented layers of wackstone (white on Fig. 17A) which emphasize the general laminoid fabric. The cement consists of fibrous aragonite which grows radially

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around the grains (Fig. 17D); the length of the fibres varies from 5 to 10 pm, although larger and smaller sizes are found. The development of the microcrystalline cement around the grains, which are deeply bored and micritized, obscures completely the original texture of the sediment (Fig. 17C). The cemented zones seem to develop from an algal mat (Fig. 17E), part of which domed when actively growing and formed internal cavities which remained as fenestrae; these mats actively bound detrital particles which are closely packed. Cementation must have been rapid as many flocs of organic matter as well as residual filaments have not been completely oxidised and can be extracted from the microcrystalline cement by gentle etching (Fig. 17F). It seems evident that cementation is bound to given algal mat phases, as no significant cement can be found elsewhere in the cryptalgal structure.

In between the units composed of the fenestrae and their cemented surrounding laminae, are layers or lenses of unconsolidated sand (Fig. 17B); these are microstratified and show definite boundstone fabrics. The micro- stratification results from an alternation of grainstone and packstone laminae, 500-1500pm thick, with algal films, 100-250 pm thick; these are not always well preserved and may be recorded as thin micritic films, as clotted or as loose and unsupported laminae.

(2) When algal layers, isolated bundles of filaments, or colonies of unicells are not calcified, their organic matter may be rapidly oxidized, mainly in coarse sediments, leaving elongate, irregular, horizontal or vertical cavities (Fig. 19; Monty and Hardie, Ch. 8.6, Fig. 14).

A particular type of fenestra with a flat floor and an irregular, saw-toothed roof, quite frequent in the geological record, may be predicted from laminated mats observed in tidal marshes of eastern Andros Island and elsewhere. In cross-section, some of these mats show an alternation between erect growths of Scytonerna and layers of detrital carbonate (Fig. 18A). In many

Fig. 17. Development of laminoid fenestral fabrics in a dome built by oscillatoriaceans (Schizothrix, Oscillatoria) in the subtidal zone, Carbla Point, Shark Bay, Australia. A. Laminoid fabric results from alignment of elongate fenestrae and is emphasized by

calcified layers (white) which enclose them. Impregnated slab, scale bar = 1 cm. B. Contact between a non-cemented grainstone lamina (appearing medium gray on A)

and the top of a cemented fenestral packstone lamina. Fenestrae here are open, which is not always the case. Plane polarized light. Scale bar = 2 mm.

c. Detail of a cemented zone roofing a fenestra; note intense micritization and boring of grains (ooids) which fade into microcrystalline aragonitic cement. Cross polarized light. Scale bar = 200 pm.

D. Production of microcrystalline fibrous aragonite around micritized grains to yield undifferentiated micritic patches such as visible on middle left of C. Cross polarized light. Scale bar = 100 pm.

E. Vestige of part of a domed algal mat (arrow), in cemented lamina, under which a fenestra formed. Plane polarized light. Scale bar = 100pm.

F. Remains of coarse algal filaments extracted from cemented laminae by gentle etching. Plane polarized light. Scale bar = 100 pm.

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A

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9

6

1 mm - Fig. 18. Formation of fenestrae with smooth floor and saw-toothed roof after oxidation of non-calcified growths of Scytonema growing on a depositional surface. Prediction from observation of laminated mats built by Scytonema ; tidal marshes, eastern Andros Island, Bahamas.

cases, Scytonerna forms discontinuous layers with more actively growing tufts here and poorly growing filaments there, although belonging to the same horizon. The uncalcified algal mat grows on the subhorizontal depositional surface of an earlier detrital layer. The top of the algal growth is quite irregular due to the variously projecting bundles of filaments. This surface is accordingly replicated by the base of the overlying corpuscular layer which has a characteristic scalloped and saw-toothed outline. Oxidation of the algal mat and cementation of the corpuscular layers would leave typical fenestrae aslshown on Fig. 18B.

(3) Algal or bacterial life within the mat may originate significant changes in COz concentration and accordingly in pH; this may result in solution of carbonate within the cryptalgal structure and the opening of cavities in given horizons (Fig. 20). (4) Animal organisms living within or on the mat may, at their death,

leave aligned round or elongate cavities (Fig. 7). This is illustrated by the tubes of chironomids in freshwater algal structures: “they form long cavities up to 5 mm long and 0.5 mm wide (. . .) They are always parallel to the annual layers” (Irion and Muller, 1968, pp. 165-166, their Figs. 4, 5).

For the fenestrae to be preserved (mainly in processes 1 and 2 described in this section) $he roofing layer needs some degree of coherence - as may be provided by desiccation - and rapid cementation (or cementation of the surrounding sediment).


Fig. 19. Formation of an open vaulted fenestra after oxidation of a flabellate non-calcified bundle of Scytonema over which a smooth Schizothrix mat draped. Lower intertidal zone. Hamelin Pool, Shark Bay, Australia. Plane polarized light. Scale bar = 500pm.

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Fig. 20. Opening of an elongate fenestra by internal solution of a microcrystalline layer of Scytonema beneath a smooth Schizothrix mat. Note characteristic solution surface cor- roding the Scytonema flabellate growth. Lower intertidal zone. Hamelin Pool, Shark Bay, Australia. Plane polarized light. Scale bar = 500pm.

Complex or composite lamination Combinations of the previously described processes are frequently met.

When they act “in phase”, the resulting fabric will not be more complex than the one yielded by each of them separately; the lamination will eventually be re-inforced. This is, for instance, the case for the Phormidium (Schizothrix) laminated domes described on p. 207; here, the sedimentary event outlines the biolamination resulting from the phototactic behavior of the alga, as detrital particles concentrate in the lamina built by resting filaments.

When the combined processes act “out of phase”, a composite lamination is formed. Such a situation can be found in the crusts built by Phormidium incrustatum; the basic fabric made of alternating laminae of strongly and poorly lithified filaments (Fig. 8) can be complicated by the superposition of various processes originating the development of sublaminae (Fig. 21). Among these processes the following can be found (together): (1) periodic blooms of calcified films of Schizothrix across the radial framework of Phormidium (Fig. 8A, 21); (2) the seasonal influx of detrital particles which contaminates the loosely cemented layer and concentrates on top of it (Fig. 21); and (3) the development of aligned elongated fenestrae due to chironomids or to the oxidation of seasonal clumps of diatoms (Figs. 7, 21).

Another type of complex’lamination can be found in the crusts and


cryptalgal laminites of western Andros (Monty and Hardie, p. 459-471), where biological, sedimentological and geochemical processes act at different moments, each contributing a particular (sub)lamina; the resulting cyclothemic lamination results from the repetition of the following series of events: (1) growth of a calcified Scytonerna mat; (2) influx of a detrital pelletal layer; (3) cementation of the upper half of this layer which becomes a pellet packstone overlying a pellet grainstone; and (4) precipitation of a micritic film.

Composite lamination may also be built by ecotonal communities; an ecotone (Odum, 1959) can be briefly defined as a transition between two or more communities. The ecotonal community contains not only many of the organisms of each of the overlapping communities but also organisms proper to the ecotone. This concept, which has already been theoretically applied to the development of stromatolitic microstructures (Monty, 1973c, fig. 4) can be clearly seen when studying the evolution of algal-mat fabrics along a cline (e.g., tidal flat or a coast). In the ecotone where two blue-green algal communities overlap, algal mats are built in which each community imprints its own microstructure, the combination of which originates new fabrics. Such a situation can be found in stromatolites of the lower intertidal zone of Shark Bay where Schizothrix communities, originating the so-called “smooth mat” (Logan et al., 1974) illustrated on Fig. 13, overlap with the Scytonema community. Columnar structures are formed (Figs. 22A, 23), in which a laminoid fabric is imprinted by the recurrence of mats of Schizothrix, which bind detrital particles, whereas the doming of the lamination results from the radial flabellate growths of Scytonema over which Schizothrix mats are draped (Fig. 22B). As a result, a composite complex fabric is formed: a gross lamination may evolve from the succession of layers built mainly by Scyto- nema and layers built by sets of films of Schizothrix (Fig. 22B, C). This does not necessarily correspond to seasonal successions but reflects the competition for space between the two dominant algae, one outcompeting the other for some time. Each layer will exhibit the fabric imposed by the component algae: the zones built by Schizothrix will show a laminoid boundstone fabric where elongate grains tend to orient themselves parallel to the mat surface. In zones built by Scytonerna, grains are not primarily bound but trapped: they fall in between the erect filaments and lie without any preferential 6rientation. Furthermore, although most of the Scy tonema are not calcified here, this alga will tend to imprint its typical radial filament fabric on its layers; this will be done in two ways: (1) development of tubular vertical microfenestrae separated by packstone, after oxidation of the filaments (Fig. 25); (2) development of moulds of filaments where microcrystalline aragonite has been precipitated in the pervading mucilages (Fig. 24). Finally, at a smaller scale, the superposition of the two fabrics may originate a reticulate microstructure where the radially growing filaments of Scy tonerna are regularly cross-cut by horizontal films of Schizothrix, as already seen on p. 202 (Fig. 25). I t should be added that in such ecotonal situations, the

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overlapping communities may not only overgrow each other as time goes by, but also juxtapose themselves to build a complex stromatolite with two different interfingering fabrics.

NON-LAYERED FABRICS

Non-laminated cryptalgal fabrics are quite varied and numerous; only three major types will be considered here: the thrombolitic, massive and radial fabrics.

Throm bolitic fabrics

Aitken (1967) proposed the term throm bolite for “cryptalgal structures related to stromatolites, but lacking lamination and characterized by a macroscopic clotted fabric” (p. 1164). “The microfabric of thrombolite consists of centimeter-sized patches or clots of microcrystalline limestone (grain size 8-20 microns) with rare clastic particles, in part indistinctly ‘laminated, separated by spaces filled either with sparry calcite or silt- and sand-sized . . . sediments” (p. 1171). He adds “in some instances, microscopic filaments including rare examples assignable to Giruanella are present in the clots and in some specimens clearly make up the bulk of the clot forming material” (p. 1172).

Present-day occurrences show that thrombolitic fabrics may originate in several ways and show a wide range of complexity.

Oxidation of mamillate and pustular algal mats Irregular fenestral fabrics have been found in association with pustular

colonies of unicellular blue-green algae, in Shark Bay, Australia (Logan et al., 1974). Colonies of Entophysalis up to 3 cm wide and 2 cm thick (Fig. 26D) form a subcontinuous mat showing a strongly mamillated surface. Sediment particles are trapped and entombed in the depressions between the adjacent colonies, as well as on their surfaces, where they are soon cemented by microcrystalline aragonite. The oxidation of dead algal colonies leaves abundant irregular voids separated by patches of pellet or intraclast packstone, originating a thrombolitic fabric (Fig. 26A, B, C). This simple scheme is, however, complicated by a number of processes.

~~ ~~

Fig. 21. Development of composite laminatton in a Phormidiurn incrustaturn head such as illustrated on Fig. 7. Here, the normal alternation of densely calcified with poorly calcified seasonal Phorrnidiurn layers (Fig. 8D) is complicated by the intercalation of (1) a lamina (B) of aligned fenestrae due to chironomids (see F on Fig. 7) overgrowing the heavily calcified spring layer (A); (2) a lamina (C) made of detrital terrigenous grains, and the superposition of Schizothrix films (E) over the next Phorrnidium growth (from D up). Plane polarized light. Scale bar = 500pm.

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(a) Microcrystalline carbonate may be precipitated within the pustular colonies or around them (Fig. 26D). In the first case, the oxidation of the algae will not leave an empty cavity but flocs of micrite. In the second, the peripheral micrite may form a coherent lining that would support the cavity when the algal material disappears.

(b) Whereas some corpuscular patches appear to be constituted almost exclusively of detrital particles (ooids, bioclasts, etc.) others seem to form in situ, and several stages can be found: many ovoid to subspherical peloids (50-150 pm across) appear to derive from hollow shells of organic matter on the inner surface of which aragonitic microspherulites are first precipitated (Fig. 27A); they are followed by successive populations that fill up the shell from inside. Some of the peloids may show several stages of growth (Fig. 27C): they are composed of an inner compact core, made of imbricated microspherulites, surrounded by an accretionary zone delimited outward by an organic film “floating” around the core while new microspherulites are precipitated in the intervening spaces, As a result of the imbrication of the constitutive microspherulites, the in situ particles often appear as micritic peloids quite similar to the incorporated but otherwise completely micritized, detrital grains (in many cases, however, gentle etching can liberate the spherulites and help distinguish the grains). As these peloids grow amid members of the algal community, corpuscular patches or clots (sensu Aitkens) are progressively formed in which autochthonous micritic grains are enclosed in a pervading, sometimes filamentous, organic matter, where they seem to have been trapped and bound. In older parts of the structure, cryptocrystalline aragonite is progressively deposited in the interstitial organic matter, and clots of pellet packstone are constructed in situ (Fig. 27D).

In the other cases, the microspherulitic precipitation is not limited to individual peloids but involves the greater part of a thrombolitic clot; in this case, a big patch of homogeneous cryptocrystalline (microspherulitic) aragonite is formed and accretes centrifugally by precipitation of new microspherulites between the clot and the surrounding organic film (Fig. 27E, F). Such in situ formation of micrite particles and clots has already been described by Monty (1965a, 1967), Dalrymple (1965), Friedman et al. (1973),

~~

Fig. 22. Stromatolites built by ecotonal communities. A. Vertical section in columnar stromatolite resulting from the overlap of the lower inter-

tidal Schizothrix mat and Scytonema mat. Hamelin Pool, Shark Bay, Australia. Impreg- nated slab. Scale bar = 1 cm.

B. Detail of Fig. 23 showing the radial growth of two Scytonema colonies (one in the lowermost part of the photo, the other in the upper half) which originated the doming of lamination (see the dark thin Schizothrix film (s f . ) doming over the upper Scyto- nema growth) and the formation of column-like structures. Plane polarized light. Scale bar = 2 mm.

C. General view of lamination: thin dense laminae built by Schizothrix binding detrital grains alternate with looser layers resulting from the entrapment of grains by tufts of Scytonema, most of which have been oxidized. Plane polarized light. Scale bar = 2 mm.

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Krumbein (1974), and Krumbein and Cohen (1974); aragonitic spherulites have been experimentally formed by bacterial action in marine waters by Oppenheimer (1965) and Krumbein (1974). It is evident that the formation of such thrombolites is not a simple matter, but results from a complex of sedimentary, organic and biochemical processes. Microspherulite precipitation as a factor of thrombolite genesis will be further substantiated below (p. 232-235).

Coalescence of calcified botryoidal colonies

(a) Simple situations. Aitken (1967), in his study of Early Paleozoic cryptalgal carbonates, reported situations where filaments of Giruanella could be found in thrombolitic clots or even could form the bulk of them. Analogous situations are found in Recent thrombolitic carbonates where the constitutive clots represent calcified algal colonies endowed with a botryoidal surface growth form; as the various algal knobs merge numerous cavities are created, originating an irregular fenestral fabric. Two examples will be given here: freshwater cryptalgal tufa deposited in Green Lake, New York, U.S.A. is “exceedingly porous or spongy and consists of more or less closely intergrown arborescent masses that are richly nodose” (Bradley, 1929a, p. 205). The botryoidal upper surface of the deposit is obvious on Walcott’s (1916) pl. 4, fig. 3 and Bradley’s (1929a) pl. 29B, whereas the resultant thrombolitic fabric is illustrated in Walcott’s pl. 4, fig. 4. The microfabric of the clots is radial or laminated: radial flabellate growths of blue-green algal filaments (oscillatoriaceans), forming the calcareous knobs, are illustrated by Bradley (1929, pl. 30, figs. A, B); fine-grained calcite is first deposited in the interstices between the filaments and cementation follows.* When the algal felts decay and disappear from the clot, irregular radially elongated cavities are left that will also be cemented.

* The carbonate is algally precipitated according to Bradley and Walcott but detrital and trapped according to Dean and Eggleston (1975). although their paper is inconclusive as to how much of the fine-grained carbonate material is really detrital and how much is precipitated in situ .

Fig. 23. General view of section cut in sample shown on Fig. 22A. Frame delimits part of Fig. 22B. Two columns are clearly visible: half of one covers the right-hand side of the photograph, part of the other appears along the left margin. They are separated by a gap infilled with detrital particles sometimes lying edgewise; in the upper half, Schizothrix films extend from columns over the depositional surface of the infilling material (arrows). Fabric results from alternation and/or interfingering of radial growth of Scytonema (often leaving vertical tubular microfenestrae (see Fig. 25) with smooth films of Schizo- thrk doming over them). The right-hand column indicates the role of Scytonema in the shaping and framework of the structure. Plane polarized light. Scale bar = 1 cm. Column margins have been partly emphasized with ink lines.

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Fig. 24. Detail of Fig. 23. Radial filament fabric of Scytonema preserved as moulds by microcrystalline aragonite. Plane polarized light. Scale bar = 500 pm.

A laminated pattern may also appear in the clots of Green Lake thrombolites, mainly near the surface (Dean and Eggleston, 1975, fig. 14, general view; Eggleston and Dean, Ch. 8.7, Fig. 7, close up). Such laminated knobs would result from the alternation of phases of algal-mat growth with phases of physicochemical precipitation, as proposed by Dean and Eggleston (Ch. 8.7).

In other thrombolites, the clots are definitely concentrically laminated as a result of the episodic growth of the constitutive bluegreen algal colonies (Fig. 28). Such cases occur among the variety of cryptalgal carbonates that are deposited in the Hoyoux river, Belgium, where algae like Schizothrk and Phormidium are growing laminated botryoidal knobby encrustations. Here, the irregular fenestral fabrics result from two main factors: first, the merging of the various algal knobs which circumscribe internal cavities; second, the presence of globose masses of diatoms encrusted by carbonate-depositing blue-green algae. When the diatoms disappear, open irregular fenestrae are left in the tufa.

(b) Complex situations. Another type of thrombolite fabric occurs in littoral deposits on the northern side of Mono Lake, California. The origin of these various deposits has been much disputed (Dunn, 1953;.Scholl and Taft, 1964). One case is here reported as an illustration of a variety of cryptalgal fabric


Fig. 25. Detail of Fig. 23. Radial filament fabric of Scytonema preserved as vertical tubular microfenestrae after oxidation of the algal bundles. Overlap of ecotonal communities results in the superposition of horizontal films of Schizothrix on the radial fabric, leading to a reticulate microfabric. Plane polarized light. Scale bar = 500pm.

which also forms by coalescence of small calcareous knobs. It might also help understanding spherulitic microfabrics such as those described by Buchbinder et al. (1974). The samples, collected in 1964, are rather porous and show in section, a series of rounded (from less than a mm to 1 cm across) to irregular or flattened crenulated carbonate flocs separated by open cavities (Fig. 29A). In thin section, these flocs have a “cellular” microstructure (Fig. 29B, C); each “cell” is composed of one or several imbricated spherulites (Fig. 29D), about 40 pm across, and is generally surrounded by darkish discontinuous organic matter. Observation near the surface of a mamillated knob shows that spherulites are actually forming within or beneath a brownish microfilamen- tous, apparently bacterial, film doming over the structure (Fig. 29F). The genesis can hence be viewed as follows: mono- or polyspherulitic microclots grow under the influence of bacterial colonies until they meet similar growing microclots in which case crowding prevents any further growth and a com- promise “cellular” microstructure progressively appears. Macroscopic clots are progressively formed by juxta- and superposition of “cells” or spherules. These clots have a very knobby surface as the juxtaposition of spherulitic units tends to produce rounded protuberances. The projection of these knobby clots (such as illustrated on Fig. 29B) merge in turn with other similar bodies and the resultant internal cavities originate a conspicuous thrombolitic

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fabric. At this stage, the cavities are soon invaded by mucilage rich in bacteria and microfilaments which either fill the whole cavity or line it with a compact film. In between this continuous fi!m and the cavity wall, carbonate precipitation resumes, no longer as discrete spherulites, but as a continuous layer of closely packed needles up to 50pm long and very reminiscent of inorganic cements (Fig. 29C, E). In this way, the “cellular” microspherulitic clots become enclosed in a subcontinuous fibrous girdle. Further complications are found in these thrombolites due to overgrowth by various algal and bacterial colonies, but they will not be considered here as this is a problem in itself.

Corrosion of deposited cryptalgal carbonate The development of irregular fenestrae leading to thrombolite-like fabrics

may also result from the chemical processes of internal solution. This is illustrated by the large calcareous nodules (5-25 cm across) that are built by blue- green algae on the bottom of the Rhine River, at Stein am Rhein, Germany (Baumann, 1913; personal observations), and which in section show a very porous fabric. These pores and holes that cut through the calcified algal mass are interpreted by Golubic (1973a) as developing “by secondary carbonate dissolution of the interior of the crust caused by local C 0 2 accumulation from bacterial action” (p. 442, see his fig. 21.4).

It should be added here that the cavities and fenestrae of all these porous cryptalgal deposits generally harbour multitudes of organisms such as insect larvae, small crustaceans and nematodes; these organisms contribute to maintaining these cavities open and may eventually enlarge them.

The massive fabrics

Trapping mechanisms Massive cryptalgal fabrics show no internal lineation or spatial organization

of their constitutive elements. Thus, they will not be easily identifiable as cryptalgal. Among the processes liable to originate massive fabrics, five will be briefly described.

Fig. 26. Thrombolitic fabrics. A. General view of a thrombolitic cryptalgal structure associated with pustular colonies of

Entophysalis major Ercegovih. Lower intertidal zone, Hamelin Pool, Shark Bay, Aus- tralia. Scale bar = 1 cm.

B. Vertical section showing the development of irregular fenestrae delimited by cemented corpuscular clots. Plane polarized light. Scale bar = 2 mm.

C. Detail of a clot made of detrital and in situ particles, commonly held together by organic matter, and larger patches of microcrystalline aragonite. Plane polarized light. Scale bar = 500 pm.

D. Pustular colony of Entophysalis major at the surface of the living mat; dark clouds and linings inside and around the colony are made of microcrystalline aragonite. Plane polarized light. Scale bar = 200 pm.

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Entrapment of detrital grains is common in communities of strong blue- green algae (e.g., Dichothrix bornetiana (Howe), species of Rivularia, Scyto- nema, and Lyngbya) which grow erect tufts of filaments between which silt- and sand-sized particles infiltrate and remain trapped (Fig. 30). Besides blue-green algae, filamentous green and red algae may also form mats actively trapping detrital grains; good instances are illustrated by the Enteromorpha (green) mat described from the Bahamas by Scoffin (1970, p. 265, fig. 22) and the Rhodothamniella (red) mat described by Hommeril and Rioult (1965, pp. 137-140, fig. 3A) from the coast of Normandy, France. Such algal filaments project freely from 0.5 to 2 cm above the sediment surface. In areas where shifting of sand is continuous, grains constantly infiltrate in between the mesh of filaments, where they accumulate. In the subtidal zone of the Bahamas and the intertidal zone of the windward lagoon of Andros Island (Monty, 196513, 1967) the deposit progressively collapses as the filaments disappear and would yield an undifferentiated massive grainstone. It should be noted, however, that in cases of rapid cementation of the sand, as happens in Shark Bay, Australia, the fabric would evolve differently with the preservation of irregular piles of packstone separated by subvertical elongate cavities left by the oxidized bundles of filaments; here, a “tubular fenestral fabric” rather than a massive fabric would be formed.

As is often the case in boundstones, the grains trapped by the algae are better sorted than the surrounding sediments and the average grain size is lower, as only the finer fractions are resuspended and transported to the mats. However, as opposed to what happens when sedimentary grains are directly bound, the elongated particles will not show any common orientation parallel to the growing surface but may often stay edgewise (which may be a typical feature of “trapstones”).

Fig. 27. Development of in situ grains and clots in Shark Bay thrombolites. A. In situ grain in process of formation: fibrous and microspherulitic aragonite is being

precipitated on the inner surface of an organic shell and accretes inward to finally yield a micritic grain. Plane polarized light. Scale bar = 50pm.

B. Collection of in situ grains in corpuscular patches such as appears on Fig. 26C. Note their microspherulitic microstructure. Plane polarized light. Scale bar = 75 pm.

C. In situ microspherulitic grain with accretionary zone made of a lobate organic film (black lining at the surface) underneath which new microspherulites are being precipitated. Plane polarized light. Scale bar = 50pm.

D. Detail of a clot made of in situ microspherulitic grains bound by pervading organic matter in which microcrystalline carbonate is being precipitated. Plane polarized light. Scale bar = 200pm.

E. Clot made of microspherulitic aragonite and surrounded by an organic dark film. The clot is accreting mainly on its higher upper left margin (see detail in F). Black spots are air bubbles. Plane polarized light. Scale bar = 500pm.

F. Detail of the upper left margin of clot shown on E. Accretion results from precipitation of new microspherulites in the interspaces between the organic film (arrow) and the older parts of the clot over which it domes. Plane polarized light. Scale bar = 50 pm.

238 C.L.V. MONTY

Fig. 28. Development of thrombolitic fabrics (continued). A. Fragment of knobby cryptalgal deposit built by clustered lobate colonies of oscilla-

toriaceans (Hoyoux river, Belgium). The merging of all these knobs and projections produces irregular internal cavities delimited by algal clots and hence a particular thrombolitic fabric. Scale bar = 5 mm.

B. Section of structure illustrated in A showing algal clots with radially growing filaments (not visible on photograph); layering of the clots results from seasonal growth. Plane polarized light. Scale bar = 500 pm.


Along with simple trapping mechanisms, other processes generally occur simultaneously. For instance, the mud fraction of the sediment, when present, does not tumble down in between the filaments, but sticks to their mucilaginous sheaths. In this case, the algal filaments may become encased in a micritic sheath and the original algal fabric will be variously replicated according to the abundance of mud and the fidelity of the mould. Also, the leading alga, which makes the sedimentary trap, forms a general framework in which other algae thrive (e.g., Schizothrix, Phormidium, Microcoleus). These will in turn overgrow and bind the detrital particles that have been trapped, and consolidate the deposit. They will not alter the “trapstone” fabric though, except that eventual precipitation of cryptocrystalline carbonate in their mucilage may alter the grainstone to a packstone.

It should be clear that trapping mechanisms do not always yield non- laminated cryptalgal deposits; if the influx of detrital particles is episodic or if the textural features of the sediment carried onto the mat vary seasonally or accidentally, a lamination will develop in the mat.

Oversedimen tation Non-laminated, massive, fabrics may also form in rhythmically growing

algal mats, like the Phormidium colonies described on page 197, when the mats are swamped by a continuous heavy influx of detrital grains, or when vertical algal movement is reduced. For instance, when traced onto exposed tidal flats, where sediment influx is very important during tidal movement and algal growth considerably slows down during low tide, or when traced toward those lagoonal floors where abundant silt is deposited, these well- laminated domes may pass to unlaminated massive formations (Fig. 31). In thin section, such bound sediments appear as an homogeneous sandy or silty deposit, composed of well-sorted particles significantly finer than the surrounding sediments, whereas tiny filaments of Phormidium run erratically from one particle to another. Beside what has been said before (p. 237), the normally good sorting of the component particles results (1) from the absence of protruding coarse filaments that could simultaneously agglutinate the fine and trap the coarser material, and (2) from the micron size of the algal filaments which is such that only the fine fraction of the bottom load can be stabilized. Furthermore, the coarse particles, which have less chance of being resuspended by normal agitation, are too voluminous or too heavy to be firmly bound during a single tidal event.

If elongated particles are present in the sediment their orientation will not be gravity determined (see p. 217), which may give us clues as to the original boundstone nature of the deposit.

In brackish or hypersaline environments micritic carbonate may be progressively precipitated in the interstitial organic matter by bacterial action and a massive packstone is formed.

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FABRIC AND MICROSTRUCTURE 24 1

Binding of su bstratal sediment Unlaminated massive corpuscular deposits can also form in another way:

sediment is not carried onto a growing algal mat, but algae invade and bind a previously deposited sediment. Putting aside the submarine algal films which periodically stretch over the sea floor and stabilize the sand (Bathurst, 1968; Neumann et al., 1970; Scoffin, 1970), I shall mention here only the case of sandy “crusts” found in the proces’s of formation on some beaches bordering the lower Fresh Creek (Andros Island, Bahamas; Monty, 1965b). These friable, white to grayish, sandy agglomerates, about 1 cm thick, are composed of coarse organoclastic material and rest free on the loose skeletal beach sand. Their surface is rough, granular and devoid of any significant algal vegetation (Fig. 32). A thin section shows a poorly sorted biocalcarenite loosely bound by filaments of Schizothrix and glued by mucilaginous masses of Entophys- alis, sometimes loaded with clouds of fine-grained carbonate. These structures originate by the local invasion of the loose beach sand by colonies of Schizo- thrix which bind the grains into friable flakes; once initiated, the sandy flakes thicken rapidly. Detrital particles, carried by the wind or the tides, are trapped and entombed in the irregularities of their rough surface where they will be bound. They produce a new and advantageous habitat for algal and bacterial colonies by providing a stable microenvironment (with no shifting of particles) in which moisture can be retained and protection against solar radiations is ensured. Therefore, most of the living material is inside the structure, amid mucilage. The organic material harbours bacteria that contribute to the precipitation of lime. Although friable, these structures harden considerably and curl up when dried, and can be incorporated into the beach deposit as chips of massive packstones.

Fig. 29. Development of thrombolitic fabrics (continued) A. General view of a littoral tufa from Mono Lake, California. Note variations in size and

shape of component clots (grayish to white) which go from subspherical (upper left) to elongate. Impregnated slab. Scale bar = 1 cm.

B. View of a small clot made of a lobate, spherulitic inner core coated with a fibrous irregular lining. Cross polarized light. Scale bar = 500 pm.

C. Detail of a clot. The groundmass is made of juxtaposed microspherulitic units (see D) originating its “cellular” microstructure. The rim of the clot is coated by a fibrous deposit (light band) growing under a probably bacterial organic film (0.f . ) Cross polarized light. Scale bar = 200pm.

D. Detail of the microspherulitic units composing the clots. Cross polarized light. Scale bar = 50 pm.

E. View of the bacterial organic film (arrow) under which is growing the fibrous deposit visible in B and C. The film split longitudinally during dehydration and staining of the preparation. Plane polarized light. Scale bar = 50 pm.

F. Initiation of spherulitic units: close up of scattered spherulites growing inside a probably bacterial envelope appearing blurred on top of photograph. Plane polarized light. Scale bar = 50pm.

242 C.L.V. MONTY

Fig. 30. Entrapment of detrital particles between bundles of filaments of Rivularia. Inter- tidal zone, eastern Andros Island, Bahamas. Plane polarized light. Scale bar = 500 /Jm.

Massive cryptalgal diatom mats Massive cryptalgal structures can also result from the continuous precipi-

tation of carbonate in growing mucilaginous algal masses. The intertidal mud flats bordering Fresh Creek, Andros Island, are locally covered with subpoly- gonal, grayish-black mats, about l c m thick and 3-5cm across (Monty, 1965a). The uppermost part of the mats is composed of a hyaline uncalcified film of diatoms, up to 100 pm thick (Fig. 33A); a faint layering may appear in this film as a result of the horizontal elongation of big clusters of diatoms. Beneath this is a layer up to 5mm thick composed of a mucilaginous mass rich in diatoms and bacteria, with scattered colonies of Entophysalis; within this is abundant cryptocrystalline carbonate characterized by crystals mostly smaller than 1 pm wide (Fig. 33B). This highly organic layer rests on an undifferentiated stratum of micrite where frustules of diatoms are still present. The superposition of these three layers illustrates the development of the mat: when conditions are favorable, diatoms bloom at the surface of the polygons and secrete a very mucilaginous stratum which is soon invaded by bacteria; calcium carbonate is progressively deposited at the expense of the mucilage and the living mat is progressively converted in a massive biomicrite.

FABRIC AND MICROSTRUCTURE 24 3

Fig. 31. Development of massive fabric due to oversedimentation on Phormidium films. Intertidal zone, cays off Fresh Creek, Andros Island, Bahamas. Plane polarized light. Scale bar = 5 mm.

, Fig. 32. Development of massive fabric due to invasion and binding of beach sand by Schizothrix while mucilaginous colonies of unicells, developing internally, contribute to gluing the particles together. View of resulting friable crust. Sandy beaches of the Fresh Creek, eastern Andros Island, Bahamas. Scale bar = 1 mm.

244 C.L.V. MONTY

Fig. 33. Development of massive fabrics in diatom mats. Intertidal beaches of the Fresh Creek, eastern Andros Island, Bahamas. A. Hyaline upper part of the mat made of piles of mucilaginous films of diatoms. Plane

polarized light. Scale bar = 200pm. B. Lower part of the seasonal mat where microcrystalline calcite is being precipitated in

the mucilage. Plane polarized light. Scale bar = 200 pm.

Similar phenomena have been reported from freshwater deposits by various authors (e.g., Wallner, 1935; Van Oye, 1938; Kann, 1940).

Diagenesis Compaction, oxidation, precipitation of micrite, micritization of grains,

dolomitization and recrystallization are among the many diagenetic processes that can blur or obliterate originally laminated or thrombolitic fabrics and produce massive crystalline fabrics. However, it is my experience that “ghost lamination”, visible only on weathered outcrop or in thin section, may persist even in crystalline fabrics.

Radial fabrics

Some algae, like Scytonema in the Recent or Orthonella and Garwoodia in the Paleozoic, frequently show a radial growth of skeletal (calcified) filaments. Such radial growths have already been described (e.g., p. 198) as part of laminated fabrics in which distinct laminae are composed of radially growing filaments. However, continuous or subcontinuous growth of such frame-building algae may lead to the formation of unlaminated to poorly


Fig. 34. Radial filament fabric in lithified calcitic dome built by Scytonema. Higher parts of freshwater marshes, western Andros Island, Bahamas. Plane polarized light. Scale bar = 1 cm.

laminated structures characterized by a typical radial filament fabric (Figs. 34, 35;Monty and Hardie, Ch. 8.6, fig. 11). This fabric may be interrupted here and there by concentric layers due to the development of horizontal growth stages of Scytonema (p. 198). Furthermore, as already described, factors such as concurrent development of horizontal films of oscillatoriaceans in ecotones, episodic influx of detrital particles or physicochemical precipitation of cryptocrystalline laminae, can alter the radial filament fabric into a reticulate fabric. The fossilization of such algal structures can lead to the development of sparry calcite in between the skeletal filaments or to the deposition of cryptocrystalline carbonate (Fig. 35).

THE IMPACT OF THE BIOLOGICAL STRATIFICATION OF THE ALGAL MAT ON THE RESULTING STRUCTURES

The layered fabric of stromatolites as it appears to us is in fact the total sum of two types of superposed stratifications, as has been developed in Monty (1965a): (1) a stratification in space or “instantaneous biological stratification” (Monty, 1973a) results from the taxonomical and/or metabolical differentiation of blue-green algal and bacterial communities which at any given time are superposed in a living algal mat; (2) a stratification in

246 C.L.V. MONTY

Fig. 35. Radial filament fabric in living Scytonerna crust (above) and subfossil one (below) where micrite has been precipitated between the calcified sheaths as well as in the lumen. Higher parts of the freshwater marsh, western Andros Island, Bahamas. Scale bar = 1 mm.

time or “historical biomineralogical stratification” results from the periodical superposition of layers.

The instantaneous biological stratification reflects the complex structure of algal mats. A good example of such stratification studied by Sorensen and


Conover (1962) is the Lyngbya mats of the Laguna Madre (Texas). These mats are 1-20mm thick and show a laminated fabric resulting from the superposition of layers characterized by different colours and different taxonomic composition. At the surface are deeply stained blue-green algae which screen the incident light, in the inner layers the protected algae are actively growing, and in the basal parts there are purple bacteria and colorless algae. Heterotrophic processes predominate in the basal zone (see also Golubic, Ch. 4.1). Even when they are built by only one dominant alga, mats may show a vertical stratification of populations of conspecific individuals characterized by different metabolisms and even different morphological and histological properties. Fritsch (1953) has for instance described the various morphologies and habits of conspecific filaments of Phormidium at different levels in a film of Phorrnidium. This stratification in space determines the individualization of superposed and clearly circumscribed metabolical and hence biochemical microenvironments. In the case of polyspecific mats it also reflects the individualization of efficient cenotic pools from which the mat will be able to differentiate seasonally and keep on growing in spite of changing climatic or atmospheric conditions (see p. 199ff). The relationships between this instantaneous biological stratification and the historical stratification may vary considerably according to the growth pattern of the stromatolite and in fine to the particularities of its basic module: the algal film or mat. Two extremes can be illustrated (Monty, 1973c, fig. 2): in the case of stromatolites resulting from the yearly superposition of rather thick mats the instantaneous stratification will generally be limited to the growing mat, the underlying ones being already in the process of diagenesis. At this level, it may originate the differentiation of the yearly layer into secondary laminae. When the stromatolite results from the rapid superposition (daily, monthly) of thin algal films, the biological stratification will extend over a considerable number of historical laminae in which it will probably induce several types of auto bioturbation.

Indeed, as the stratification in space is an energetic stratification, it may initiate fundamental reworking of the original fabrics and microstructures, for example:

(1) Encysted populations within the mat may suddenly start growing, their optimal ecological threshold being suddenly reached, and thicken the mat from the inside, disturbing the original fabrics and structures (Fig. 36).

(2) Filamentous populations moving upwards, as the stromatolite grows, to remain in the optimal energetic level, may originate the collapse of the laminae they were previously colonizing.

(3) Formation of internal cavities, disorganizing the microstructure, as a result of the consumption of mucilages and organic accumulations by heterotrophic organisms.

(4) The presence, at particular levels, of given metabolical organic acids may control the preferential precipitation of calcite, aragonite, or various

248 C.L.V. MONTY

Fig. 36. Autobioturbation of original lamination by localized and active internal growth of Microcoleus (stained with methylene blue) in laminated mats built by Scytonema. Tidal marshes, eastern Andros Island, Bahamas. Note how this phase of internal growth deformed both over- and underlying laminae. Plane polarized light. Scale bar = 500pm.

other permanent or fugitive carbonate phases; indeed, the degree of complexing of the calcium by the organic acids as well as the rate at which the calcium can be released, influences the rate of precipitation of the carbonate, hence the mineral phase. Furthermore, the release of hydroxyl groups by organic acids during the complexing of calcium ions, may originate at places significant variations of pH (see Trichet, 1967, 1972b).


(5) The size, the crystallographic habit as well as the morphology of crystals (angles, sharp, rounded, etc.) are directly ruled by the physical and chemical properties of the algal mucilages in which precipitation takes place; in their turn, these properties vary with the horizon considered in the mat, with the seasonal metabolical modifications of the given cenose, with the taxonomic composition of the stratified cenoses, and so on. Finally, local variations in pH, concentration of C02, nitrates, sulfates can originate various degrees of corrosion in given parts of the mats, and reprecipitation of carbonate elsewhere.

I t is clear that all these processes of autobioturbation (see also Golubic, Ch. 4.1) can also proceed in other cryptalgal situations, for instance in thrombolites.

Accordingly, without considering the action of external physical factors, a cryptalgal structure is always the total sum of two conjugated and interrelated processes : surface outward accretion (determining in stromatolites the superposition of successive historical laminae), followed by constant re- adjustment of the inner mat-dwellers which distribute themselves, from top to base, into distinct biological and energetical zones. Because of the multiple and sometimes antagonistic properties of these stratified microcommunities, the net result of the interference of this stratification in space on the accretionary stratification in time, will vary. Vertical accretion will be enhanced if, for example, one of the inner communities develops an internal growth phase, or induce a microenvironment in which precipitation and cementation will take place and consolidate the structure. The net result will be negative when one of the microcommunities takes the lead to determine solution processes.

REMARK ON MUD MOUND

Studies in progress on Devonian micritic mud mounds strongly suggest that these would be unlaminated blue-green algal bioherms as evidenced by the abundance of filaments in the micrite, and the organic-like aspect of associated peloids. In first approach, they could then illustrate a mega- development of massive fabrics.

ACKNOWLEDGEMENTS

The present study has been achieved thanks to Grant No. 10215 from the National Foundation for Scientific Research (Belgium) for which I am very grateful.

I am particularly indebted to Conrad Gebelein, Steve Golubic and Phil Playford for exchanging experience on Recent marine stromatolites as well as for their extensive collaboration in sending me samples from Shark Bay without which this paper would have been more incomplete yet.


Chapter 5.2

AN A’ITEMPT TO CLASSIFY LATE PRECAMBRIAN STROMATOLITE MICROSTRUCTURES

J. Bertrand-Sarfati

INTRODUCTION

The stromatolite primary lamination reflects the growth pattern of the algal coenose and the habit of the carbonate precipitated or trapped within the filament framework, among other things. The fossil stromatolite microstructure includes furthermore the imprint of subsequent diagenesis.

In fact, the significance and the functioning of the algal community building up the stromatolite laminae have been often neglected by geologists. In the ancient and especially in the Precambrian, because of the bad preservation of the “microstructure”, two ways have been followed. A carefully detailed description of the actual fabric of the microstructure mixing up original and diagenetic features (Russian school of Moscow) leads to an excessive number of microstructural patterns and a very artificial classification. A detailed description of the recognizable organic remnants (Vologdin, 1962), which are often suspect, neglects the building framework of the laminae. Recently an attempt has been made to introduce a genetic element to interpret more precisely the laminations (Bertrand-Sarfati, 1972c; Preiss, 1972b, 1973b; Walter, 1972a; etc.). The need of modern models was great because of the excessive attention paid by the geologists to the sedimentary and physical processes. Only few of them (Hommeril and Rioult, 1965; Monty, 1965a, 1967; Gebelein, 1969; Davies, 1970b; Neumann et al., 1970; Walter et al., 1972, 1973; Awramik, 197313) described precisely the modem algal laminae and the role of the algae or the algal community. And even then, the fossilization of the laminae is often not observed.

Having had the chance of working on Late Precambrian stromatolites with well-preserved microstructures, I proposed a classification that was well adapted for some of them: (1) simple microstructures where the laminations or the dominant fabrics follow the rhythmical growth pattern of the coenose; and (2) complex microstructures where the historical succession of laminations presents microstructural changes due to seasonal differentiation of the algal coenose.

h Fig. 1. Microstructure occurring as dark films. A. Conophyton ressoti var. jacqueti Sougy. B. Baicalia mauritanica Bertrand-Sarfati, 9 early Late Riphean, northern edge of Taoudenni Basin (scale-bar: 0.5 mm). g

9 =!

FABRIC AND MICROSTRUCTURE

SIMPLE MICROSTRUCTURE

253

Film microstructure

The essential feature of this type of microstructure is represented by regularly banded dark, thin (mode 0.003 mm), micritic films. These micritic films alternating with clear sparitic (or microsparitic) laminae (0.03-0.05 mm) constitute the first-order lamination (Conophyton ressoti Menchikov, Fig. 1A). A second-order lamination occurs with the intercalation of much thicker layers of clear sparite (up to 2 mm wide). A variant of this last one shows groups of films, disrupted by diagenetic processes (Baicalia mauritanica Bertrand-Sarfati, Fig. 1B) in large sparry carbonate layers.

Similar microstructures are found in some forms from the late Middle Riphean and the early Late Riphean: (a) Baicalia lacera Semikhatov, films composed of fibers (“razdel’no voloknistaya”: Raaben, 1972, in Raaben anc Zabrodin, 1972), with sporadic clear layers containing peloids; (b) Cono- phyton garganicum australe Walter (banded microstructure), where the films are regular with a thickness of 0.03-0.05mm. In some other forms of the Late Riphean and the Vendian, thin films of dark micrite occur within laminations of a more complex constitution : Inzeria nyfrieslandica, Tungussia indica, and Poludia polymorpha Raaben.

We can find two present-day examples which, when fossilized, would display a film microstructure: (a) In mats dominated by Schizothrix calcicola in the sink holes of the Bahamas (Monty, 1965a, p. 204, fig. 20). The algal laminations were formed by the “daily” superimposition of thin calcified films, composed of very thin filaments arranged parallel to the lamination. (b) In the effluent of a geyser in Yellowstone National Park, stromatolites resembling Conophyton (Walter et al., 1972; Ch. 6.2 herein) are built of very thin films (5-10pm). They are made of filaments of bacteria and cyanophytes. In the two cases we find thin filaments and the repetition of the same laminae as for our ancient stromatolites.

Tussock microstructure

In this type an irregular lamination is defined by the juxtaposition of separate hemispheric tussocks of different size (0.01-1 mm). Although composed of interlocking crystals, they retain traces of the primary radiating filament fabric. The tussocks are commonly overgrown either by a cement of pure sparite (Tungussia globulosa Bertrand-Sarfati, Fig. 2A) or a dark film (Tarioufetia hemispherica BertrandSarfati) or embedded in carbonate cement and detrital quartz (Serizia radians Bertrand-Sarfati). On these substrates, new tussocks begin to grow and so they define roughly parallel laminations. Some of the tussocks have a concentric growth pattern.

254 J. BERTRAND-SARFATI


Outside the Taoudenni Basin (W. Africa), the only other comparable known microstructure comes from the Late Riphean: Alternella hyperboreica Raaben (intertwining films, “uzorchato-plenochnoy ”: Raaben, 1972, in Raaben and Zabrodin, 1972). However, in this case, the tussocks seem to be without filaments and occur as flat pillows superimposed randomly. The dark films are moulded onto the surface of the pillows and seem to anas- tomose.

Recent colonies of Riuularia which form tiny tussocks encrustating various hard substrates in streams (Kann, 1941a; Bertrand-Sarfati, 1972c, plate XXIX, 4) have a microstructure of juxtaposed hemispheric tussocks. The Riuularia colonies are formed by radial growth of numerous trichomes, periodically embedded in calcite crystals (Fritsch, 1945; Monty, Ch. 5.1). The colonies, when indurated, often show a concentric pattern, due to different stages of growth. When the growth is interrupted, the colony may become covered by a coating of sediments.

Vermiform microstructure

In this kind of microstructure, thick, regular, pale and dark layers have the same composition, both consisting of “narrow, sinuous, pale-coloured areas (usually of sparry carbonate) are surrounded by darker, usually fine-grained areas (usually carbonate)” (Walter, 1972a, p. 14). Such a microstructure occurs in Madiganites mawsoni Walter (Fig. 2B). Diagenetic alteration breaks up the sinuous patches giving the layers a gnunelous aspect.

Vermiform microstructures are found in Vendian and Cambrian forms: (a) Acaciella angepena Preiss, more regularly laminated, (b) Boxonia gracilis Korolyuk ( “globulyarnaya”: Raaben, 1972, in Raaben and Zabrodin, 1972), with granules more rounded than vermiform; (c) Uricatellu urica Korolyuk, with a polygonal network; (d) Boxonia diuertata Korolyuk, with a more variable network. Older forms such as Minjaria procera Semikhatov can show this structure sporadically.

According to Walter (1972b, p. 159), the vermiform microstructure can be compared with the filaments of Giruanella, in spite of the fact that the clear patches of the vermiform microstructure are moulds of large filaments. Giruanella is not well defined with reference to an actualistic model. The structure has been compared (Monty, 1965a, p. 231, plate 35) with that produced by compaction and collapse (3 cm below the surface) in laminae composed of intertwined filaments of Scytonema and Schizothrix. This microstructure is most probably a polyspecific one and remains obscure.

Fig. 2. A. Microstructure consisting of filamentous tussocks: Tungussia globulosa Bertrand-Sarfati (scale-bar: 1.25 mm). B. Vermiform microstructure: Madigunites mawsoni Walter, Middle to Late Cambrian, Jay Creek Limestone?, Jay Creek, Amadeus Basin. Width of field of view is 2.7 mm.

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COMPLEX MICROSTRUCTURE

J. BERTRAND-SARFATI

The stromatolite is built of variable laminae and is probably controlled by one or several species or several coenoses.

Microstructures consisting of micritic mats

Basically the laminae are thick, darkish in colour and composed of micrite; their upper boundary may be straight, crinkled or scalloped. The laminae are formed (Tungussia nodosa Semikhatov Fig. 3A) of two different layers, the upper one (micritic) is darker than the lower one (microsparitic). Many variants of this type are found in the same stromatolites: dark films at the top of the micritic laminae; rhythmical thick micritic layers separating lessdefined micritic mats; trapping (?) of peloids in the basal laminae. Fre- quently, sheet cracks separate adjacent laminae; these are filled with clear calcite. Diagenetic processes also disrupted the mats, which now appear patchy.

Complex microstructures of this type are recorded in numerous forms: Jurusania burrensis, Baicalia burra and Acaciella augusta Preiss (banded microstructure) ; Jurusania nisuensis and J . cylindrica Krylov (disc-like microstructure, “plastinchataya”, or amorphous films, “amorfno-plenochnoy”: Raaben, 1972, in Raaben and Zabrodin, 1972; J. judomica Semikhatov), Baicalia capricornia and Inzeria intia Walter, and Gymnosolen ramsayi Steinmann (this one often has crinkled mats). Numerous forms of this type are known throughout the Riphean and Vendian; their usefulness in stratigraphy may be increased when more information on their possible range of variation is available.

This kind of microstructure is quite difficult to interpret. Is the micrite an original pattern of growth of the calcite crystals or diagenetic micritization? Is the “doublet” the differential growth of one species of algae, or the succession of two dominant species? One possible comparison is proposed with Scytonemacrustaceum mats in the Bahamian upper tidal flats (Monty, 1965a, fig. 33, Pi. 55). The erect filaments form thick mats followed by calcitic cryptocrystalline layers accumulated around unicells and scarce Scyto- nema filaments. The erect filaments in the mat are completely reduced afterwards and enclosed within crystals as impurities. The variations can be due to trapping of particles within the calcitic layer or intercalation of a felt of Schizothrix.

Ca tagrap h- bearing mats

These thick mats appear periodically in a succession of micritic mats. They contain catagraphs in a microsparitic matrix (Tifounheia ramificata Bertrand-Sarfati, Fig. 3B). Catagraphs are quite regular spheroids with a dark

Fig. 3. A. Microstructure in micritic mats: Tungussia nodosa Semikhatov (scale-bar: 0.5 mm), Late Riphean, northern edge of Taoudenni Basin. B. Catagraph-bearing microstructure: Tifounkeia ramificata BertrandSarfati, late Late Riphean, northern edge of Taoudenni Basin (scale-bar: 0.5 mm). N

Cn 4

258 J. BERTRAND-SARFATI

envelope and clear calcite infilling. They are present also in the carbonate between the columns, where they are bigger. Sometimes they cannot be distinguished from micritic peloids.

This microstructure is a dominant feature of a number of Vendian forms: Boxonia grumulosa, B. lissa Komar, Linella simica Krylov and above all Box- onia ingilica Semikhatov, where the laminations include numerous real catagraphs. Some other forms occasionally contain peloid-bearing layers (Linella ukka Krylov and Baicalia rara Semikhatov); these peloids seem to be trapped particles.

This model is not very different from the preceding one, as far as the micritic laminae are concerned, but it differs by the presence of the catagraph layer. Are they ooids or peloids, or really calcified colonies of algal unicells with organic matter migrating on the surface? At present unicells proliferate at the bottom of Scytonema mats (in fresh water) and during recrystallization show this kind of feature (Monty, 1965a, plate 60). In this case, the catagraphs developed in situ, inside the micritic mats. But it is also possible that catagraphs are trapped on the surface of the mat, in the way that trapping occurs in tidal-flat algal mats (Hommeril and Rioult, 1965; Monty, 1965a; Gebelein, 1969; Neumann et al., 1970; etc.).

DISCUSSION

(1) It is now clear that consideration of the microstructure must enter into the definition and classification of stromatolites. However, beyond the simple description of the mineralogy and petrofabrics of the microstructure, an attempt must be made to elucidate its genesis, and also the diagenetic alterations.

(2) The nature of the microstructure has become more and more important in the definition of stromatolite forms (Conophyton, Komar et al., 1965) and even groups, especially in the case of simple microstructures. The delimitation of groups is more difficult in the case of complex forms and usually the stromatolite has to be determined on the basis of several morphological attributes (Hofmann, 1969a; Walter, 1972a).

(3) Attempts at comparing fossil stromatolite microstructures with those of present-day algal communities are faced with several difficulties: (a) a paucity of description of present-day algal microstructures; (b) overemphasis of the importance of intertidal polyspecific algal mats whose structure is largely controlled by sedimentation rates; (c) difficulties for ancient stromatolites in defining basic types of microstructure due to frequent recrystallization, and lack of models.

(4) At a first glance, some kinds of microstructure are easy to interpret: those presenting regular simple microstructures such. as films or tussocks or more complex vermiform microstructures, which are evenly repeated all


along the stromatolite column. In the other cases, microstructures are composed of a historical succession of various kinds of laminae and the basic features of microstructure and the relation with other (variant) laminae are difficult to understand, particularly in the Precambrian. The seasonal and climatic variations played an unknown role in this kind of microstructure.

(5) Various approaches have been used to study the relationship between morphology and microstructure. Krylov (1967a) was the pioneer. For Serebryakov and Semikhatov (1974, p. 567) “they are a combination of various structures originating under various conditions as the result of the life activity of one community (or species) of algae (Krylov, 1967a, 1972; Serebryakov, 1971; Serebryakov et al., 1972)”. Monty ( 1 9 7 3 ~ ) used the terminology of eurybiontic stromatolites for which different morphologies have originated due to environmental factors, and stenobiontic stromatolites which show very few morphological variations and are restricted to a constant environment.

(6) It is already clear that stromatolites with simple microstructures are widespread and seem to have a restricted stratigraphical range:

(a) Stromatolites with dark film microstructure appear all over the world, with a large range of morphological variation, but in a narrow interval of time, from the late Middle Riphean, to the early Late Riphean. They show great adaptability of the coenose to a wide variety of environments.

(b) Filamentous tussock microstructure building columns with little morphological differences seem to inhabit more restricted environments during the Late Riphean.

(c) During the Vendian and later, all the Cambrian vermiform microstructures show very little morphological variation also suggesting growth in a restricted range of environments.

In conclusion, the definition and classification of stromatolite microstructures of the Late Precambrian confirm that the lamination and therefore the construction of stromatolites is often controlled by the algae (Raaben, 1964a; Monty, 1965a, 1972, 1973c; Krylov, 1967; Bertrand-Sarfati, 1972c; Walter, 1972a; Walter et al., 1973; Serebryakov and Semikhatov, 1974). This is in conflict with the assertion that lamination (Logan et al., 1964), columnar structure and ramification (Roehl, 1967) merely reflect variations in the activity of currents and sedimentation rates.

6. MORPHOGENESIS

Chapter 6.1

STROMATOLITE MORPHOGENESIS IN SHARK BAY, WESTERN AUSTRALIA

Paul Hoffman

INTRODUCTION

Hamelin Pool, a hypersaline lagoon in Shark Bay, is apparently the only occurrence of modern stromatolites of comparable size and diversity with Proterozoic stromatolites. The Hamelin Pool stromatolites grow in environmental settings ranging from wave-raked shorelines of high physical stress, to protected tidal ponds of low stress. The most impressive stromatolites are in the high-stress environments and their morphogenesis is controlled by physical processes. Smaller stromatolites and sheet-like cryptalgal deposits occur in settings of low stress, where morphogenesis is dominated by the biotic processes of the algae.

The Hamelin Pool stromatolites were first described by Logan (1961); enlarged upon by Logan et al. (1964); and treated comprehensively by Logan et al. (1974). What follows is a condensation, eliminating much important detail, of the most recent publication.

BATH YMETRY AND ENVIRONMENTAL PROCESSES

Shark Bay (Fig. 1) is a bilobate embayment of the Indian Ocean on the west coast of Australia. Hamelin Pool, the southeasternmost lobe, is silled by a sea-grass barrier at its mouth. The climate is semi-arid and freshwater nm-off from inland is negligible. For these reasons, the salinity of Hamelin Pool is nearly twice that of normal seawater. In the less saline parts of Shark Bay, invertebrates graze the sediment surface and prevent the growth of algal mats in the lower intertidal and sublittoral zones. But in Hamelin Pool, the grazing invertebrates are inhibited by the high salinity and, as a result, algal mats grow in a broader range of environments than elsewhere.

262 P. HOFFMAN

supratidal flats

beach ridges

intertidal algal mats

sublittoral shelf

10 krn O A

’ ..._... .... ... . p y F ..... ;I.. ;!

HAMELIN POOL

Fig. 1. Map of Shark Bay showing the location of Hamelin Pool and the sea-grass banks and sills (stippled).

Fig. 2. Bathymetric subdivisions of Hamelin Pool.

Bathymetric subdivisions

Hamelin Pool (Fig. 2) consists of: (a) a central basin floor, 8-10m deep; (b) a sublittoral shelf, less than 5 m deep; and (c) a coastal strip, in the intertidal and supratidal zones.

The coastal strip is made up of skeletal and ooid sand, produced mainly on the sublittoral shelf but driven onshore by waves during cyclonic storms and the prevailing southerly gales. The coastal strip is differentiated along its length according to different intensities of wave attack:

(1) The most intense wave attack is at r o c b headlands, where the sublittoral shelf is relatively narrow and deep. Wave attack is multidirectional. The coastal strip is less than 100 m wide, and sediment is coarse and transi- ent.

MORPHOGENESIS 263

(2) Intermediate are the beaches and sandflats located in arcuate bights, where the sublittoral shelf is relatively broad and shallow. Waves are refracted onto an onshore track during their passage across the shelf so that wave scour, concentrated along the seaward margin of the coastal strip, is perpendicular to the shoreline. Sedimentation rates are high and the intertidal sandflats prograde, producing a moderately broad supratidal zone, typically lined with beach ridges of bivalve coquina. Continued progradation will ultimately engulf the headlands with sediments and straighten the coastal strip, reducing its lungitudinal differentiation.

(3) Most protected from wave attack are shorelines that face north (e.g. Nilemah Embayment) and those around tidal ponds behind barrier spits (e.g. Hutchison Embayment). Tidal currents may be strong in the outlets of the ponds but wave action is negligible. The supratidal zone is very broad, up to 2000 m, and generally lacks beach ridges.

Factors controlling algal-mat distribution

Algal mats are most prominent in the intertidal zone but, in places, extend into the lower supratidal and shallow sublittoral as well. There are seven basic types of mat (Fig. 3), each a complex community of several species of blue-green algae. Each mat type is dominated by a unique combination of species. Each is distinctive in its appearance, in the way it traps and/or binds sand to produce stromatolites, and in the laminated or unlaminated internal fabric it imparts to the stromatolites.

The distribution of mats and their zonation by type are controlled by the following physical processes :

(1) Desiccation. The zonation of mats is controlled most of all by the desiccation gradient (Fig. 4). As desiccation increases, the number of algal species decreases, thus allowing different combinations of species to dominate each successive mat type. The gradient is related mainly to elevation, the higher tidal flats being less frequently flooded than the lower ones. How- ever, even the higher flats have depressions where water is ponded and desiccation is low.

(2) Sedimentation. A secondary control of the mat zonation is the sedimentation rate (Fig. 4). In certain areas, particularly where the sublittoral shelf is very broad, sediment influx to the coastal strip is so rapid that algal growth cannot keep pace, and the intertidal zone is barren of mats.

(3) Cementation. Precipitation of intergranular aragonite results in extensive lithified crusts in the intertidal and, to a lesser extent, sublittoral zones. Typically, the mats colonize exposed crusts because they offer a more cohesive substrate than loose shifting sand. Once established, however, the mats resist erosion in even the most turbulent environments. Furthermore, the stromatolites are themselves rapidly lithified during growth, accounting for their remarkable wave resistance. Thus, cementation is doubly important,

264 P. HOFFMAN

’ILM

Hormatonema cryptocrystalline rinds vrolacco -rugrum on lithified crusts

- - .. - - - - -- 3LISTER

_. - inorganic laminations - -- Microcoleus chtonop/sstes ----- - - with algal partings

etc

TUFTED

Lyngbya mstuerii etc

scallop fabric

GELATINOUS ,--

Symploce iwte-viriciis - ribbon fabric with ? etc - - ----

polygonal prism cracks

SMOOTH

Schizothrix helve laminated fabric etc with fine fenestrae

COLLOFORM

Mictvcokws digitate fabric with coarse fenestrae tmnwrimus

etc

Fig. 3. Name, dominant bluegreen algal species, and characteristic internal fabric of the seven mat types in Hamelin Pool.

both in providing sites for the initiation of mats and in subsequent lithification of mat-bound sediment.

(4) Erosion. Wave scour and, to a lesser extent, tidal run-off serve to expose cemented crusts to mat colonization. Wave scour also tends to prevent accretion in the depressions between stromatolites, thus allowing the tops of the stromatolites to grow well above the surrounding substrate. Where waves are strongly refracted, unidirectional wave scour results in elongate stromatolites.

MORPHOGENESIS 265

supratidal

intertidal SMOOTH MAT

I sublittoral COLLOFORM MAT

(open coastline) (tidal pond)

Fig. 4. Distribution of mat types on an environmental process grid. Mat types in capital letters produce columnar structures; those in lower case only stratiform cryptalgal sheets.

DISTRIBUTION OF STROMATOLITE MORPHOTYF'ES

The stromatolites are most readily classified according to the mat types that build them: (a) colloform mat structures; (b) smooth mat structures; (c) pustular mat structures. The other 4 types produce only stratiform sheets. The distribution of stromatolite morphotypes is summarized in Fig. 5.

Colloform mat structures

Stromatolites of diverse morphology occur on the sublittoral shelf to depths of 2 m below mean sea level. All are formed by the colloform mat and inherit its coarse fenestrate internal fabric of multiconvex laminations. The mats support a dense growth of.the green alga AcetabuZaria and are populated by serpulid worms and small bivalves. These disrupt the mat growth and contribute to the generally cavernous margins and irregular morphology of the stromatolites.

At rocky headlands, irregular columns, up to 1.0m in relief, having discrete bladed branches with flattened tops, grow on cemented crusts of bivalve coquina. In places, the columns coalesce to form bioherms, many meters in diameter, with a continuous outer shell.

In the bights, the stromatolites become more prolate and generally have only 0.5 m relief. Individual structures are' elongate perpendicular to the shoreline (i.e. parallel to the direction of wave scour), but they may be arranged in rows parallel to the shoreline. The structures commonly coalesce laterally along the length of the rows, producing compound masses with elongation normal to that of the primary structures.

On the sublittoral floor of the tidal pond in Hutchison Embayment, stromatolites with a digitate columnar internal structure occur beneath ovoid patches, less than 0.5m in diameter, of colloform mat. Unlike the

266 P. HOFFMAN

HEADLAND

. . . . . . . . .

BIGHT EMBAYMENT

Fig. 5. Generalized distribution of stromatolite morphotypes.

free-standing stromatolites on the open shelf, these masses are almost buried in pellet mud.

Smooth mat structures

Stromatolites built by the smooth mat have a fine fenestrate internal fabric of simple convex laminations. Except in protected embayments, the lower intertidal substrate is so unstable that the mats initially colonize only lithified crusts and crust fragments. Once established, the mat-bound sediment is itself rapidly cemented. Discrete columnar structures prevail because the mats are unable to breach the loose sand in depressions between the columns.

The stromatolite columns are largest at rocky headlands, attaining a relief of 0.8 m. The columns are circular in plan, and tend to broaden and coalesce at the top (Fig. 6a), particularly where they project into the upper intertidal zone of pustular mat.

Toward the bights, the relief of the columns is greatly reduced and the stromatolites become strongly elongate perpendicular to the shoreline (Fig. 6b). Some bights have bands, oriented parallel to the shoreline, of

MORPHOGENESIS 267

“ridge-and-rill” structures (Fig. 6c). The “ridges” are extremely elongated stromatolites of low relief, oriented perpendicular to the shoreline, separated by “rills” of rippled sand. The seaward ends of the ridges tend to be swollen and oversteepened, and some of the ridges project seaward into rows of small discrete columns tilted seaward (Fig. 6d). Many of these columns widen upward and branch (Fig. 6e). In places, calyx-like structures occur where stromatolite columns have been decapitated, their less well cemented centers hollowed out, and their r i m s recolonized by mats.

Pustular mat structures

Upper intertidal structures built by pustular mat have an irregular fenestrate fabric without laminations. Most of the structures are inherited from lower intertidal ones as a result of their upward growth. They range from high-relief circular columns at headlands, through medium-relief elongate columns, to low-relief ridge-and-rill structures in bights. The upper intertidal ridges tend to be broader and more sinuous than those in the lower intertidal. In places, the rills are colonized by gelatinous or tufted mats, whereby the unlaminated ridges become flanked by laminated cryptalgal deposits.

Stratiform cryptalgal sheets

Wherever wave or current action is too weak or impersistent to prevent mats from colonizing loose sand, mats form continuous undifferentiated sheets that grow without forming structures with significant relief. This occurs throughout the intertidal zone of protected embayments and in protected parts of the upper intertidal zone in bights. The internal fabric of the sheets is controlled by the mat type’ (Fig. 3), distributed as follows: (a) smooth mat in the lower intertidal zone; (b) pustular mat in the upper intertidal zone; (c) blister mat in the lower supratidal zone; (d) gelatinous mat in poorly drained depressions and ponds; and (e) tufted mat in better drained depressions.

STROMATOLITE MORPHOGENESIS : SUMMARY AND CONCLUSIONS

From Fig. 5 , it is clear that the gross morphology of the stromatolites is not controlled by the type of algal mats. Smooth mat alone, for example, produces structures ranging from stratiform sheets to high-relief columns. Conversely, similar morphologies, the distinctive ridge-and-rill structure for example, are built by as biologically distinct mat types as the filamentous smooth mat and the coccoid pustular mat. The mat type does control the internal fabric of the stromatolites, filamentous mats producing laminated structures and coccoid mats unlaminated ones. But even the distribution of

268 P. HOFFMAN

MORPHOGENESIS 269

Fig. 6. Stromatolites near Carbla Point, 10 km south of the Hutchison Embayment. a. Circular columns, 0.8 m in height, at a rocky headland. b. Columns elongate perpendicular to the shoreline, in a bight near a headland. c. Low-relief ridge-and-rill structures, seaward to the left, in bight away from a headland. d. Small tilted columns, seaward to the left, in a bight. e. Detail of branching columns in a bight.

270 P. HOFFMAN

mat types is ultimately related, as is the stromatolite morphology, to physical processes of the environment.

The following generalizations may be useful in interpreting the environment of ancient stromatolites:

(1) Stratiform cryptalgal sheets occur where wave and tidal scour are weak. (2) Discrete columnar structures occur where wave and tidal scour are

strong. (3) Relief of columnar structures is proportional to the intensity of wave

action. (4) Elongation of simple stromatolites is parallel to the direction of wave

and tidal scour, generally perpendicular to the shoreline. (5) Elongate stromatolites tend to occur in rows or bands parallel to the

shoreline. (6) Stromatolites tend to be oversteepened or tilted seaward into the

oncoming waves. Thus, shorelines can be subdivided into three bathymetric types according

to their stromatolite assemblages (Fig. 7): (a) shorelines fully exposed to waves have large columns (diameter 30-100 cm) of high relief (greater than

Fig. 7 . Change in stromatolite morphology from circular columns at a headland to bands of ridge-and-rill structures in a bight.

MORPHOGENESIS 27 1

50 cm); (b) shorelines with a shallow sublittoral shelf offering partial protection from wave attack have small tilted columns (diameter 5-15 cm) and ridge-and-rill structures of moderate relief (10-30 cm); ( c ) shorelines completely protected from waves by barrier islands or spits have stratiform sheets of low relief (less than 5 cm).

CAUTIONARY POSTSCRIPT

Although Hamelin Pool contains the most diverse assemblage of modern stromatolites, many ancient stromatolites doubtless grew in environments not represented there. This is particularly true for sublittoral settings of low environmental stress, where the conclusions regarding stromatolite morphogenesis in shoreline environments, listed above, may not be applicable.

6. MORPHOGENESIS

Chapter 6.2

MICROBIOLOGY AND MORPHOGENESIS OF COLUMNAR STROMATOLITES (CONOPHYTON, VACERRILLA) FROM HOT SPRINGS IN YELLOWSTONE NATIONAL PARK

M.R. Walter, J . Bauld and T.D. Brock

INTRODUCTION

The hot springs of Yellowstone National Park in Wyoming constitute an environment that in some ways is analogous to Precambrian marine environments. There is widespread inorganic deposition of silica, as during the Precambrian (Holland, 1971; Walter, 1972b). At higher temperatures there are no animals and the biota consists of cyanophytes and bacteria, as during much of the Precambrian (e.g. Schopf, 1970). I t is in this “primitive” environment that we have found two kinds of columnar stromatolites; one of these is Conophyton, a group previously thought to have stopped forming near the end of the Precambrian; the other is Vacerrilla, a previously undescribed group containing only one form.

The great amount of microbiological data on the Yellowstone springs (e.g., Castenholz, 1969; Brock, 1970) provides an excellent opportunity to study in depth the microbiology of columnar stromatolites. Furthermore, the significant environmental features of the outflow channels and pools can be determined readily. I t is thus possible to attempt to relate stromatolite morphology to both the environment and the constructing microbiota, and to attempt to resolve the conundrum of stromatolite morphogenesis, at least for two groups of columnar stromatolites.

Earlier results of our work showed that. stromatolites can be primarily siliceous (Walter, 197213) and that bacteria can build stromatolites (Walter et al., 1972; Doemel and Brock, 1974), thus fulfilling the primary objectives of the study. The discovery of Conophyton was unexpected but is of great interest. This is one of the most distinctive groups of Precambrian stromatolites. We now have the opportunity to analyse its morphogenesis to discover what biological and environmental information is stored in the Precambrian forms.

Walcott (in Anonymous, 1916) and Barghoorn and Tyler (1965) have also looked to Yellowstone for Recent analogues of the Precambrian stromatolites

274 M.R. WALTER ET AL.

they were studying. The Yellowstone stromatolites were first described in detail by Weed (1889a), in a comprehensive monograph on Yellowstone sinters. Korde (1961) has commented that Weed’s illustrations show some Conophyton-like stromatolites. The Yellowstone conophytons are similar to those described from the Precambrian by Hofmann (196913) and the Precam- brian conophytons illustrated by Vlasov (1970) are astoundingly like those in Yellowstone.

The interdisciplinary nature of this study presented certain problems based largely on the different experience of the three investigators involved. All of these problems were resolved except one: we were unable to agree that stromatolites should be given Linnean binary names. For this reason the stromatolites are named in an appendix to this chapter, under the sole author- ship of M.R. Walter. In his defence this author refers to the arguments of Cloud and Semikhatov (1969b), which apply as much to Recent as to ancient stromatolites.

METHODS

(1) Study areas

Investigations were carried out at the following localities in Yellowstone Park: Main Spring, Jupiter Terrace (Mammoth Geyser Basin); Queens Laundry Spring (Sentinel Meadows); “Weeds Pool” (Fountain-Paint Pots area, 600 m W of Celestine Pool); Grand Prismatic Spring and an unnamed spring, 900 m SE of there on the east bank of the Firehole River (Midway Geyeer Basin); “Conophyton Pool”, 1OOOm W of the south end of Goose Lake, “Column Spouter”, 800 m NNW of the north end of Goose Lake, two unnamed springs 900m NW of the north end of Goose Lake (Fairy Creek meadows); White Crater Spring (Shoshone Geyser Basin). The names in quotation marks are unofficial. For localities see U.S. Geological Survey topographic map “Yellow- stone National Park” and the 15’ quadrangle maps.

(2) Environmental parame ters

Temperature, pH and light intensity were measured as described in Bauld and Brock (1973). Water-flow patterns and velocities were determined by injecting into the water a concentrated solution of methylene blue dye, and by photographing the distribution of the dye at intervals of 1-3 seconds. In some cases it was sufficient to time the passage of the dye along a measured course, without photographing the distribution pattern. Vertical flow gradients in relatively deep pools were determined by injecting the dye at several depths.

MORPHOGENESIS 27 5

(3) Growth rate studies in the field

To study the rate of formation of the stromatolites and the temporal significance of their laminae, fine carborundum powder was sprinkled over them to provide reference laminae. About fifteen sites in various pools were marked and sampled on several occasions. Soft, flat microbial mats were sampled with a cork borer and heavily mineralized mats were cut with a knife. The mats were marked during August 1971 and April, June and July 1972. Samples were taken during April, August and November 1972. One- to three- day experiments to check for diurnal lamination were run at three localities.

(4) Identification of the microorganisms

Morphology Morphological observations were made using a Zeiss RA or a Zeiss Univer-

sal microscope equipped with phase-contrast and differential interference contrast optics. The cyanophytes were classified using Copeland's (1936) descriptions of Yellowstone species.

Enrichment and isolation procedures Stromatolite samples from a variety of hot-spring effluents and pools were

treated in the following manner. Algal enrichments were obtained by inocu- lating 10 ml of medium D (Castenholz, 1969) with natural material. The glass, screw-capped culture tubes containing the inoculated medium were incubated at temperatures of 30-40°C,' or 50°C, under fluorescent illumination. Cultures were examined by phase-contrast microscopy after 3-4 weeks. Unicellular blue-green algae are preferentially selected in liquid enrichment cultures (Stanier et al., 1971); to obtain pure cultures of the filamentous alga Phormidium , macerated stromatolite material was spread along a single line on the surface of an agar block (1% Difco Bacto-agar, medium D). The inoculated agar surface was then placed in a light gradient. Phormidium filaments moved towards the light by gliding motility, away from immotile algae and bacteria. The most active filaments were transferred to a fresh agar surface and the process was repeated several times. Purity was assessed microscopically and by the absence of bacterial growth upon inoculation into suitable test media. Chloroflexus was enriched from stromatolite samples by inoculation into screw-capped culture tubes completely filled with medium D supplemented with 0.05% (w/v) yeast extract, 0.05% (w/v) NazS.9Hz0 and buffered with 0.5g/l of glycyl-glycine. The tubes were incubated at 5OoC under tungsten illumination. Cultures were examined microscopically after 3-4 weeks. Pure cultures of Chloroflexus were obtained under aerobic conditions as described by Bauld (1973).


Pigment analyses Pigments were extracted from stromatolites and mat samples with meth-

anol or acetone as described in Bauld and Brock (1973). Absorption spectra for the pigments were obtained with a Beckman DB-G grating.spectrophotometer equipped with an R136 photomultiplier tube having an extended response in th‘e red region. The in uiuo spectra of chromatophore preparations (Pierson and Castenholz, 1971; Bauld and Brock, 1973) were determined by R.W. Castenholz, University of Oregon, using a Cary model 14R recording spectrophotometer.

Fluorescence microscopy Microorganisms from stromatolite samples were examined using a Zeiss

Universal microscope equipped with a mercury arc lamp. When Phormidium and other algae are illuminated with blue exciting light, they can be seen because of the red fluorescence of their chlorophyll a. Chloroflexus, on the other hand, contains bacteriochlorophylls a and c which fluoresce in the near infra-red region of the spectrum and not in the visible region (Pierson and Howard, 1972).

( 5 ) Relative abundances of the microorganisms in the stromatolites

Lengths of algal and bacterial filaments were estimated by a method similar to that of Bott and Brock (1970) using a Zeiss Universal or a Zeiss RA phase- contrast microscope. Samples fixed in 4% formalin at the time of collection were homogenized and placed in a Petroff-Hausser counting chamber; the number of intersections made by filaments with the vertical and horizontal lines of the grid was recorded. Intersections with non-filamentous microorganisms were also recorded to provide an approximate measure of their abundance. Between 700 and 2700 intersections were counted for each sample.

(6) Measurement of 14C02 fixation (photosynthesis) in the stromatolites

Sampling and incubation Stromatolitic tufts 1-2 mm high from “Conophyton Pool” were excised

from the adjoining flat mat and placed in spring water taken from the sampling site. Tuft preparations were placed in 5-ml glass vials in “Conophyton Pool” and allowed to equilibrate for 10 min before adding 0.1 ml of 2.3mM NaH 14C03 (specific activity 8.7 pCi/pmole). The samples were incubated in the environment from which they were obtained. Preliminary experiments had shown that isotope incorporation was linear for up to 2 hours under such conditions.

MORPHOGENESIS 27 7

Differentiation between algal and bacterial photosynthesis In order to estimate bacterial photosynthesis in the presence of Phormid-

ium, 3-( 3,4dichlorophenyl)-1,ldimethylurea (DCMU) was added to the incubation vials to give a final concentration of M , sufficient to inhibit algal photosynthesis completely (Bishop, 1958; Bauld and Brock, 1973). Some incubations were carried out under a Wratten 88A filter in the form of a 10 x 10 cm gelatin layer sandwiched between glass plates. The transmission of light at wavelengths below 720nm was about 0.1% of the total incident light. Use of this filter effectively precludes algal photosynthesis, which does not occur at wavelengths above 700nm. However, transmission of light at 740 nm (in uiuo absorption peak of b;acteriochlorophyll a) was virtually 100% (manufacturer’s specifications).

Processing of radioactive samples and determinations of pigment content in these samples were as described by Bauld and Brock (1973).

(7) Testing for phototaxis

Macerated stromatolite samples, or laboratory cultures, were spread along a line on an agar surface. The agar, contained in a petri plate wrapped in aluminium foil, was placed in a light gradient and oriented so that the “spread line” was normal to the direction of the light. A small hole in the foil allowed light to enter the container. Phototaxis of Phormidium filaments was quantified by measuring the rate of advancement of the front of algal filaments moving towards the light source. The rate of gliding motility of Phormidium filaments from macerated samples of conical stromatolites was also determined by placing a portion of the material on an agar surface, and measuring the movement of individual filaments using a Zeiss RA phase-contrast microscope equipped with a long-working-distance condenser and an ocular micro- meter.

( 8 ) Laboratory growth of conical stromatolites

A bacteria-free culture of Phormidium isolated from a conical stromatolite was used. Cultures were grown under fluorescent lights in a 37°C incubator room where the heat of the lights raised the temperature of the culture vessels to about 40°C. Erlenmeyer flasks containing a thick layer of solid medium D were inoculated with a suspension of filaments. Sufficient inoculum was used so that a confluent growth of the alga occurred after about four days of incubation. At this time, liquid medium D. was carefully added to fill the flasks, and the organisms were allowed to grow up into the liquid. Within 24-48 hours after addition of the liquid medium, small nodes of algal growth appeared on the surface of the solid medium at the bottoms of the flasks. To simulate the periodicity of sunlight a light-dark cycle of 12 hours was used, the light being controlled by an automatic timer. This resulted in the formation of two laminae per 24 hours.


(9) Identification of opaline silica

The siliceous composition of the stromatolites was determined by energy dispersal X-ray microanalysis. X-ray diffractometry showed that the silica is opal A (Jones and Segnit, 1971). Scanning electron microscopy showed that the silica in’ these samples has the spherical habit characteristic of many “opaline” silicas (Jones et al., 1964).

(10) Embedding and thin-sectioning the stromatolites

The stromatolites were dehydrated through an ethanol series, infiltrated with propylene oxide, and impregnated with Spurr Low-Viscosity Embedding Medium B (a blend of epoxy resins). This medium was chosen because of its low viscosity and high solubility in ethanol and propylene oxide (Spurr, 1969) ; these characteristics facilitated thorough impregnation of the samples. “Curing” of the resins was achieved by heating the samples at 70°C for 8 hours in a vacuum oven. In samples thicker than about 2 cm the resins did not polymerize fully and remained soft. After embedding, thin sections of the stromatolites (80-100 Vm thick) were prepared by conventional methods (Kummel and Raup, 1965).

RESULTS

Description of columnar stromatolites

The stromatolites described here are formally named in the Appendix to this chapter. Here the two types are referred to as flat-topped columns and conical columns. Terms used to describe stromatolite morphology are defined in the Glossary (pp. 687-692).

Flat-topped columns (Vacerrilla walcotti)

Mode of occurrence. This stromatolite occurs in sheets 2-3cm thick. The shape of the sheets is determined by the environment in which they occur. In outflow channels they are long, narrow and frequently lobate. On extensive terraces they are irregularly lobate. In the examples known the sheets have maximum horizontal dimensions of about 10 m (see Ch. 8.8).

Column shape. The columns are erect and subcylindrical with flat to gently convex tops (Fig. 1). They are 0.5-1.0 cm in width and have about 1 cm of relief above the substrate. Adjacent columns may coalesce so that in transverse section they are lobate. No branching of columns has been seen.

MORPHOGENESIS 279

Fig. 1. Flat-topped columnar stromatolites in an outflow channel of “Column Spouter”, Fairy Creek meadows. The columns are about 1 cm wide.

Lamina shape. Laminae are gently convex with abruptly deflexed margins. They are rectangular, rhombic or, infrequently, steeply convex (Figs. 2, 3). They frequently coat the column margins, forming a wall (Figs. 2-4). Bridging between columns is moderately frequent and may produce a pseudocolumnar structure.

Microstructure. The lamination is predominantly banded with pale laminae 30-150pm thick and dark laminae 20-60pm thick (Figs. 2-4). A striated microstructure is present locally within some columns: dark lenses 30-150 pm thick are set in a pale matrix. The walls are up to 600pm thick (Figs. 2-4). In them lamination is absent or is diffusely banded to streaky.

Filament orientation. Where the microstructure is banded the algal and bacterial filaments alternate from predominantly subvertical, in the pale laminae, to predominantly parallel to laminae, in the dark laminae (Fig. 5). In the walls the filaments are subvertical ji.e. parallel to the walls). In the dark lenses of the striated microstructure the filaments appear to be randomly oriented. In the light matrix the filaments are subvertical.

Distribution. This stromatolite was studied in the outflow channels of two hot springs: “Column Spouter” (Fairy Creek meadows) and near “Weeds

z. Figs. 2-4. Longitudinal thin sections of flat-topped columnar stromatolites from "Column Spouter". Transmitted light..Scale bar 2 mm. Figs. 2 ,3 , CPC 15460; Fig. 4, CPC 15461. 8

K 3 s

MORPHOGENESIS 281

Fig. 5. Thin section normal to the lamination of a flat-topped columnar stromatolite from “Column Spouter”, showing Phormidium and Chloroflexus filaments (indistinguishable in this figure) forming two laminae, a lower one with horizontal filaments (which formed at night) and an upper one with erect filaments (which formed during the day). The filaments are about 1 pm wide. Photographed with differential interference contrast optics. CPC 15460

Pool” (Fountain-Paint Pots area); it was subsequently recognised in a number of other springs which are mostly unnamed.

Columns with conical tops (Conophyton weedii)

Mode of occurrence. This stromatolite occurs in sheets a few millimetres to 30 cm or so thick. The shape of the sheets is determined by the environment of growth, ranging from long and narrow (in outflow channels) through very irregularly lobate (within large springs) to annular (around the margins of hot springs with little flow). The sheet-like bodies are biostromes in the sense of Walter (1972a). A few are so narrow and thick that they can be called bioherms. The occurrence is further discussed and illustrated elsewhere in this volume (Walter, Ch. 8.8).

Column Shape. The columns are subcylindrical and erect and have conical tops (Figs. 6-16). Transverse sections are subcircular. The columns are about


6

7

Figs. 6, 7 . Tufts, small cones, and ridges at early growth stages of conical stromatolites, “Conophyton Pool”. Scale divisions are 1 cm.

MORPHOGENESIS 283

Fig. 8 . Small cones from an outflow channel of “Column Spouter”. Pine needles indicate scale.

Fig. 9. Subaqueous cones and ridges in “Conophyton Pool”. The nail marks the site of an accretion rate experiment and the dark patch is carborundum powder. Scale divisions are 1 cm.


Fig. 10. Subaqueous cones in Jupiter Spring, Mammoth area. Scale divisions are 1 cm.

0.1-3.5 cm wide and from less than 0.1 to 10 cm high. Most are 1-2 cm wide by 2-3cm high. The smaller examples are usually at least l c m apart; the larger columns are contiguous, or nearly so. Columns less than about 0.1 cm wide are referred to here as tufts (Figs. 6, 7).

This basic shape is frequently modified by the presence of tiny spinose projections and radially arranged, narrow vertical ridges (Figs. 12-16). In some springs (e.g. “Column Spouter”) no ridges or spines are developed; in others there are only very weakly developed ridges (e.g. “Weeds Pool”, White Crater Spring); ridges may be strongly developed (“Conophyton Pool”, Main Spring), in some places to the extent that no columns are recognizable (parts of “Conophyton Pool”) ; spines are restricted to the ridged stromatolites (“Conophyton Pool”).

Lamina shape. Laminae are irregularly conical (Figs. 17-25), with apices ranging from sharp (where the laminae have a radius of curvature of less than 0.2 mm) tp blunt (where the laminae are horizontal for 2-3 mm). The apical line, linking successive apices, is straight to gently wavy (with amplitudes of the flexures ranging up to 1 mm); abrupt displacements of the apical line are moderately frequent. The laminae are mostly ‘smooth, but some are slightly wavy or wrinkled.

MORPHOGENESIS 285

Fig. 11. Large cones from “Weeds Pool”, Fountain-Paint Pots area. Scale divisions are 1 cm.

Laminae in the spines (Figs. 25,26) are widely variable in shape; they range from almost planar (oriented parallel to the associated cone flanks) through irregularly convex to obtusely conical.

Crestal zone. All three types of crestal zone recognized in Precambrian Cono- phyton stromatolites (Komar et al., 1965a) are present in the Yellowstone cones (Figs. 23,24,27-29). The zone of thickening and contortion of the laminae is 0.5-2.0 mm wide in the five specimens in which it was measured. The degree of thickening of the laminae in the crestal zone is difficult to measure because the laminae are frequently diffuse, but it appears that most laminae are about twice as thick in the crests as on the flanks of the cones: thickenings of up to five times were noted.

Microstructure. The microstructure ranges from banded to streaky and from distinct to diffuse; differences in appearance depend partly on the magnification at which the microstructure is viewed (Figs. 17-31). At low magnifi- cations, it normally appears distinctly banded, with laminae 30pm or more thick. Laminae range down to only 1 pm in thickness but most are 5-15 pm

28 6 M.R. WALTER ET AL.

Fig. 12. Ridged cones from “Conophyton Pool”. Scale divisions are 1 cm.

thick. The thinnest laminae are composed of single algal or bacterial filaments encrusted with silica. These laminae were excluded from the measurements presented in Fig. 32 (p. 298).

Subspherical to lenticular thickenings of laminae are relatively irlfrequent and irregularly distributed (Figs. 30, 31). They are mostly 0.5-1.5 mm thick by 0.5-9.0 mm long.

Filament orientation. The algal and bacterial filaments are usually arranged parallel to the lamination. There are three exceptions: (1) subspherical to subcylindrical masses of tangled unoriented filaments occur at column bases (Figs. 17-19); (2) locally on column flanks there are subspherical to lenticular thickenings in which the filaments are rolled up into concentric layers (Figs. 30, 31); (3) in the axial zones filaments tend toward a vertical orientation. This last case has several expressions: there may be a tuft of more or less vertical filaments at the apex of a cone - these are perpendicular to the lamination and may penetrate through overlying laminae as though injected into them (Fig. 29); there may be a mass of unoriented filaments; filaments may be slightly oblique to the laminae, and sigmoidally purved (Fig. 28).

Distribution. Conical stromatolites are known from the following hot springs and geysers in Y ellowstone National Park: “Conophyton Pool”, “Column Spouter” and two unnamed springs (Fairy Creek meadows); Queens Laundry

MORPHOGENESIS 287

13

14

Figs. 13, 14. Cones with ridges and spines, “Conophyton Pool”. Scale divisions are 1 cm.


Fig. 15. Top view of ridged and spinose cones from “Conophyton Pool”. Width of field of view about 10 cm.

Spring (Sentinel Meadows) ; “Weeds Pool” (on the flank of the Fountain-Paint Pots group); Grand Prismatic, Tromp Spring and an unnamed spring (Midway Geyser Basin); White Crater Spring (Shoshone Geyser Basin); and Main Spring (Jupiter Terrace, Mammoth area). In addition, they have been seen in a number of other small, unnamed springs in Yellowstone. Similar structures have been collected from a spring at Hveravellir in central Iceland by T.D. Brock.

Conditions o f stromatolite formation

The stromatolites described here occur in thermal waters with temperatures between 32°C and 59°C and with pH values of 7-9. It was demonstrated by carborundum marking that the stromatolites are actively growing in the environments in which they were observed. Filamentous cyanophytes of the genus Phormidium, which are an essential component of the biotas of both the flat-topped and conical columnar stromatolites, are apparently absent from waters hotter than 59” C (the filamentous photosynthetic bacterium Chloroflexus aurantiacus Pierson and Castenholz forms stratiform and nodular stromatolites at higher temperatures: Doemel and Brock, 1974). Below about 32”C, coarsely filamentous cyanophytes such as Calothrix coriacea Copeland dominate the stromatolitic mats, which are stratiform or nodular.

MORPHOGENESIS 289

Fig. 16. Cones with well-developed radial ridges, “Conophyton Pool”. Scale divisions are 1 cm.

Animals are present at temperatures up to 48-51OC (Brock, 1967c; Casten- holz, 1969; Wickstrom and Castenholz, 1973). These include rotifers, ostracods, spiders and flies. Although some of these organisms graze the microbial mats (Brock et al., 1969), they apparently are not effective in shaping the stromatolites we have studied, as morphologically comparable stromatolites also form at higher temperatures than those at which the animals can survive. Encrustations of silica may protect some of the stromatolites from grazing animals.

Air temperatures in Yellowstone Park average 15°C during mid-summer and - 10°C during mid-winter. The hot spring waters remain free .of ice and snow during winter but, where measured, the water in pools and outflow channels was 2-3OC cooler in winter than in summer. The maximum mid- summer light intensity of 6.1*1051ux drops to a mid-winter maximum of around 1.9*1051ux. This has a marked effect on the pigmentation of many of the cyanophytes in the springs: during summer they are yellow-orange (carotenoid pigments dominating) whereas in winter they are dark green due to an increased chlorophyll content (Brock and Brock, 1969b).

The flat-topped columns were studied during 1971-1972 in three outflow channels: two at “Column Spouter” (Fairy Creek meadows) and one immedi-

290 M

.R. W

AL

TE

R E

T A

L.

Figs. 17-19. Thin sections of tufts and flat mats, which are the initial accretion stages of cones, "Conophyton Pool". Transmitted light. The black layer near the bottom of each photograph and the black, angular grains above the laminated mats are carborundum powder introduced during accretion-rate experiments. The filaments near the tips of the tufts are the cyanophyte Calothrix. CPC 15462.

z s

d t'

F3n s 20 3

Figs. 20-22. Longitudinal thin sections of cones from "Conophyton Pool". Lateral surfaces of the cones were broken during collection and thin-sectioning. There were many bridging'laminae between adjacent cones. Transmitted light. Scale bars 1 cm. Fig. 20, . CPC 15463, Fig. 21, CPC 15464; Fig. 22, CPC 15465

24 23 hl (D Figs. 23, 24. Longitudinal thin sections of cones from “Conophyton Pool”, showing the crestal regions of the laminae. Note the

deformed cone tip immediately above the scale in Fig. 23. Transmitted light. CPC 15466. w

z %

3 2

Figs. 25, 26. Longitudinal thin sections of cones from "Conophyton Pool" showing surficial spines. The spine in Fig. 26 projects $ .

r

s M c3

from the side of a'cone. Transmitted light. Fig. 25, CPC 15467; Fig. 26, CPC 15468.

295

28

27

29 Figs. 27-29. Longitudinal thin sections of cones from “Conophyton Pool” showing the crestal regions of the laminae. Fig. 28 is an enlargement of part of Fig. 27. Note the vertical tuft of filaments in Fig. 29. Transmitted light. Scale bar in Fig. 27, 1 mm. CPC 15468.

z 2'

8 $

Figs. 30, 31. Longitudinal thin sections showing knoh of coiled filaments within the flanks of cones from "Conophyton Pool" Transmitted light. CPC 15464. 2'

3 g

MORPHOGENESIS 297

ately above “Weeds Pool” (Fountain-Paint Pots area). Stromatolites from all these localities are siliceous. “Column Spouter” erupts approximately hourly. During eruptions the columns are submerged by several centimetres of water with a temperature of up to 47’C. Eruptions last about 45 min. For about 15 min between eruptions there is very little water flow and during this period the stromatolites, which occur near the margins of outflow channels, are exposed to the atmosphere, and the temperature of any water flowing around the columns drops markedly, commonly to 20-25°C. Similar conditions prevail in the channel near “Weeds Pool”.

In 1971-1972 the conical stromatolites were abundant in ten springs and geysers. In nine, silica is precipitating ftom the thermal waters; in the tenth (Main Spring, Jupiter Terrace, Mammoth area), calcium carbonate is precipitating in the form of aragonite (Friedman, 1970). The various waters range in temperature from 32’ to 50°C and in pH from 7 to 9. As most previous studies of stromatolites have attempted to relate their morphology to environmental influences, especially to directional water currents, particular attention was paid to these factors in Yellowstone. The waters in which conical stromatolites are growing range from intermittently flowing in. outflow channels, through gently and irregularly flowing sheets of water in shallow pools, to nearly still in pools up to 30 cm deep. Measured flow rates varied considerably. In the outflow channels of “Column Spouter” the cones are submerged in several centimetres of water flowing at 5-20 cm/sec at peak flow, but are exposed to the atmosphere between eruptions. In “Conophyton Pool” the cones are continually submerged in 2-10cm of water in which current directions and velocities are variable, depending on wind direction and intensity as well as on the rate of outflow from the spring (which dis- charges continually). In this pool there are large differences in flow rates between the water surface and the bottom: on one occasion flow at the surface was 3-6cm/sec, but one centimetre beneath the surface no flow was detected. Flow rates of 1-10 cm/sec were measured at the tips of 2-3 mm high cones on a windy day. Flow around the cones in White Crater Spring was measured only once and was about 1 cm/sec. On a windy day, flow in “Weeds Pool” was 5-6 cm/sec at the surface and 0.5-1 cm/sec around the cones 5-10 cm below the surface; this pool is 30 cm deep, and is the deepest pool known to contain conical stromatolites.

Microbiology

The microbiology of the flat-topped columnar stromatolites is very poorly known because before they could be studied in detail slight changes in outflow patterns destroyed them all. They were constructed by a cyanophyte which in one sample was identified as Phormidium truncutum var. thermule Copeland, accompanied by the photosynthetic filamentous bacterium Chloroflexus uuruntiucus Pierson and Castenholz. The unicellular cyano-


55-

60 - 46- - s Y

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I ih icknrr

5.. * . .. . r. . . . %

D 129 sets of measurements

lWl40 IW18017Ol90 190200 220 140 u m l

129 sets of measurements C

l?6

20

a !

I 5

10

a

0

MORPHOGENESIS 299

phytes Synechococcus liuidus Copeland and S. mineruae Copeland were also present. Other filamentous cyanophytes were sporadic.

The conical stromatolites were studied in ten hot springs and geysers. In each case the microbiota was essentially the same: the dominant organism was the filamentous cyanophyte Phormidium tenue var. granuliferum Cope- land. Next most abundant was the bacterium Chloroflexus aurantiacus. The unicellular cyanophyte Synechococcus Naegeli was usually present, with S. liuidus much more abundant than S. mineruae. Other filamentous cyanophytes were sporadic; these included forms resembling Pseudanabaena and Isocystis. Organisms resembling myxobacteria were present in some samples and Spirillum-like bacteria and spirochetes occurred rarely. The microbiota of the calcareous cones from the Mammoth area was somewhat different from that of the siliceous forms. The Phormidium species present was slightly coarser (trichomes 1.0-1.7 pm wide) than P. tenue var. granuliferum (trichomes 0.8-1.2 pm wide) though in other ways it was indistinguishable.

As is shown in Table I, P. tenue var. granuliferum was usually much more abundant in the cones than in the mats between the cones. Conversely Chloroflexus was less abundant in the cones than in the mats. P. tenue var. granuliferum was most abundant in the tiny tufts and in the smallest, most recently developed cones. The ridges and spines frequently present on the cones had a microbiota that was indistinguishable from that of the cones generally.

To further our analysis of the morphogenesis of the conical stromatolites we determined the relative photosynthetic activities of natural populations of Phormidium and Chloroflexus in the stromatolites from Conophyton Pool; these are the two most abundant organisms in the cones. To do this, light- stimulated 14C02 fixation (photosynthesis) was measured under the prevailing environmental conditions (see Methods section). The results are presented in Table 11. These data show that Phormidium was responsible for about 90% of the total 14C02 fixation in the tufts; the remainder (DCMU-insensitive photosynthesis) was attributable to Chloroflexus. Incubation under a Wratten filter which excluded light of wavelengths suitable for algal photosynthesis resulted in an inhibition of algal photosynthesis in the tufts similar to that observed in the presence of DCMU. A statistical analysis of forty samples that included both tufts and the surrounding mat showed that the mean chlorophyll a content of the tufts, when standardized per unit protein, was about 1.5 times that of the mat from which they arise. This is consistent with the microscopic observation of more dead cells in the mat than in the tufts.

It is apparent that Phormidium is primarily responsible for the construc-

Fig. 32. Graphic plots of lamina thicknesses in Conophyton weedii. For the contouring of diagram D, the number of points within squares with 5mm sides were counted (in this reproduction the squares have 4.0-mm sides). This is the standard manner of data presentation for Conophyton stromatolites.


TABLE I

The relative proportions of various microorganisms in the conical stromatolites from two springs, expressed as percentages of the total microbiota (see Methods, section 5 )

P. C. S.1. S.m. “P.” “I.” m. “S.” s.

“Conophytoa Pool”

Tufts 82.9 Mat between 64.0

Cone tips 72.1 Mat between 59.7

tufts

cones

“Column Spo u ter ”

Small cones 83.7 Mat between 85.6

Large cones 54.1 Mat between 23.2

small cones

large cones

8.2 25.2

16.3 28.7

13.0 12.7

30.9 29.8

2.3 6.5

4 .O 8.2

3.3 1.8

8.3 24.6

0.3 5.4 0.6 0.2 3.1 0.5

0.4 6.4 0.4 1.7 1.1 0.5

- 0.4 - 1.4 - 0.5

0.4 - 0.3 -

-

-

0.5 - 0.2 -

- -

5.8 0.6 -

20.2 - 0.3

P. = Phormidium tenue var. ganuliferum ; C. = Chloroflexus aurantiacus; S.1. = Synecho- coccus liuidus; S.m. = Synechococcus mineruae; “P.” = Pseudanabaena-like cyanophyte; “I. ” = Isocystis-like cyanophyte; m. = myxobacteria (?); “S. ” = Spirillum-like bacteria; s. = spirochetes.

tion of the cones. Chloroflexus may play some role by utilizing the excretory products (e.g. mucus) of Phormidium as a carbon source (Bauld and Brock, 1974), thereby affecting the physical properties of the mats and cones (e.g. by reducing the cohesiveness of the aggregates of filaments). However, we have demonstrated in the laboratory that Chloroflexus-free cultures of P. tenue var. granuliferum are capable of independently forming conical structures (Figs. 33’34); these grow rapidly under ideal laboratory conditions. Conversely, natural Chloroflexus mats are frequently nodular but are never conical (Doemel and Brock, 1974).

Controls of stromatolite morphogenesis and lamination production

Biological influence Many species of Phormidium are present in Yellowstone waters (Copeland,

1936) but apparently only P. tenue var. granuliferum constructs conical stromatolites of the kind described here. The flat-topped columnar stromatolites seem to be built by a different species, P. truncatum var. thermale (this correlation of cyanophyte morphotypes with stromatolite morphotypes has a validity independent of that of the cyanophyte taxonomy; Copeland’s (1936) taxonomy and nomenclature are used for convenience, with the

MORPHOGENESIS 301

TABLE I1

Photosynthetic 14C02 fixation in conical tufts from “Conophyton Pool” (see Methods, section 6)

Component cPm/Pg % total Relative Chl a fixation abundance (%)

1. Phormidium -I- Chloroflexus 9419 100 100 2. Chloroflexus (4- DCMU) 681 7 9 2a. Chloroflexus (88A filter)* (1295) (13) 9 3. Phormidium** 8738 93 91

It is assumed that Phormidium and Chloroflexus account for all of the fixation of I4CO2. The amount of fixation is expressed as counts per minute (cpm) standarized to the chlorophyll a (Chl a ) content of the samples. The relative abundances of Phormidium and Chloroflexus in the samples are expressed as percentages, assuming that these two groups of microorganisms are the only components of the tufts (see Methods, section 5). Preven- tion of algal photosynthesis was by addition of DCMU (column 2) or incubation under a filter (column 2a). Each value for I4CO2 fixation is the average of four incubations.

*No dark controls. The percentage of 14C02 fixation given in parentheses was corrected only for absor tion. **Values for’ CO2 fixation by Phormidium were obtained by subtracting component 2 from component 1.

!

understanding that the named cyanophyte morphotypes may not represent true genetic taxa). These facts clearly suggest biological control of stromatolite morphogenesis, an interpretation that is supported by the growth of the conical stromatolites in a variety of hydrological regimes (including a laboratory flask).

Effects of light Light may be expected to play a key role in stromatolite accretion since

these structures are constructed by photosynthetic microorganisms. Previous studies of marine stromatolites (Monty, 1967; Gebelein, 1969) have shown that the orientation of cyanophyte filaments within stromatolites can be controlled by light and that lamination can be produced by the diurnal variation in intensity of light. In these examples, laminae with vertical filaments form during the day whereas those with horizontal filaments form at night. Diurnal light variation is responsible for similar lamination in the high-temperature bacterial stromatolites in Yellowstone, but in this case the filaments migrate upward at night (Doemel and Brock, 1974). Two interrelated aspects of our study involved investigating the role of light in morphogenesis and the temporal significance, if any, of the lamination. I t was difficult to predict the effects of light because both bacteria and cyanophytes (which may respond differently) are abundant in the stromatolites studied and furthermore the filamentous bacterium present (Chloroflexus) is only facultatively photo- trophic (Pierson and Castenholz, 1974).

MORPHOGENESIS 303

To investigate stromatolite growth in the dark, a long-term experiment was conducted from August 1971 to April 1972 in which a large light-proof black plastic box was used to shade about a square metre of cones in “Cono- phyton Pool”. The box did not prevent water movement; however, deprivation of light did prevent algal and bacterial photosynthesis and photoheterotrophism. The darkened stromatolites and those outside the box were marked with carborundum powder. After seven months the uncovered columns had accreted 0.5-2.0 mm and the new deposit was finely laminated. The columns under the box had a very uneven and almost completely unlaminated coating up to 0.5 mm thick above the carborundum. These results demonstrate that the conical stromatolites will not accrete in the dark. In another experiment, run for three months, an infra-red filter was used to prevent photosynthesis by the cyanophytes but not by the bacteria. This also stopped the accretion of the stromatolites. It is apparent from these two studies that light of the wavelengths used for algal photosynthesis is necessary for accretion.

The intensity of incident sunlight is subject to two major variations, diurnal and seasonal, which may be reflected in patterns of stromatolite accretion. Like the diurnal laminae of some marine stromatolites (Monty, 1967; Gebelein, 1969) the laminae of the flat-topped columns in Yellowstone consist of alternations of vertical and prostrate filaments. The laminae composed of vertical filaments are pale and 30-150 pm thick, whereas those composed of prostrate filaments are dark and 20-60 pm thick. A fourday experiment at “Column Spouter” demonstrated that these laminae also form diurnally, with the filaments standing vertically during the day and lying prostrate at night. Other experiments in the same hot spring showed no accretion over the same period. A 12-month experiment in “Column Spouter” during which carborundum layers were emplaced on three occasions also indicated that the lamination is diurnal. However, laminae do not form every day. During summer about 70% of the days are represented by pairs of laminae (pale and dark), but very few laminae formed during winter. The following results were derived from a single specimen in which the three marked layers were visible. Between 2 August 1971 and 20 April 1972 only 27 laminae formed, mostly near the beginning and end of this period (i.e. not during winter). Winter is represented by a 1 mm thick, dark, coarsely and diffusely laminated layer. Between 20 April and 20 June 1972 (61 days), 45-49 laminae formed. Between 20 June and 21 July (31 days), 20 laminae formed. In other examples there was either no growth during winter or else Calothrh coriacea overgrew the columns (because of a temperature drop resulting from changing water-flow patterns). Monty (1967) has observed that lamination is not strictly diurnal in Bahamian stromatolites either.

Figs. 33,34. Conical structuresformed by a pure culture of Phormidium tenue var.gronuli- ferum. Figures show cones as they appeared twelve days after addition of liquid growth medium to a preformed algal mat (see Methods section).


Fig. 35. Submerged, inclined cones in White Crater Spring, Shoshone Geyser Basin. Scale divisions are 1 cm.

No clearly diurnal lamination was found in the cones. One example from “Column Spouter” shows possibly diurnal lamination, but was overgrown by Calothrix coriacea before the accretion experiment was completed. Three scales of lamination are usually discernible in the cones. Laminae about 1 pm thick are ubiquitous; these consist of single layers of silica-encrusted filaments and their thickness is determined by that of the filaments (0.7-1.5 pm). Up to six such laminae can form in one day. The other two scales of lamination may be distinct or diffuse and are not always present. Laminae 5-15 pm thick occur frequently. They are difficult to .count but those from marked stromatolites in “Conophyton Pool” may be crudely diurnal. Fre- quently there are more laminae than days elapsed. The thickest laminae can be termed macrolaminae. In one experiment 7-8 macrolaminae 35-60 pm thick formed in 8% months. Thus, formation of laminae in the cones is not clearly controlled by diurnal variations in light intensity.

It has been suggested that the inclination of some fossil stromatolite columns fesults from the phototropic (heliotropic) growth of the constructing organisms (Norderg, 1963; Vologdin, 1963). In Shark Bay, Australia, the inclination of stromatolite columns is due to tidal current action (Hoffman, Ch. 6.1) or other, unknown, factors (Playford and Cockbain, Ch. 8.2). In the environment of hot pools in which water movement is relatively gentle, heliotropism might be expected to be a significant process. In the relatively high

MORPHOGENESIS 305

latitude of Yellowstone Park (45’) heliotropism, if present, should be apparent. However, with only one exception, all columnar stromatolites observed there accrete more or less vertically. The exception occurs in White Crater Spring where the conical columns are inclined about 20” from the vertical in a direction of S 45’W (Fig. 35). Water flow in the pool was measured at 1 cm/sec and the columns were inclined in the same direction, suggesting that deviation from the vertical in this example may have resulted from directional water flow. Possibly the lightdiffusing effects of clouds, refraction at the air-water interface, and scattering within the water, combine to reduce any inclination due to heliotropic growth. Inclination of stromatolite columns due to heliotropic accretion has not yet been demonstrated in either ancient or Recent stromatolites.

On a smaller scale heliotropism may be significant. The vertical orientation of filaments in the flat-topped columns during daylight hours is best interpreted as a heliotropic response. In the cones the filaments predominantly are parallel to the column lamination, which is very steeply inclined (mostly 50-80’ below the horizontal), or even vertical. At cone apices the filaments frequently form vertical tufts. In the laboratory it was shown that the most abundant filamentous cyanophyte in the cones, Phormidium tenue var. granuliferum, is phototactic and will glide towards a light source at’up to 16 pm/min. Such phototaxis will explain the observed filament orientation as well as the gross form of the stromatolites (as is discussed below).

Effects o f water flow The maximum height (synoptic, or growth, relief) of both types of stro-

matolites is determined by the depth of water in which they form. None extend above maximum water levels. They are both taller and broader in deeper water. The maximum water depth in which the cones were found was 30 cm; in this pool the cones were about 3.5 cm wide and had a synoptic relief of about 10cm. Greater size would probably have been attained had not the temperature of this pool fluctuated markedly over long periods, periodically allowing invasion by the coarse, filamentous cyanophyte Calo- thrix, which stopped cone accretion. Presumably, the maximum size of the structures is also limited by their mechanical strength, which would vary according to the amount of silica precipitated within the pools.

The flat-topped columns are subcircular in transverse section (Fig. 1) despite growing in water with a strong unidirectional current flow. Similarly, the cones are almost invariably subcirculh to subquadrangular in transverse section irrespective of the environment in which they were growing (Figs. 8-16). The development of ridges and spines on the cones varies from pool to pool; the origin of this variation is unknown. Ridges are rarely developed to the extent that columns become obscured. In one part of “Conophyton Pool”, ridges were both parallel and perpendicular to water flow directions. In the one case where the cones are inclined (Fig. 35), the inclination is in the direction of water flow (i.e., downstream).


a a u -e-Q m A b a a ! m

1 2 3

6 5 4

Fig. 36. Proposed morphogenetic sequence of the Yellowstone conical stromatolites: 1 , A mat of horizontal filaments; 2, Development of a knot of filaments due to tangling during gliding; 3, A tuft of erect filaments forms over the knot (due to phototaxis), 4-5, Tuft enlarges, 6, Occasional lateral filament movement produces bridging laminae (some of which are suspended above the substrate). Silicification occurs continually during the growth process.

Suggested morphogenesis of the conical stromatolites

The following morphogenetic sequence (Fig. 36) is a provisional interpretation. Everywhere in this section the word “filaments” refers to P. tenue var. granuliferum.

(1) A flat mat forms by the random gliding of filaments over their substrate. An example of this is the horizontal filaments forming the flat laminae visible in Fig. 17. (2) Here and there the randomly gliding filaments become entangled,

forming knots 100 pm or so wide and high, which project above the surfaces of the mats (Figs. 6, 7). Tangling due to gliding is a well-known phenomenon (Castenholz, 1969). The tangled knots of filaments are visible in Figs. 17 and 18. (3) Filaments gliding over the mats may encounter the knots and be

deflected upwards over them, towards the light source (the sun). This re-

MORPHOGENESIS 307

orientation may initiate positive phototopotaxis and additional upward gliding may result. It is known that in at least some species of Phormidium there is no phototopotactic response when filaments are perpendicular to the light rays (as in a flat mat under water), but phototopotaxis does occur when filaments are in other orientations (Haupt, 1965). Self-shading within the knots of filaments may also initiate positive photophobotaxis. Thus tufts of erect filaments form above the knots (Figs. 17,18).

(4) As a result of steps 1-3, the substrate is now dotted with vertical protuberances. Thus, increasing numbers of randomly gliding filaments are diverted up the sides of the tufts, which therefore become enlarged. At this stage the tufts are about 1-2 mm high. Tuft formation as described in stages 1-4 occurred within a 3-month period (from August to November 1972) in a small area of “Conophyton Pool”.

( 5 ) There will be a strong tendency for the most actively gliding filaments to be concentrated in the tufts, because they are the most likely to encounter them. The most phototopotactic filaments would probably glide the farthest, to the tips of the tufts. Those that exhibit weaker phototopotaxis would be left at the base or on the flanks of the tufts. Thus the tufts become pointed, and maximum accretion is concentrated at the tips. In cone formation,. the process of phototopotaxis is augmented by cohesion among the mucus-coated filaments (the filaments cohere like the hairs of a wet brush).

(6) The cones accrete upwards at rates determined by the supply of nutrients and solar energy. Variations in these factors, particularly in light intensity, produce intermittent accretion, and therefore lamination.

(7) If the directions of gliding of the filaments in a flat mat with protruding knots is not entirely random, then more filaments may impinge on each knot from some directions than from others. Thus lateral ridges will form on the knots. Once a slight ridge has formed it will enlarge because it will deflect more and more filaments upwards. This explanation is consistent with the occurrence of ridges converging on small tufts (Figs. 6, 7) as well as on larger cones (Figs. 11-16).

(8) Local entanglement of filaments can occur on cone flanks (Figs. 30, 31), producing knots which may develop into spines.

(9) At cone tips, filaments tend to stand vertically, forming tufts or thickenings within the laminae (Figs. 28,29). Filaments accumulate fastest at cone tips. The tips, therefore, are less silicified than the rest of the cone and so mechanically weaker, rendering them more susceptible to structural defor- mation (Fig. 23).

The three essential processes in the morphogenetic sequence suggested here are gliding, phototaxis and cohesion, all of which demonstrably occur in filaments of P. tenue var. granuliferum.


DISCUSSION

The banded-streaky microstructure of the conical stromatolites is a result of parallel alignment of the Phormidium and Chloroflexus filaments, which in turn results from their phototaxis and cohesion. Lenticular and subspherical thickenings in the laminae on the cone flanks form by tangling of filaments during upward gliding. The apical complexities result from the rapid accretion and tendency for tuft formation in this region, again due to phototaxis. Thus the microstructure is essentially a biologically determined characteristic.

The morphogenetic analysis of both of the columnar stromatolites described here stresses the definitive role of responses to light by the constructing microorganisms. In environments with only gentle water movement, light responses can determine stromatolite morphology. This has two applications in the study of ancient stromatolites. Morphogenetic analysis of ancient stromatolites can provide physiological data on the constructing microorganisms; these data may be correlated with microfossil types in those Precambrian stromatolites where microfossils are preserved (e.g. Awramik, Ch. 6.3). Secondly, stromatolites which formed in quiet-water environments (e.g. lagoonal, deep subtidal) will in many cases have morphologies determined by the physiological characteristics of the constructing microorganisms. The morphology of such stromatolites can be expected to change in response to evolutionary changes in microbiotas, and the stromatolites therefore should be biostratigraphically useful.

Conical stromatolites (Conophy ton) are abundant in Precambrian rocks but are not certainly known from Phanerozoic rocks. They are presently forming in hot springs probably because these are sufficiently free of animal grazers to permit development. I t is likely that similar stromatolites will be found in other hot-spring areas. They can also be expected to occur in other environments relatively free of animal grazers. It is probable that photosynthetic, filamentous, phototactic organisms other than P. tenue var. grunuli- ferum will be found constructing similar stromatolites. The abundance and diversity of Conophy ton stromatolites through geologic time was probably directly related to the availability of quiet-water environments largely free of grazing animals. The advent of active grazers near the end of the Precambrian (Awramik, 1971a) may well explain the apparently complete elimination of marine subtidal Conophy ton stromatolites at that time.

CONCLUSIONS

In the hot springs of Yellowstone National Park two types of columnar stromatolites, conical and flat-topped, are currently forming. Here these are described (and in the Appendix are named Vucerrilla wulcotti gr. et f. nov.

MORPHOGENESIS 309

and Conophyton weedii f. nov.). Prior to this discovery Conophyton stromatolites were thought not to have formed since the Precambrian. Field observations and laboratory experiments strongly suggest that the morphology and microstructure of the Conophy ton stromatolites result from the phototaxis and cohesion of the principal constructing microorganism, Phormidium tenue var. granuliferum. Conophyton-like structures have been induced to form in the laboratory, using pure cultures of this cyanophyte. The results of our work indicate that morphogenetic analysis of at least some fossil stromatolites can provide physiological data on the constructing microorganisms.

APPENDIX (by M.R. Walter)

Stromatolite systematics

Vacerrilla gr. nov.

Type form : Vacerrilla walcotti f . nov. Name: From the Latin for small posts, referring to the appearance of the stromatolite columns. Characteristics : As for the type form. Content: V. walcotti only. Age : Recent

Vacerrilla walcotti f . nov. (Figs. 1-5)

Columnar stromatolites (partim) Walter et al. (1972, figs. 1, 2a, b )

Holotype: CPC 15460. Name : In honour of Dr C.D. Walcott, who recognised the analogy between Precambrian stromatolites and those in Yellowstone hot springs. Diagnosis: Stromatolite with subcylindrical, unbranched, patchily walled columns, convex laminae and predominantly a banded microstructure except in the walls, where lamination is indistinct or absent. Description: See p. 278. Comparisons: In gross shape V. walcotti resembles forms of Colonnella Komar. It is distinguished from this group by its distinctive differentiated microstructure which changes from predominantly distinctly banded within columns to diffuse and frequently unlaminated on the margins. This same microstructural feature and its small size distinguishes it from Linocollenia angarica Korolyuk. Age : Recent.

Conophyton Maslov

Type form : Conophyton lituum Maslov. Diagnosis: Extremely rarely branching columnar stromatolites with conical laminae, many of which are thickened in their crestal parts. Content: See Walter (1972a, p. 102). Age: Precambrian, Recent.

Conophyton weedii f. nov.


Conophyton f. nov., Walter, Bauld and Brock, 1972

Holotype: CPC 15468. Name: In honour of W.H. Weed, a pioneer of American geology, who described in great detail the sinters of Yellowstone National Park. Diagnosis : Conophyton with a diffusely banded and streaky microstructure in which laminae predominantly are less than lOpm thick. Description: See p. 281. Comparisons : This stromatolite belongs to the group Conophyton Maslov (emend. Komar, Raaben and Semikhatov) because of its erect, parallel, non-branching columns built of conical laminae with upwardly directed apices and thickenings in the crestal zone. Micro- structure is used to divide the group into three subgroups (coarsely banded, striated and clotted). The microstructure of C. weedii cannot be classified into any of these types, although it is similar to the striated microstructure of C. garganicum Korolyuk (emend. Komar, Raaben and Semikhatov). In addition, the modal thickness of the laminae is in the 5-10pm range, less than half the thickness of the modes in previously described forms. It is these microstructural features that are used to define forms within the group Conophyton, and therefore a new form is here distinguished. To allow comparisons with fossil conophytons the data on lamina thickness are presented in the standard form (Fig. 32). Age: Recent. Depository : Type specimens of the stromatolites are deposited in the Commonwealth Palaeontological Collection (CPC), Bureau of Mineral Resources, Geology and Geophysics, Australia.

ACKNOWLEDGEMENTS

The U.S. National Park Service permitted field work in Yellowstone National Park. Financial support for Walter was provided by the Department of Geology and Geophysics, Yale University, and for Brock and Bauld by the U.S. National Science Foundation; Bauld also received an Australian- American Educational Foundation Travel Grant (Fulbright). J.H. Oehler critically read the manuscript and made many helpful suggestions.

6. MORPHOGENESIS

Chapter 6.3

GUNFLINT STROMATOLITES: MICROFOSSIL DISTRIBUTION IN RELATION TO STROMATOLITE MORPHOLOGY

S . M . Awramik

INTRODUCTION

Aside from the search for metazoans, Precambrian paleontology has suffered from a great dichotomy of research concentrations. One area of research has been the detailed study of microorganisms preserved in cherts while the other has been the study of stromatolites and their use in paleoecology and biostratigraphy. Very few investigators have concerned themselves with the possible relation between the microfossils contained within stromatolitic laminae and the stromatolites themselves. A major problem is that most stromatolites are preserved as carbonates containing little if any direct evidence of the microorganisms responsible for their accretion (see Gebelein, 1974, for an interesting discussion of microstructure in relation to major groups of blue-green algae). Vologdin (1955c, 1962) and Korde (1953b) have identified what they interpret as microorganisms preserved within the laminae of carbonate stromatolites from the Precambrian and Cambrian of the U.S.S.R. Serious doubt exists whether all the microstructures illustrated in their publications are the actual cellular remains of the algae or degraded organic matter diagenetically and mineralogically altered (Korolyuk, 1960c, 1963; Krylov, 1963; Semikhatov et al., 1963a; Kirichenko, 1964; Raaben and Komar, 1964; Walter, 1972a). By far the most compelling evidence for cellular remains preserved within stromatolitic laminae is in the microfossiliferous stromatolitic limestones and cherts of the Transvaal Dolomite (MacGregor et al., 1974; Nagy, 1974), Gunflint Iron Formation (Barghoorn and Tyler, 1965; Cloud, 1965; Licari and Cloud, 1068), Belcher Islands (Hofmann, 1974), Paradise Creek (Licari and Cloud, 1972), and Bitter Springs Formation (Schopf, 1968; Schopf and Black, 1971) to name a few (see Schopf, 1975, for a more detailed list). Not all microfossiliferous cherts are stromatolitic, those of the Fig Tree Group (Barghoom and Schopf, 1966) and Pokegama Quartzite (Licari and Cloud, 1972) being notable exceptions. Aside from Vologdin (1962), little attention has been paid to the possible relationship between preserved microorganisms

WALTER first proof, Chapter 6.3, page 1

312 S.M. AWRAMIK

and stromatolite morphology (Awramik, 1971b, 1973a; Licari and Cloud, 1972; see discussion in Walter, 1972a).

The stromatolites of the Gunflint Iron Formation with their well-preserved microflora are the source of valuable information for determining the possible relationships between microorganisms, the physical environment, and stromqtolite morphology.

I have observed and concluded the following about the Gunflint stromatolites: (1) they grew subtidally in shallow, frequently agitated, silica-rich water; (2) coccoid cyanophytes predominate over filamentous forms in stratiform stromatolites, whereas filamentous cyanophytes predominate over coccoids in columnar stromatolites; (3) a gradational change in the proportion of coccoids to filaments occurs in the change from stratiform to columnar stromatolites; (4) turbulence does not correlate with stromatolite morphology or contained organisms; and ( 5 ) each distinctive stromatolite morphology examined contains a distinctive assemblage of microorganisms peculiar to that morphology, reflecting, at least in part, environmental conditions and biological response to those conditions.

THE STROMATOLITES

The Middle Precambrian Gunflint Iron Formation trends ENE to WSW along the northern shore of Lake Superior, from Ontario into Minnesota (Fig. 1). The age of the Gunflint has been variously assessed at somewhere between 1,600 and 2,000 millions years although mounting evidence suggests an age around 2,000 million years (see Awramik, 197313, for a discussion on the age of the Gunflint).

Stromatolites are found along the entire strike of the formation, but are restricted to two stratigraphic levels: the Lower Cherty and Upper Cherty (Broderick, 1920; see also Hofman, 196913). The stromatolites in the eastern facies are generally black, including reduced iron compounds along with organic matter, while those from the western facies are red, with iron in the sesquioxide state (Barghoorn and Tyler, 1965). The microorganisms in the stromatolites of the eastern black cherts are preserved with greater fidelity than those from the western jasperoid facies. I have concentrated on stromatolites from the eastern facies.

Four morphologically distinct stromatolites were studied with a view to specifying their systematics and the types of microorganisms present. The stromatolite systematics have been reported elsewhere (Awramik, 197313) and here need only be reviewed briefly. Three types of columnar and one type of stratiform stromatolites were identified : Gruneria biwabikia schreiberi, Kussiella superiora, an unnamed form, and Stratifera biwabikensis (Figs. 2, 3, 4 and 5). Vertical transitions between the unnamed stromatolite, Stratifera, and Kussiella were found (Fig. 6). Horizontal transitions between different stromatolite forms were not observed.

MORPHOGENESIS 313

ONTARIO

Fig. 1. Map showing general areal distribution of .the Gunflint Iron Formation (dotted region). “Frustration Bay”, “Discovery Point” and “Flint Point” are informal locality names.

THE MICROORGANISMS

To date, 29 morphotypes have been described from the microfossiliferous cherts of the Gunflint Iron Formation (Korde, 1958; Barghoom and Tyler, 1965; Deflandre, 1968; Hofmann, 1971b; and Edhom, 1973). Of these 29 different morphotypes, 17 are of doubtful biological origin, doubtful taxonomic assignment and/or are morphologically indistinguishable from previously described morphotypes from the same rock unit (Awramik and Barghoom, 1975).

Continuing research on the microorganisms by Barghoom and myself has shed new light on the biological specificity and taxonomic position of many of the organisms previously described and presently under study. Briefly, we have found that:

(1) The microflora is wholly prokaryotic. (2) There are two different microfossiliferous chert facies. One, a thin-

bedded chert, contains what is interpreted as a planktonic assemblage of microorganisms dominated by Kakabekia urn bellata Barghoom, Eoastrion simplex Barghoom, and E. bifurcatum Barghoorn. The other is stromatolitic (Fig. 7). The microfloras found in these two facies are different, with very few common elements.

(3) Huroniospora Barghoorn may not be represented by three species based on surface sculpture as previously reported (Barghoom and Tyler, 1965). Surface sculpture may be a product of mineral precipitation on the surface of the cell while growing or after death, cellular degradation, or diagenesis. Size, although dependent on growth and reproductive stage, may be

314 S.M. AWRAMIK

Fig. 2. Polished slab of Gruneria biwabikia schreiberi. Bar equals 1 cm. Fig. 3. Polished slab of Kussiella superiora. Bar equals 1 cm. Fig. 4. Polished slab of unnamed stromatolite. Bar equals 1 cm. Fig. 5. Polished slab of Stratifera biwabikensis. Bar equals 1 cm. Fig. 6. Vertical transition; from the bottom: unnamed form + Stratifera + Kussiella. Bar equals 1 cm. Fig. 7. Thin-section photomicrograph of a microfossiliferous stromatolitic chert of the Gunflint Iron Formation. Bar equals 10pm.

MORPHOGENESIS 315

a more reliable criterion. Based on size, I found two populations of Huroni- ospora in the four stromatolites studied systematically; one smaller than 3 pm and one larger than 3 pm. Schopf (1975) reports the presence of three populations of Huroniospora in Gunflint cherts with the following diameters: 1-2 pm, 3-8pm, and 9-16pm. There was no clear-cut population of Huroniospora in the 9- 16-pm size range in the four stromatolites I studied.

(4) Six new morphotypes have been discovered bringing the diversity of the Gunflint to eighteen forms (Awramik and Barghoom, 1975). Only eight of the eighteen morphotypes recognized have probable taxonomic affinities with extant prokaryotes.

(5) Enterosphaeroides amplus Barghoom may not be a distinct taxon but a degradation product or empty cyanophycean sheath with spherical microorganisms occupying the microenvironment within the sheath (Awramik, Golubic and Barghoorn, unpublished).

The dominant microorganisms found in the stromatolitic cherts are the coccoid and filamentous blue-green algae Huroniospora spp. and Gunflintia minuta Barghoom, respectively (Plate I, 2, 4, 6, 7 and 9). The possibility that G. minuta is a photosynthetic flexibacterium has been suggested by Walter et al. (1972) but well-defined structures morphologically comparable to heterocysts and akinetes have been found (Licari and Cloud, 1968) making this suggestion, at present, untenable. All other microorganisms illustrated in Plate I, 1, 3,4, 8, lO and 11 are found within stromatolitic laminae but are very rare. The presence or absence of extremely rare forms is not considered critical to the interpretation of the possible relationship between stromatolite morphology and microbiota. Conceivably, the rare forms were a minor component of the microbiota contributing and depleting various resources, but present and preserved in such minor amounts as to not affect the stromatolite biocenose as a whole. In living algal mats it is quite common to find statistically rare microorganisms present with no apparent role in mat accretion living in an unoccupied niche or microenvironment as a “guest” (a term used by S. Golubic to describe this phenomenon).

BIOGENICITY OF THE GUNFLINT STROMATOLITES

Hofmann (1969b, and personal communication, 1974) raised an interesting and critical question: Are the microorganisms found within the stromatolitic laminae responsible for the biological growth of the stromatolite, or do they represent a detrital accumulation of microorganisms? Hofmann notes that the filaments appear as a “hash” with random attitudes and that Huroniospora is found solitary and not in clumps as would be expected with mats built by coccoid algae.

Although we may never be certain that the microfossils preserved within the laminae were responsible for the growth of the Gunflint stromatolites, or

316 S.M. AWRAMIK

MORPHOG ENESIS 317

any other microfossiliferous stromatolite for that matter, circumstantial evidence is consistent with a biogenic interpretation. By far the most compelling argument is that the filaments and Huroniosporu are restricted to the stromatolitic laminae while the microfossils characteristic of the non-stromatolitic facies are rarely found both within laminae and in the interspaces of stromatolites. Gunflintiu rninuta, the dominant morphotype, is occasionally found both prostrate and erect within laminae. I consider the Gunflint stromatolites biogenic.

METHODOLOGY

Thin sections for optical microscopy were prepared from each of the four stromatolites studied. The microscope was fitted with a Zeiss mechanical pointcounting stage set for 0.5 mm intervals for both the 3c- and y-axes. Thin sections were examined under oil immersion (ca. 1OOOx) and all the microorganisms within the field of vision (150pm in diameter) were counted for each of the 0.5-mm intervals. The entire thin section was studied in this manner. The numerical density of the microorganisms varied at each stop a;S did the preservation. Upwards of 50,000 microorganisms were counted in this manner.

The proportion of the different microorganisms encountered for each of the stromatolites examined (Table I) is presented in terms of the Braun-Blanquet (1932) scale of abundance as presented in Mueller-Dombois and Ellenberg (1974): 5 = > 75%; 4 = 50-75%; 3 = 25-50%; 2 = 5-25%; 1 = < 5%; + is rare.

PLATE I

Thin-section photomicrographs of microfossils from the stromatolitic cherts. All photomicrographs taken at the same magnification. Bar equals 10pm. 1. Eosphaera tyleri Barghoorn. 2. Gunflintia grandis Barghoorn. 3. Animikiea septata Barghoorn. 4. New form (Awramik and Barghoorn, in prep.). 5. Gunflintia minuta Barghoom. 6 . Huroniospora sp. 7 . ?Huroniospora. 8. New form (Awramik and Barghoorn, in prep.). 9. Huroniospora sp. 10. Archaeorestis n.sp. (Awramik and Barghoorn, in prep.). 11. New form (Awramik and Barghoorn, in prep.).

318

TABLE I

S.M. AWRAMIK

Microorganism composition of Gunflint stromatolites (numbers and + refer to Braun- Blanquet scale of abundance as presented in Mueller-Dombois and Ellenberg, 1974)

Gruneria biwabikia schreiberi Kussiella superiora Unnamed stromatolite Stmtifera biwabikensis

3 3 3 4

__ - -

4 4 4 3

__ - .~

+ 7 1 + - ~

T- 30 - 34% 42-46% 44 - 48% 65 - 68%

u)

Y

Eh

6 s B

a 0 E d

0 0

u

66-70% 54-58% 52-56% 32- 35%

DISCUSSION

The three columnar stromatolites studied contained a microbiota dominated by Gunflintia minuta while the stratiform stromatolites contained an assemblage dominated by Huroniospora spp. Turbulence has been considered the factor responsible for mat and stromatolite growth into columns and domes, without attendant microbiological changes, in the Recent algal mats and stromatolites in Hamelin Pool, Western Australia (Logan et al., 1974). In order to determine if turbulence was a factor in shaping the Gunflint stromatolites, I measured the apparent maximum diameter of oolites and clasts trapped within laminae and in interspaces (see Friedman, 1958) (Table 11). Stromatolites growing in turbulent waters would have the largest size fac- tion of particles associated with them. There was no relation between stratiform stromatolites and degree of turbulence; Stratifera grew in waters varying from quiet (as evidenced by the lack of oolites or clasts) to agitated and turbulent (as evidenced by the presence of abundant oolites within the laminae, some approaching 1.5 mm in apparent maximum diameter). There were no corresponding changes in the microbiota with differences in degrees of turbulence. In contrast, columnar stromatolites are invariably associated with oolites and clasts of varying diameter. Turbulence may be associated with columnar growth, but other environmental factors were probably responsible

MORPHOGENESIS 319

TABLE I1

Summary of Gunflint stromatolite data

Stromatolite Algae Range of Average max. diameter oolites and clasts max. apparent diameter

oolites and clasts

Gruneria biwabikia schreiberi coccoid 30- 34% 0.1-0.9mm 0.5 mm

K ussiella superiora coccoid 42-46% 0.3-11 mm 4.0 mm

Stmtifera biwabikensis mccoid 65-68% 0.0- 1.5 mm 0.8 mm

Unnamed stromatolite coccoid 44 -481 0.2-2.0mm 1.2mm

filamentous 66-70%

filamentous 5 4 - 5 8 1

filamentous 32-35%

filamentous 52 -56%

for the differences in the assemblages of the three columnar stromatolites studied. Unfortunately, these environmental factors appear to have left no detectable records.

The presence of a distinct assemblage within a given stromatolite morphology suggests (but does not prove) a biotic influence on stromatolite morphology. Similarly, in Proterozoic stromatolites from the Paradise Creek Formation of Queensland, Australia, Licari and Cloud (1972) reported the presence of one or two distinctive fossil morphotypes in each of the different stromatolite morphologies examined ; diversity was extremely low and preservation relatively poor.

In the four Gunflint stromatolites examined, a given stromatolite morphology does not contain microorganisms unique to that morphology (in contrast to the Paradise Creek stromatolites). The biotic changes in the Gunflint stromatolites are subtle and consist of changes in the relative abundance of microorganisms. Rare microorganisms are found too infrequently to give useful data at present. However, as the microorganism composition of a stromatolite changes, the morphology of that stromatolite also changes. This is clear in the vertical transition from the unnamed stromatolite to Stratifera and then to Kussiella through a distance of 4-5 cm (Fig. 5 ) . The biocoenose changes from one dominated by Gunflintia minuta in the unnamed columnar form to a biocoenose dominated by Huroniospora spp. in Stratifera. As Gun- flintia minuta becomes more abundant relative to Huroniospora spp., doming of the laminae becomes apparent. Finally, where filaments constitute approximately 55% of the microbiota, well-defined columns of Kussiella are present.


Four morphologically distinct stromatolites examined from the approximately 2,000 m.y. old Gunflint Iron Formation, north shore of Lake Superior, Canada, were found to contain distinctive microbiotic assemblages peculiar to the stromatolite morphologies. Assemblages are dominated by

320 S.M. AWRAMIK

coccoid and filamentous morphotypes of presumed blue-green algae. Colum- nar stromatolites contain an assemblage of preserved microorganisms dominated by Gunflintia rninuta while stratiform stromatolites contain assemblages dominated by Huroniospora spp. Changes in the dominance of one morphotype over another are accompanied by a corresponding change in the morphology of the stromatolite. No detectable paleoecological changes are associated with these macrostructural and microbiotic transformations. How- ever, biologically, one cannot separate ecological changes from biological changes affecting stromatolite morphology at a given instant in time. In Recent algal and bacterial mats, changes in temperature, availability of water, water chemistry and exposure to solar radiation effect changes in the microbiology (Golubic, 1973a; Logan et al., 1974). Such ecological changes would be undetectable in the fossil record. The relative degree of turbulence (which is paleoecologically detectable), had little or no observable effect on the morphology assumed by the Gunflint stromatolites or on the microorganisms responsible for accretion. Thus, subtle environmental changes are thought to have occurred while the stromatolites were growing in the Gun- flint sea. These environmental changes effected changes in the microbiota and, in turn, the stromatolite morphology changed. The evidence presented here and by Licari and Cloud (1972), coupled with the presence of time- restricted stromatolite forms in the Proterozoic (Serebryakov and Semikhatov, 1974), suggests that microbiota does indeed influence stromatolite morphology.

The Gunflint stromatolites and living pertinent examples provide data for a biologicalenvironmental approach to interpretation of stromatolite morphology. In the Gunflint, it appears that minor changes in the physical environment caused changes in the microbial assemblage and stromatolite morphology. Viewing the distribution of stromatolites in time and space during the Proterozoic, I consider that the evolution of new members or the extinction of existing members in cyanophytic and bacterial mat-building communities among taxa possibly as low as the specific and generic level, in interaction with the environment, may be responsible for the increase in diversity and the time-restriction of these stromatolite morphologies.

ACKNOWLEDGEMENTS

This paper is from my doctoral dissertation completed at Harvard Univer- sity, the Department of Geological Sciences. I thank E.S.Barghoom, P.E. Cloud, and H.J. Hofmann for critically reading the manuscript. D. Doerner (UCSB) prepared the plates. Financial assistance was generously supplied by the Department of Geological Sciences (Harvard), Committee on Evolutionary Biology (Harvard), a Grant-in-Aid of Research from Sigma Xi, and NSF Grant GA-37140 to E.S. Barghoorn.

6. MORPHOGENESIS

Chapter 6.4

BIOTIC AND ABIOTIC FACTORS CONTROLLING THE MORPHOLOGY OF RIPHEAN STROMATOLITES

S.N. Serebryakov

INTRODUCTION

There are two main concepts of morphogenesis of stromatolites. The first, which can be called ecological, is based principally on the evidence of Recent stromatolites; it was formulated most clearly by Logan et al. (1964). Accord- ing to this concept, the morphology of stromatolites depends entirely on the conditions of their formation. The second, the biotic concept, outlined mainly in the works concerned with ancient stromatolites by Fenton and Fenton (1933, etc.) and Maslov (193713, 1939a, etc.) postulates the existence of some relationship between the morphology of stromatolites and the taxa of algae to which they owe their formation.

The ecological concept of morphogenesis cannot provide a satisfactory explanation for the empirical observation of a temporal succession of Riphean stromatolites throughout the world or the presence of similar stromatolites in heterofacial coeval deposits (for details see Ch. 7.1, 7.2, 10.8). It is this evidence that indicates indirectly the existence of some relationship between the succession of Riphean stromatolites and the evolution of stromatolite-forming algae and, consequently, the existence of biotic control over the morphology of stromatolites. Observation of a correlation between the gross morphology of fossil stromatolites and the constructing algae, in the rare cases when it can be accomplished (see ch. 6.3), does not contradict this conclusion.

The mechanism of biotic control over the form of stromatolites is still uncertain. The taxonomic composition of the algal community is generally considered to be the determining factor on which the form of the stromatolite lamina depends (Korolyuk, 1963; Krylov 1963; Raaben, 1969a, b; Walter, 1972a; etc.). The dependence of the microrelief of the mats on the kinds of algae present was found to hold for Recent stromatolites (Logan et al., 1974). Raaben (1969b), Hofmann (1969a) and Walter (1972a) showed a correlation of some gross-morphological features of ancient stromatolites with the morphology of their laminae.

322 S.N. SEREBRYAKOV

The part played by algae in the morphogenesis of the Riphean stromatolites, however, cannot be restricted to determining the morphology of the laminae, especially since this itself may vary even within one column (e.g. Krylov, 1963) and can be identical in morphologically different stromatolites. In all probability, the morphology of the structures as a whole was biotically controlled. This is supported by the following. The structure of the ancient columnar stromatolites resulted from a long-continued successive accretion of laminae formed contemporaneously with the accumulation of the enclosing sediments. Presumably, there was some balance between the rate of stromatolite formation and the rate of sedimentation which could be maintained by interruptions to the growth of the structures and/or partial destruction of the laminae. The frequent occurrence of the latter process can be judged from the occurrence of fragments of stromatolite layers which frequently fill the intercolumnar spaces in bioherms. Intermittent growth reported by Monty (1967) for Recent algal mats seems to be typical for ancient stromatolites also. Of particular importance for us is that in spite of the highly probable irregularity of their growth and the periodic destruction of laminae in fossil stromatolites the structures formed were of a definite shape with the morphological features typical for the given taxon.

One possible mechanism of biotic control was suggested by Vlasov (1965, 1970). In his opinion the features typical for stromatolites are the “manifestation of the colonial life of mono-specific organisms”, the term “colony” implying not a simple multitude of individuals, but a “unification of individuals into a higher single whole” (Vlasov, 1970, p. 152). Various parts of such a colony were specialized to create definite intercorrelated elements of the structure. The colony itself underwent three stages of development (youth, maturity, old age) which resulted in the vertical morphological zoning of stromatolites. Unfortunately, this interesting hypothesis is based on numerous assumptions and practically ignores the role of environment in morphogenesis.

The biotic concept of morphogenesis of the Riphean stromatolites does not, in principle, exclude the direct effect of environmental factors on the morphology of the stromatolites (Fenton and Fenton, 1933; Fenton, 1943; Keller et al., 1960; etc.). But the successful stratigraphic application of stromatolites in the early 1960’s resulted in the underestimation of the role of these factors by some researchers, who held the view that the environments affected the form of the structures only indirectly, by determining the distribution of coenoses of algae (Krylov, 1963; Nuzhnov, 1967). This view, in its extreme, inevitably requires the correlation of any changes in the morphologiqal characteristics of stromatolites with changes in the taxonomic composition of the algae which formed them (Raaben, 196913).

Abundant evidence has been accumulated in recent years of the intrabiohermal morphological variability of stromatolites (Krylov, 1965, 1967a, 1972; Shapovalova, 1965, 1968; Shenfil, 196513; Vlasov, 1965, 1970;

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Bertrand, 1969; Bertrand-Sarfati, 1970, 1972b, c; Serebryakov, 1971, 1975; Serebryakov et al., 1972; Walter, 1972a; Preiss, 1973a; etc.). This variability has two principal features which suggest that it depended largely on environmental factors. Firstly, intrabiohermal variability is manifested in a similar way in taxonomically different stromatolites; secondly, it is not normally accompanied by changes in microstructures.

FACTORS INFLUENCING THE MORPHOLOGY OF RIPHEAN STROMATOLITES

The following is an attempt to consider, by giving some examples, the influence of biotic and environmental factors on the morphology of Riphean stromatolites and to determine the relative importance of these factors. The analysis is based mainly on the stromatolites of the Riphean deposits of various areas of Siberia.

The effect of environmental factors on morphology is quite obvious for many Recent stromatolites (see Ch. 6.1), and in some cases can be regarded as proved and in others as highly probable also for Riphean stromatolites. The simplest example of such an influence is the elongation of bioherms or some individual structures of various stromatolites. Hoffmann (1967) and Trompette (1969) showed that the long axes of these structures are oriented along the direction of paleocurrents reconstructed from crossbedding in associated sediments. The Siberian observations support this; there we have measured the predominant orientation of ripple marks and channels with carbonate flat-pebble conglomerates. The bioherms concerned have, in plan view, the form of biconvex lenses and they are arranged in a chess-board fashion in such a way that the distance between them is practically constant. The bioherms normally have a similar orientation over a considerable area; the orientation is not related to ephemeral changes in the direction of currents as shown by ripple-mark patterns. For example, the orientation of ripple-marks in dolomites of the Platonov suite (Turukhan region, the Sukhaya Tunguska river) varies even in adjacent thin (1-3 cm) layers. Mean- while the same orientation of the bioherms, elongated in plan view and enclosed within these dolomites (over a thickness of 10-15m), has been traced over almost 20 km. The elongation of individual columns within bio- hems is less constant and can successively follow several directions (see Serebryakov and Semikhatov, 1974, fig. 4).

The lamination form of stratiform stromatolites seems to be associated, in some cases, with water movement. The bedding relief of the upper surface of Stratifera sometimes cannot be distinguished from ripple crossbedding. In the opinion of Hofmann (1969a) this is also true of Gongylina.

Currents might be effective in causing asymmetry of structures as a result of a gradual displacement of their crests along the long bioherm axis (Hoffman, 1967; etc.). It is not improbable that this can be regarded as one

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S.N. SEREBRYAKOV

Fig. 1. A block-diagram of a bed including elongated Platella stromatolites from the Debengda suite (Middle Riphean) of the Olenek Uplift. These Platella form the upper part of a large “Jacutophyton cycle”; they grade into columns which are isometric in plan view. The front of the blockdiagram is a drawing made from a large-scale photograph.

of the causes of the formation of uniformly inclined columns within some bioherms. This mechanism, however, is by no means universal. Specifically, it cannot explain the inclination of the elongated Platella columns shown in Fig. 1.

One of the best examples of more complex effects of environmental factors on the morphology of structures is provided by the small bioherms of columnar stromatolites where both vertical and horizontal morphological zoning can be observed (Krylov, 1959a, 1965, 1967a; etc.). Horizontal zoning reflects lateral variations from the central to the peripheral parts of bioherms. In the marginal zones, inclined, horizontal and declined columns are frequently formed, while in a-parallel-branched unwalled stromatolites, the columnar structures may become stratiform here. The marginal zone is also characterised by morphologically indistinct columns which frequently have numerous bridges. Such changes seem to have been caused by the specific conditions in the marginal part of the growing bioherm (such as increased mobility of the water and the ability of the mat to extend laterally).

The height of the bioherm during its lifetime was also of great importance. This depended, during early stages of bioherm formation, on the relief of the substrate, and subsequently on the relative rates of growth of the stromatolite and accumulation of the sediments. I t could be comparable with, or greatly exceed, the height of individual columns during their lifetime. In the first case the columns were situated subvertically (Fig. 2A); in the second they were arranged like a fan. The height of the bioherm sometimes changed repeatedly during its formation as a result of change in the type of sediments or as a result of irregularity in the rates of sedimentation or the growth of

MORPHOGENESIS 325

B

Fig. 2. Drawings from photographs of two neighbouring bioherms with Kussiella kussiensis in the lower subsuite of the Kotuykan suite (Lower Riphean) of the Anabar massif. A. A bioherm in a lens of carbonate endoclastic and phytoclastic rocks. Stromatolites began growing on the concave section of the solid substrate. The height of the bioherm during its lifetime did not significantly exceed the height of individual columns during their lifetime. The columns are approximately vertical. B. A bioherm which formed on an erosional mound. In the marginal zone (on the right) the columnar structures become stratiform.


Fig. 3 . A section of the marginal zone of a large bioherm with Baicaliu. Various relations between the enclosing limestones and the bioherm reflect changes in its height during its lifetime. The orientation of the columns changes accordingly. Schematic drawing, Neruen suite (Middle Riphean) of the Uchur-Maya region.

the stromatolites. These changes, which are manifested in the relation of the enclosing stratified rocks with the bioherms, coincide with changes in the orientation of the columns in the peripheral zones of the bioherms (Fig. 3).

The change from columnar to stratiform stromatolites in the marginal parts of bioherms is also related to the height of the bioherm. The division of a wide structure (or the stratiform base of a bioherm) into narrower columns, typical of the a-parallel-branched unwalled stromatolites, seems to have been facilitated by the accumulation of mechanical sediments in hollows in the mats. In the marginal parts of relatively high bioherms, where the inclination of the mats reached some critical value (in the studied examples, about 50’) this sediment is likely to roll down the slope, so allowing the formation of extensive layers (Fig. 2B).

Finally, the height of bioherms during their lifetime determined, to some extent, their general form. Bioherms that rose markedly above the substrate normally have distinct outlines and can be easily distinguished from the enclosing rocks. Bioherms whose height during their lifetime was small normally have no clear lateral boundary and their outlines are irregular.

The vertical zoning of bioherms can, in its simplest form, result from the fact that the form of structures in their lower and upper parts is less distinct and less “morphologically defined” than the form of the stromatolite in the central part. The study of the composition, texture and structure of the enclosing and intrabiohermal rocks showed that this phenomenon is linked, to some extent, with changes in the environmental conditions occurring

MORPHOGENESIS 327

during bioherm formation. These changes were, in the majority of cases, of transgressive or transgressive -regressive character. The bioherms normally overlie erosion surfaces (as a rule subaqueous) or may be separated from them by clastic and microphytolitic rocks. The lower and middle parts of bioherms are normally enriched in coarse carbonate fragments, their size gradually decreasing up the section. There are also variations in the composition of fragments; in the lower parts fragments of the substrate predominate, while higher, fragments of stromatolites are most abundant.

The existence of directed changes of the environment are most pronounced where the stromatolites are members of small-scale sedimentary rhythms. Such stromatolite-bearing rhythms found in rocks of various ages have a rather uniform structure (Sarin, 1962; Korolyuk, 1966; Szulczewski, 1968; Serebryakov, 1971, 1975; Serebryakov and Semikhatov, 1974; and other authors). The rhythms are composed of various clastic (terrigenous and carbonate) rocks, phytogenic and chemogenic carbonates, and occasionally of argillites. Whatever the type (transgressive, transgressive-regressive or regressive) or the age of rhythms, stromatolites and microphytolites take a definite place in them, lying between the clastic rocks on the one side and the chemically precipitated ones on the other (Serebryakov, 1971, 1975; Serebryakov and Semikhatov, 1974). Columnar stromatolites often occur in a lower position in the transgressive rhythms than stratiform stromatolites (Fig. 4, see also Ch. 10.8, fig. 3). This suggests that the latter have been formed under calmer hydrodynamic conditions. Yet stratiform structures are sometimes present both at the top and the base of a stromatolite member of the rhythm. One has to admit that in this case there were two different ranges of conditions favouring the formation of stratiform stromatolites. A similar phenomenon has been described by Gebelein (1969) for the Recent sublittoral Castle Roads stromatolites (Bermuda).

When studying the vertical structure of bioherms one can often observe some insignificant changes in the morphology manifested in a similar way and simultaneously in a group of neighbouring columns or even throughout the whole bioherm. The Siberian observations provide the following examples of this type of change: (1) a sharp increase in the convexity of laminae without the appearance of any axial zone in Baicalia, Inzeria, etc.; (2) appearance of axial zones without any significant increase in the convexity of laminae in Kussiella, Baicalia and Inzeria; ( 3 ) change of column orientation from subvertical to inclined and then subvertical again at particular levels within bioherms of Kussiella, Baicalia, Platella and Colonnella; (4) Sharp changes (by 2-10 times) in the size of the columns of Colonnella, Anabaria, etc.; (5) temporary narrowing of columns of various stromatolites (Fig. 5); (6) simultaneous branching of structures, e.g. Kussiella - some examples of this type have been described in the literature (Krylov, 1967a; Raaben, 196913; Walter, 1972a; etc.). These changes in morphology are in some cases accompanied by changes in the enclosing rocks (e.g. enrichment in


- -

5

Fig. 4. Sedimentary rhythms of the upper subsuite of the Yusmastakh suite (Upper Riphean) of the Anabar massif. Legend: I = chemically precipitated dolomites; 2 = stratiform stromatolites; 3 = columnar stromatolites (Kotuikania); 4 = oncolitic dolomites; 5 = elastic dolomites; 6 = clayey dolomites; 7 = silty dolomites; 8 = sandy dolomites; 9 = sandstones; 10 = submarine erosional surface.

argillaceous or coarser elastic material), but more often they are apparently not reflected in the character of the abiogenic sediments.

Changes in the form of structures similar to those described above frequently occur many times in the vertical section of the bioherm, producing a cyclic structure. Obviously, they are the consequence not of a unidirectional “development of the bioherm” (Krylov, 1959a, 1965,1967a; Vlasov, 1965, 1970), but of some periodic changes in the environment which are often

MORPHOGENESIS 329

Fig. 5. Section of a bioherm with Kussiella kussiensis, lower subsuite of the Kotuykan suite (Lower Riphean) of the Anabar massif. One can see a sharp decrease in the diameter of the four left columns which coincides with the appearance of very thin clayey layers in the enclosing dolomites. The column on the right had a greater height during its lifetime and did not experience any changes. Evidently, this is also the reason why the decrease in the diameter in the two left columns occurs a little later than in the two central ones.

reflected only in the character of the stromatolites. This phenomenon may be called stromatolitic cyclicity.

A more complex variety of stromatolitic cyclicity is the repeated alternation within a bioherm of extensive horizons composed of different morphological groups of stromatolites. For instance, in one of the bioherms of the Burovaya suite (Turukhan region) there is an alternation of horizons of Baicalia, Ornachtenia and stromatolites resembling Kussiella (see fig. 5 of Serebryakov and Semikhatov, 1974). In the bioherms of the lower members of the Kotuykan suite (Anabar massif), levels with Kussiella, Colonnella and Stratifera frequently alternate with specific columnar stromatolites which somewhat resemble Baicalia. The different strbctures are in both cases linked by gradual transitions and share a similar microstructure.

A good example of stromatolitic cyclicity is provided by the so-called “Jacutophyton cycles” (Serebryakov et al., 1972), which are the rhythmic alternation of Conophy ton with various branching structures, observed by many researchers (Rezak, 1957; Votakh and Chayka, 1962; Shapovalova, 1965, 1968; Bertrand, 1969; Hoffman, 1972; Serebryakov e t al., 1972). “Jacutophyton cycles” are most typical of the Middle Riphean. Each


Fig. 6. A section of a transitional zone between the horizons of Baicalia (at the bottom) and Jacutophyton (at the top). The zone of lamina-shape change in two neighbouring columns differs in height by 7-8 cm. The tablet with the figure is 6 x 3 cm. The Derevnya suite (Middle Riphean) of the Turukhan region.

bioherm representing such a cycle comprises several horizons (tens of centimetres to several metres each) of Conophyton, Jacutophyton, Baicalia, and, rarely, Colonnella and Strutifera. At the boundaries between these horizons there are thin (5-20 cm) transitional zones within which the morphology of the stromatolites changes rapidly (Fig. 6). The transition from Conophyton to Baicalia always occurs via a Jacutophyton stage. The Buicalia horizon is normally confined to the top of the complete cycles. It is covered by a con- formable bed of non-stromatolitic rocks, separating the bioherm from the overlying one and producing a general rhythmicity in the section (alternation of stromatolitic and non-stromatolitic carbonates, Fig. 7). There are no regularities in the alternation of horizons other than those mentioned, but the order of this alternation is specific for each of the successive “Jucutophyton cycles”.

Within tectonically and facially homogeneous zones, the individual features of the cycles persist over some area in spite of local changes in the structure of individual bioherms (Bertrand-Sarfati, 1972c; Serebryakov et al., 1972). Thus, application of only morphological features of stromatolites and their order of succession in cycles allowed precise correlation of four lithologically homogeneous sections of the Neruen suite (Uchur-Maya region);

MORPHOGENESIS 331

Fig. 7. Correlation of sections of the middle subsuite of the Neruen suite (Middle Riphean) of Uchur-Maya region according to “Jacutophy ton cycles”. (Amended from Serebryakov et al., 1972.) Legend: 1-3 = Stromatolite horizons with: I = Conophyton; 2 = Jacutophyton; 3 = Baicalia; 4 = non-stromatolitic limestones; 5 = argillities of the lower subsuite; 6-7 = boundaries of: 6 = cycles; 7 = cycle elements; 8 = boundaries of horizons with stromatolites differing in morphology. The two right columns are each composed of two parallel close sections; each of them gives an idea of local changes in the bioherm structure.

the sections were 2-50 km apart (Fig. 7). This shows that periodically operating factors causing cyclic changes may be expressed unambiguously over an extensive area, controlling the appearance in the section of structures of one shape or another. In all the four sections shown in Fig. 7 the horizons containing Buiculiu account for 58-60% of the total thickness of the stromatolitic horizons.

The analysed features of the “Jucutophyton cycles”, and especially the rhythmic structure and persistence over some area, strongly suggest abiogenic causes for stromatolitic cyclicity. The question which is more difficult to answer is which ecological factors are responsible for these phenomena.

The changes in morphology of stromatolites making up the “Jucutophy- ton cycles” normally are not accompanied by macro- or microscopic changes in the enclosing rocks; frequently stromatolites of the same morphology do cross the boundary between two lithologically different horizons (Fig. 8;


Fig. 8. Contrasting levels of changes in the morphology of stromatolites and lithology of the rocks in the Derevnya suite (Middle Riphean) of the Turukhan region. Legend: 1 = level of changes in the morphology of stromatolites within the “Jacutophy- ton cycles”; 2-3 = strornatolite horizons with: 2 =Baicolia; 3 = Jacutophyton; 4 = surface of erosion; 5 = non-stromatolitic limestones; 6-9 = character of rocks: 6 = light chocolate porcelanous crypto-to microgranular limestones and dolomitic limestones; 7 = reddish-brown irregularly grained crypto- to fine-grained limestones and dolomitic limestones; 8 = brownish-grey (on weathered surfaces they are brick red) fine-grained calcareous dolomites and dolomitic limestones; 9 = grey irregularly grained, crypto- to fine- grained limestones and dolomitic limestones.

MORPHOGENESIS 333

Serebryakov et al., 1972). The dependence of the morphology of stromatolites on factors of the environment which are not expressed by features of the sediment has been noted by Maslov (1959) and Preiss (1973a).

We are often unable to determine conclusively which environmental factors affected particular stromatolites, neither can we record all their possible combinations. For instance, according to the African evidence (Bertrand, 1969), and, to a lesser extent, the Siberian evidence (Serebryakov et al., 1972), Baicalia was found to be associated in the “Jacutophyton cycles” with the most rudaceous varieties of the carbonate rocks. This as well as the cases of replacement of Jucutophyton by Baicalia in the marginal zones of bioherms suggests that hydrodynamics affect the morphology of stromatolites. Yet observations in the Derevnya suite (Turukhan region) showed that the change of groups cannot be related exclusively to changes in water movement. Passing up-section in this suite, there is increasing evidence from one cycle to another of the turbulence of the basin water. At the same time, the character of cycles remains essentially the same (provided the absence of horizons with Conophyton in the uppermost cycles is disregarded).

The quest for some single cause responsible for the changes in the morphology of stromatolites often results in mutually exclusive conclusions. Both Rezak (1957), who considered alternations of Collenia and Conophy- ton, and Shapovalova (1965, 1968), who analysed the alternations of horizons of Baicalia, Jacutophyton and Conophyton, attributed the changes in morphology of the stromatolites to changes in water depth. However, Rezak believes that Collenia was intertidal and Conophyton subtidal, whereas Shapovalova is of the opinion that Conophyton formed under shallower conditions than Buicalia. At the same time, the data now available suggest that the effect of bathymetry on the morphology of the stromatolites in “Jacuto- phyton cycles’’ was insignificant. The occurrence of stromatolitic cyclicity cannot be attributed solely to depth fluctuations, particularly in cases of the lateral replacement of Conophyton by Baicalia in the marginal zones of bioherms (Krylov and Shapovalova, 1970a; Serebryakov et al., 1972).

In some other cases, however, it is the fluctuation of depth that seems to be the main cause responsible for the changes in the morphology of the stromatolites. For example, the replacement of columnar-layered stromatolites by stratiform stromatolites in the sedimentary rhythms of the Omakhta suite (see Ch. 10.8, fig. 3) is likely to have been associated with the gradual deepening of the basin. The hydrodynamics of the basin do not seem to have undergone significant changes: the size of detrital quartz grains is practically the same in all rock-types comprising the rhythm, including the stromatolite dolomites (Serebryakov, 1971).

From what has been said it follows that stromatolites have a considerable intrabiohermic morphological plasticity which can be attributed to the effect of environmental factors. The effects of these factors are frequently not reflected directly in the character of the enclosing sediment which makes


their simple interpretation difficult. The normally uniform microstructure of all morphological variety of stromatolites of each bioherm suggests that the effect of the environment upon the morphology of the structures was direct and did not depend on changes in the community of the stromatolite-forming algae. This conclusion rests on the assumption that the features of microstructure (except for those having an obviously superimposed character) reflect, at least to some extent, the primary textures and structures of the original stromatolite and are correlated in some definite way with the taxonomic composition of the algae which gave rise to it (see Ch. 7.1).

Consequently, that morphologic variability of stromatolites which is not associated with changes in the microstructure, does not seem to be dependent on changes in the taxonomic composition of the algal communities either. Thus, we have good reasons to suppose that stromatolites of various morphologies but identical microstructure formed as a result of the life activity of a single community or species of algae (Komar and Semikhatov, 1965; Komar, 1966; Krylov, 1967a, 1972; Serebryakov, 1971, 1975; Bertrand-Sarfati, 1972c; etc.).

However, some instances have been observed where the changes in morphology of stromatolites are accompanied by changes in the microstructure (Raaben, 196913; Serebryakov et al., 1972). In these cases, the microstructure normally does not change completely, but the relative abundances and shapes of its various elements alter. In such cases, those ecological factors which caused the gross morphologies to change can also be supposed to have caused changes in the algal communities (for example, a change in the dominant species); i.e., we are dealing here with “ecads” in the sense of Fenton and Fenton (1933) and Fenton (1943).

Though the form of stromatolites has been found to change‘under the influence of environments, it should be emphasised that the range of these changes is not infinite. It seems to be controlled by the taxonomic composition of the stromatolite-forming algae. As we have previously mentioned, the existence of such a control can be seen not only from the existence of the definite succession of similar assemblages of Riphean stromatolites in the deposits of various regions. I t can be also established by the direct study of the intrabioherm variability of the stromatolites.

Krylov (1967a) was the first to notice the regular combination of structures of different morphology in the bioherms of certain stromatolites. He coined the term “bioherm series” (Krylov, 1972, p. 56) to designate the combination of “the principal morphological varieties of structures belonging to the same bioherm or to uniform biohenns within a single bed”. Simi- lar bioherm series occur in coeval deposits over vast areas. Thus, intrabioherm variability does not prevent the use of stromatolites in stratigraphical zonation; on the contrary, it opens up new possibilities for their stratigraphic application (comparison of bioherm series). Similarly, the taxonomic composition of the sets of stromatolites which occur in the otherwise identical

MORPHOGENESIS 335

stromatolite rhythms developed at different stratigraphic levels proved to be different. These sets differ greatly from one another, particularly in the morphology of their columnar structures (Serebryakov, 1971, 1975).

A good example of the biotic control over the morphology of stromatolites is provided by the “Jacutophyton cycles”. The morphology of the branching members of such cycles is specific for the Lower, Middle and Upper Riphean (Serebryakov et al., 1972). In the Lower Riphean the cycles have KussieJa; in the Middle Riphean, Baicalia and Tungussia; in the Upper Riphean, Inzeria, and, according to Bertrand-Sarfati (1972c), Gymnosolen, Tilemsina and distinctive Baicalia.

Thus, in spite of the variability of the morphology under the influence of environments, we can observe that due to biotic control over the morphology of stromatolites each Riphean subdivision has specific, unrepeated assemblages of structures.

CONCLUSIONS

The morphology of Riphean stromatolites depends both on the taxonomic composition of the stromatolite-forming algae and on the environment. Macroscopic features of stromatolites can be subdivided into two groups. The features of one group (orientation of structures, elongated or isometric form and, possibly, the diameter of columns, etc.) were, evidently, fully controlled by abiotic factors. Others (style of branching, margin structure, shape of columns and laminae) are evidently of a dual nature, biotic and abiotic. In other words, they depended on the taxonomic composition of algae and could, at the same time, change within some definite range under the influence of the environment. Thus, in spite of the presence of inclined columns in the marginal zones of bioherms of Kussiella, Baicalia, Gymnosolen, etc., they can be easily distinguished from one another by characters of the second group. Though the transition from non-branching stromatolites with conical laminae to branching structures with convex laminae can be observed at various stratigraphic levels, the morphological features of the branching stromatolites (as well as the microstructures) change in time.

The consequence of the dual nature of the morphological features of stromatolites is, in all probability, the specific “convergence” of features; i.e. similar features might be the result of the operation of different factors. For instance, the appearance, in the Lower Riphean, of rare stromatolites with a smooth lateral surface and more or less distinct wall could be attributed to specific conditions of their formation, in particular to the fact that, in some localities, sedimentation might lag behind the rate of growth of the structure. On the other hand, the extensive development of such stromatolites in the Upper Riphean can be more reasonably attributed to the specific features of


the stromatolite-forming algae, because it is unlikely that the total rate of sedimentation was significantly reduced in the Upper Riphean.

However, the principal manifestation of the dual nature of the morphology of stromatolites is the existence of regular, definite sets of morphologically various stromatolites, fitting the diagnoses of various groups, supergroups and types in the current classification of stromatolites. The morphologically different members of each such set have a uniform microstructure and, evidently, were formed by the same community or the same species of algae. Included among the sets under consideration are Krylov’s (1972) bioherm series, which correspond to the sets of single bioherm or to several bioherms of a single layer.

The analysis of the taxonomic composition of the sets indicates that each community (or species) of algae has a most typical and most frequently encountered morphological manifestation of the stromatolites formed by it. Therefore, the relative abundances of different stromatolites within a bioherm are almost always very different. One of the morphological modifications is normally dominant and comprises the bulk of the stromatolite body. This modification apparently results from optimal conditions for the development of the particular algae.

In the sets under consideration there are, apart from taxa of extensive vertical distribution (e.g. Strutifera), stromatolites typical of limited stratigraphic intervals. I t is these that are responsible for the specific appearance of each of the sets and allow their stratigraphic application.

7 . STROMATOLITE BIOSTRATIGRAPHY

Chapter 7.1

EXPERIENCE IN STROMATOLITE STUDIES IN THE U.S.S.R.

M . A . Semikhatov

INTRODUCTION

Most of the modern concepts in stromatolitology originated in the U.S.S.R. in Maslov’s works. Considering stromatolites as organo-sedimentary structures, he came to the conclusion that there was a relationship between the species composition of the stromatolite-building algal mats and the gross morphology of the stromatolites; he made the first attempt to clarify the main trends in stromatolite evolution and successfully used them for the first time for inter-regional correlation (Maslov, 1937b, 1938, 1939a, b, 1950; etc.). During the post-war years Maslov (1945, 195313, 1959, 1960) began carefully appraising the possibility of stromatolites acting as time markers and demonstrated the variability in their morphology and microstructures under the effects of the environment (in the case of the Ordovician forms). The concept that stromatolite morphology is completely dependent upon the environment was predominant in the U.S.S.R. during the 50’s. The most diligent in developing this concept were Korde (1950,1953b) and Vologdin (1955b, 1962) who concentrated their attention on the search for cellular algal remains in the carbonate stromatolite structures.

EMPIRICAL EVIDENCE OF THE STRATIGRAPHIC SIGNIFICANCE OF STROM ATOLITES

A new phase in these studies, which has undoubtedly provided evidence of the stratigraphic significance of stromatolites, began in the U.S.S.R. in the late 50’s. This phase was started by the work of Korolyuk (1958, 195913, 1960). She distinguished three types of stromatolites (layered, nodular and columnar), described several form-genera and -species for each type, and demonstrated for the first time the diagnostic significance and temporal variation of the margin structure of columns. The method suggested by Krylov (1959b, 1963) for reconstructing the form of columnar structures by serial sectioning (“graphical reconstruction”) made it possible to identify two new

338 M.A. SEMIKHATOV

diagnostic features of columnar stromatolites: their branching style and the gross morphology of the columns. The application of this method in studying the Uralian stromatolites has shown that there is a distinct change in the morphological varieties of stromatolites through the geological column. The branching type, the margin structure and character of the microlamination change in a regular manner with time (Krylov, 1960b). Similar results were obtained with the Siberian material (Nuzhnov, 1960; Semikhatov, 1960). Somewhat earlier, it had been noticed that massive developments of conophytons were confined to the lower part of the Upper Precambrian (Riphean) in the U.S.S.R. (Dragunov, 1958; Keller and Khomentovskii, 1958, 1960) and specific branching structures were described for the Upper Riphean within the northeastern and eastern margins of the Russian Platform (Krylov, 1960c; Raaben, 1960).

The generalization of the information gathered by 1960 had shown that there are three stromatolite complexes in the key sections of the Riphean in Siberia and the Ural (Keller et al., 1960) and that the composition and succession of these complexes coincide over vast areas. This coincidence was considered to be evidence that the vertical change in the stromatolite complexes probably reflects the evolution of stromatolite-building algae and not changes in the local sedimentary environment. The first isotopic datings obtained by the K-Ar method confirmed this idea, and proved to be similar in the rocks containing similar stromatolites both in Siberia and the Ural.

The conclusions regarding the existence of time-dependent, consistent associations of Riphean stromatolites have been further developed in the works of a large group of geologists who investigated all the main platform and miogeosyncline sections in northern Eurasia (Korolyuk et al., 1962; Semikhatov, 1962; Korolyuk, 1963; Krylov, 1963, 1966c, 1967a, 196813, 1969, 1972; Semikhatov et al., 1963a, b, 1967a, 1970; Komar, 1964,1966; Komar et al., 1964, 1965a, 1970, 1973; Raaben, 1964a, b, 1967, 196913, 1972a, b; Furduy, 1965,1969; Keller et al., 1965; Kirichenko, 1965; Nuzhnov and Shapovalova, 1965, 1968; Semikhatov and Komar, 1965; Shenfil, 196513; Golovanov, 1966, 1970; Golovanov and Raaben, 1967; Nuzhnov, 1967; Krylov et al., 1968, 1971; Komar and Semikhatov, 196813; Korolyuk and Sidorov, 1969, 1971; Raaben and Zabrodin, 1969, 1972; Dolnik and Vorontsova, 1971, 1972; Rabotnov et al., 1971; Khomentovskii et al., 1972a; and others). New data obtained have confirmed the conclusion that the changes in the composition of the stromatolite complexes are irreversible in the Riphean section and make it possible to distinguish a fourth complex in addition to the above three. The fourth complex defines the uppermost strata of the Precambrian, designated the Vendian or Yudomian, on the basis of historical geological data (Sokolov, 1952,1964; Keller and Sokolov, 1962) and the assemblages of microphytolites (Zhuravleva and Komar, 1962; Zhuravleva, 1964a). It is important to emphasize that the levels at which the complexes change are largely independent of historical-geological boundaries.

STROMATOLITE BIOSTRATIGRAPHY 339

This is seen best of all in the Siberian sections (Semikhatov, 1962; Komar, 1966; Khomentovskii et al., 1972a).

The stromatolite classification that served as a basis for these biostratigraphic schemes was suggested by Krylov (1961, 1963). I t is a formal one and it provides for the specification of three categories of taxa - types, groups (form-genera) and forms (form-species) - and for Latin binary names for the groups and forms. Maslov’s suggestion (195313,1960) that three-fold or five-fold nomenclature should be applied to stromatolites has not been accepted, since such nomenclature is complicated, cannot replace descriptions and is no more informative than binomial nomenclature (Korolyuk, 1960c; Krylov, 1963). Columnar stromatolites are basic to the Subdivision of the Riphean. In Krylov’s classification they are grouped into form-genera according to a combination of three features: the gross morphology, branching style and margin structure. I t has been suggested that forms should be distinguished by details of column morphology. However, microstructure has become a basic factor in their practical identification (e.g., Semikhatov, 1962; Komar et al., 1965a; Komar, 1966). Raaben (1964b) and Komar (1966) suggested slightly different schemes based on Krylov’s classification (see Ch. 2.4 and Hofmann, 1969a). Analyses of the diagnostic stromatolite features were published by Krylov (1963), Komar (1966), Raaben (1965, 1969a, b) and Raaben and Zabrodin (1972).

The stromatolite distribution in the Riphean reference section of northern Eurasia is shown in Fig. 1 constructed on the basis of data available from the Geological Institute of the U.S.S.R. Academy of Sciences and critical analysis of published information. With a minor exception, endemic forms are not plotted in Fig. 1. It should be emphasized that there are different concepts of a number of groups (form-genera) in the current literature (see Ch. 2.4). In compiling Fig. 1 the author kept as close as possible to the original concepts of the taxa. The published data on the vertical distribution of the most important groups of Proterozoic stromatolites are generalized in Fig. 2. Substantiation of the correlations implied in the construction of Figs. 1 and 2 is given in Geochronology of the U.S.S.R. (Anonymous, 1973b) and in Semikhatov (1974). Four successive stromatolite complexes and adequate microphytolite associ-

ations (Zhuravleva, 1962, 1964a, 1968) served as a basis for constructing the general four-fold scale for the Riphean on paleontological grounds (Keller, 1964, 1966a; Komar et al., 1964; Semikhatov, 1966). This scale was derived from an earlier stratigraphic scheme which had been based on historical-geological criteria and used the unconformity-bounded lithostratigraphic rock groups of the Ural and Russian Platform as the types for chronostratigraphic units (Keller, 1952; Sokolov, 1952). The persistence of this thinking has caused a duality in concepts of the Riphean subdivision. On the one hand, the paleontological criterion is now regarded as fundamental in their distinction, and these subdivisions are accordingly understood to be

340 M.A. SEMIKHATOV

Fig. 1. Distribution of stromatolites in the Riphean reference sections in northern Eurasia. Legend: 1 = carbonate and terrigenous-carbonate formations; 2 = terrigenous formations; 3 = the pre-Riphean formations; 4 = discontinuities; 5 = acid intrusions; 6 = basic intrusions; 7-12 = stratigraphic boundaries: 7 = the Cambrian lower boundary; 8 = the Terminal Riphean lower boundary; 9 = the Upper Riphean lower boundary; 10 = the Middle Riphean lower boundary; 1 1 = the Lower Riphean lower boundary; 12 = bound- ariesof formations; 13 = unconformity; 14-24 = isotopic age in m.y.; 14 = K-Ar method (glauconite); 15 = K-Ar method (hydromica); 16 = K-Ar method (whole-rock, effusive rocks); 17 = K-Ar method (minerals of metamorphic and intrusive rocks); 18 = Rb-Sr method (whole-rock, granitoids); 19 = Rb-Sr isochron method (whole-rock, granitoids); 20 = U-Pb method (zircon); 21 = isochron U-Pb method (uranium from cement of clastic rocks); 22 = U-Pb isochron method zircons 23 = &Pb method (clastic zircons); h~ = 0.557 lO-''year-'; h ~ b = 1.39 * 10- year . An.r. = Anabaria radialis Komar; B. = Baicalia sp., B.b. = B. baicalica (Maslov), B.in. = B. ingilensis Nuzhnov, B.1. = B. lacera Semikhatov, B.m. = B. minuta Komar, B.mc. = B. maica Nuzhnov, B.p. = B. prima Semikhatov, B.r. = B. rara Semikhatov, B.un. = B. unca Semikhatov; Bx.ab. = Boxonia allahjunica Komar and Semikhatov, Bx.g. = B. grumulosa Komar, Bx.1. = B. lissa Komar; C1. = Colleniella sp., C1.s. = C. singularis Komar; Col. = Colonnella sp., C01.c. = C. cormosa Komar, Co1.d. = C. discreta Komar, Co1.f. = C. frequens (Fenton), Co1.k. = C. kylachii Shapovalova, Col.1. = C. lineata Komar, Col.ul. = C. ulakia Komar; Con. = Conophyton sp., C0n.c. = C. cylindricum Maslov, C0n.g. = C. garganicurn Korolyuk, C0n.m. = C. miloradovici Raaben, Con.ml. = C. metula Kirichenko, Con.1. = C. lituum Maslov; G. = Gymnosolen sp., G.ab. = G. altus Semikhatov, G.as. = G. asymmetricus Raaben, G.f. = G. furcatus Komar, G.r. = G. ramsayi Steinmann, G.t. = G. tungusicus, G.l. = G . levis Krylov, Gn.d. = Gongylina differenciata Komar, Gn.m. = G . mixto Komar, Gn.n. = G . nodulosa Komar and Semikhatov, Gn.z. = G. zonata Komar; In. = Inzeria sp., 1n.c. = I. confragosa (Semikhatov), 1n.d. = I. djejimi Raaben, 1n.n. = I. nimbifera (Semikhatov), 1n.t. = I . tjomusi Krylov, 1n.vr. = I . variusata Golovanov; Jac. = Jacutophyton sp., Jac.m. = J. multiforme Shapovalova, Jac.r. = J. ramosum Shapovalova; Jur.c. = Jurusania cylindrica Krylov, Jur.j. = J. judomica Komar and Semikhatov, Jur.s. = J. sibirica (Yakovlev); Jur.t. = J. tumuldurica Krylov; K. = Kussiella sp., K.k. = K. kussiensis (Maslov), K.vt. = K . uittata Komar; Kt.t. = Kotuikania torulosa Komar;L.uk. = Linella ukka Krylov, L.s. = L. simica Krylov; Min.s. = Minjaria sakharica Komar, Min.ur. = M. uralica Krylov. M1.m. = Malginella malgica Komar and Semikhatov; M1.z. = M. zipandica Komar; N.in. = Nucleella inconformis Komar, N.f. = N. figurata Komar, Om. = Omachtenia sp., Om.om. = 0. ornachtensis Nuzhnov; Pan. = Paniscollenia sp., Pan.em. = P. emergens Komar; P1.p. = Platella protensa Komar; Prm.aim. = Parmites aimicus (Nuzhnov), Prm.c. = P. concrescens Raaben; Svsv . = Svetiella svetlica Shapovalova; St. = Stratifera sp., St.f. = S. flexurata Komar, St.ir. = S. irregularis Komar, St.un. = S. undata Komar; T. = Tungussia sp., T.b. = 2'. bassa Krylov, T.cl. = T. colcimi Raaben, T.cn. = T. confusa Semikhatov, T.n. = T. nodosa Semikhatov; Tn.p. = Tinnia patomica Dolnik. I-IV = formations; I = Kartochka; I1 = Krasnya Gora; III = Ostrov; I V = Migoedikha; V = Nemakit-Daldyn.

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348 M.A. SEMIKHATOV

beds with a definite stromatolite assemblage. On the other hand, there is still an endeavor to determine the time-range of these units by the time-ranges of those lithostratigraphic rock groups which had been suggested as the types of the Upper Precambrian subdivisions. The situation is complicated by the fact that (a) not all the members of the stromatolite succession described in Siberia have been identified in the Uralian section historically chosen as the type for the Riphean and its subdivisions; and (b) there is some uncertainty in the precise correlation of this section with the Siberian ones. The above facts explain the difference in viewpoints on the subdivision of the Riphean and on the position of the boundaries of its subdivisions (discussions can be found in the works of Keller and Semikhatov, 1968, pp. 26-28, 90-92; Keller, 1971; and Semikhatov, 1974, pp. 260-264). , In spite of these differences, the four-fold stratigraphic scale of the Upper

Precambrian based on stromatolite and microphytolite assemblages is more widely used in the U.S.S.R. than the other scales. Since 1961 it has served as a basis for all the correlation schemes approved at conferences held on Riphean stratigraphy in the main regions of its development in the U.S.S.R. and is used in a number of important generalizations (e.g., Anonymous, 1963, 1973b; Garris et al., 1964; Vinogradov, 1968; Keller et al., 1968). The four main subdivisions of the Upper Precambrian in this scale are considered to be chronostratigraphic. Following Keller (1966a), a number of geologists consider them to be units of a special rank: protosystems or phytems. The lower three phytems are usually described as the Lower, Middle and Upper Riphean, while the fourth one (Vendian or Yudomian) is either included in the Riphean as its terminal member (e.g. Keller and Semikhatov, 1968), or distinguished as an independent unit, sometimes referred to the Phanerozoic (e.g. Sokolov, 1968, 1 5 7 2 ~ ) . The Yudomian is considered in this chapter as the Terminal Riphean.

The data on the isotopic geochronology of the Late Precambrian of the U.S.S.R. obtained mainly by the K-Ar method, in general confirmed the stratigraphic conclusions based on stromatolites and made it possible t o date the boundaries of the complexes at 1350 f 50,950 f 50 (1000) and 675 k 25 (680 k 20) m.y., respectively (Garris et al., 1964; Keller, 1964; Semikhatov and Chumakov, 1968; Anonymous, 197313).

Pre-Riphean stromatolites are of a very limited distribution in the U.S.S.R. and are poorly studied in comparison to those of the Riphean (Metzger, 1924; Butin, 1960, 1965, 1966; Krylov, 1966a; Vologdin, 1966; Kirichenko et al., 1967; Lyubtsov, 1973; Dolnik and Nikolskii, 1974). That is why the information now available about them is mainly based on the data obtained in other countries. However, the application of comparative methods is only beginning. Nevertheless, we can state even now that in the Aphebian as well as specific taxa there exist widespread groups with long time ranges extending through the Riphean and even into overlying formations.


DETAILED FOUR-FOLD SCALE OF RIPHEAN SUBDIVISION ACCORDING TO STROMATOLITES

The very considerable time-range of the four main subdivisions of the Riphean based on the assemblages of groups of stromatolites, explains the striving towards a detailed general scale. Three approaches to the use of stromatolites for this purpose may be outlined. In the first approach subdivisions are distinguished in one (Kabankov et al., 1967; Golovanov, 1970) or in several neighbouring sections (Nuzhnov and Shapovalova, 1965,1968) and are characterized by one or two forms. Thus, a relatively detailed subdivision is reached, but the problem of the lateral persistence of the subdivisions and of the time-ranges of the forms has still to be solved. It is not clear therefore whether the peculiarities of the distribution of forms are purely accidental and dependent on local conditions.

The second approach is more promising. It has led to the present scheme of subdividing the Middle and Upper Riphean and made it possible to outline the two-fold subdivision of the Yudomian in Siberia. It provides for selecting certain assemblages of the stromatolite groups and forms on the basis of empirical data on their distribution in a series of sections in two or more regions. For stratigraphic purposes, only an integral picture of vertical distribution of the taxa is used, which summarizes the data available from the largest possible area. This makes it possible to balance, to some extent, the particular features of stromatolite distribution associated with environmental influences and to recognize the time-ranges of the taxa. The fossiliferous strata are grouped then with overlying and/or underlying formations into stratigraphic subdivisions so that the upper boundary of each subdivision coincides with the lower boundary of the next (Semikhatov et al., 1967a, 1970; Krylov et al., 1968; Raaben, 1969b, 1972; Raaben and Zabrodin, 1969; Komar et al., 1970; Krylov and Shapovalova, 1970b). As the stromatolite- bearing strata usually form only a certain part of the section, peculiarities of the strato-type (such as the position of unconformities, sharp changes in lithology, etc.) are considered to be of decisive importance in determining the location of these boundaries. It is clear that tracing of such boundaries between regions is a very complicated matter and sometimes leads to contradictions.

The third approach to Riphean stratigraphy lies in a detailed interregional correlation on the basis of similar stromatolite forms discovered (actually, of similar microstructures) (Komar and Semikhatov, l965,1968a, b; Semikhatov and Komar, 1965; Komar, 1966, 1973). Such an approach has two weak points. Firstly, microstructure becomes the basic diagnostic stromatolite feature according to this viewpoint; this leads to a violation of the principles of distinguishing taxa in the present classification (see below). Secondly, the tedious task of determining the distribution of stromatolites with certain microstructures in each of the regions must precede their use in the

3 50 M.A. SEMIKHATOV

inter-regional correlation. The fact is that the persistence of the succession of microstructures observed in a single section over a whole region is only one of the possible ways they can be distributed (see Ch. 10.8), as was revealed by the detailed investigations carried out in the Uchur-Maya Region (Komar et al., 1973). Moreover, some of the forms with a peculiar microstructure in one area may be considered as age datum markers in one succession but those in another area may be encountered in the reverse succession (e.g., Inzeriu tjornusi and I . confrugosu in the Uchur-Maya and Turukhan areas, Fig. 1 and Ch. 10.8, fig. 5 ) .

The most reliable results are obtained through the second of the approaches under consideration. This approach was used in establishing the phytems and now is applied in distinguishing their subordinate subdivisions, usually called “horizons”. They are established on stable assemblages of forms of different groups and sometimes even by specific groups. Three horizons are usually distinguished in the Middle Riphean, viz. the Svetlinian, Tsipandian and Lakhandian horizons (Krylov et al., 1968, 1971; Komar et al., 1970; Krylov and Shapovalova, 1970b), and two in the Upper Riphean, viz. the Katavian and Minyarian (Krylov and Shapovalova, 1970a; Krylov, 1972), or, to some extent of different content, the Biryanian and Minyarian (Raaben, 1969a, 1972). The other two approaches in detailing Riphean stratigraphy using stromatolites are aimed at detailed subdivision of the horizons. Their subordinate units are distinguished with confidence only on the regional scale at the present stage.

All the Middle and Upper Riphean horizons, except the Svetlinian, extend over a very large area in northern Eurasia, and some of them can be traced beyond there (Krylov, 1972; Semikhatov, 1974). Nevertheless, they cannot be recognized in all the stromatolite-bearing sections in the U.S.S.R. since the specific stromatolite composition makes it possible, in a number of cases, to distinguish only a phytem as a whole. Moreover, Gymnosolen and Minjuriu urulicu, typical of the Minyarian horizon, occur in some of the sections above the beds with Inzeriu tjornusi and Jurusuniu cylindricu (the Katavian horizon), while in the other continuous sections the Katavian stromatolites are missing and the Upper Riphean begins with the beds containing Minjuriu urulicu and/ or Gymnosolen (see Fig. 1). This fact is probably explained by environmental causes as the first association of forms tends to occur in grey-coloured massive dolomites and the second one in yellow and red argillaceous dolomites and limestones. Some other stromatolites are also confined to certain facies (Semikhatov et al., 1970). However, there is also a contrary trend, i.e. spread of particular forms and certain associations of taxa into different facies. This trend, particularly in the Katavian association, is stressed by its discovery in clean, dark limestones on the Korean Peninsula and in the Uchur-Maya area (see Ch. 10.8). Preiss (1973a) arrived at similar conclusions for a number of taxa from the Adelaidean of South Australia.


CRITERIA FOR ESTABLISHING STRATIGRAPHIC BOUNDARIES ACCORDING TO STROM ATOLITES

The investigations carried out in recent years have considerably widened knowledge of the systematic composition of the successive complexes of Riphean stromatolites, and have shown that the complexes have more groups in common than considered previously (see Fig. 2). As a result, establishing criteria for determining the stratigraphic boundaries according to stromatolites has become urgent.

The majority of the Soviet stromatolite workers consider that the boundaries should be determined by the change of stromatolite complexes, i.e. by the change of the empirically established stable assemblages of certain groups and forms. Thus, the boundary of the Middle and Upper Riphean is defined, on the one hand, by the disappearance of the complexes of Lakhandian baicalias, conophytons, jacutophytons, tungussias and associated groups and, on the other hand, by the spread of another complex with the predominance of certain forms of gymnosolens, minjarias, inzerias, and jurusanias. The older groups are greatly reduced in the new complex and, as a rule, represented by new forms. This boundary is clearly traced in Siberia and, beyond its territory, established in the northern part of the Korean Peninsula and in some other areas (Semikhatov, 1974). However, the exact position of the boundary is not quite clear in the Ural. The Katavian suite contains the first Upper Riphean stromatolites and the underlying stromatolite-bearing (the Tsipandian according to Krylov, or Svetlinian, according to Komar) part of the Middle Riphean deposits are separated here by the thick clastic Zilmerdak suite and its antecedent unconformity to which are confined gabbro-diabases with K-Ar age of 1,015-1,195 m.y. (Fig. 1). On the basis of historical- geological and geochronologic data this boundary in the Ural is placed either at the base of the Zilmerdak suite (Raaben, 1964a, 1967, 1969a, 1972b; Krylov and Shapovalova, 1970a, b; Krylov, 1972) or at its top (Anonymous, 1973b; Semikhatov, 1973, 1974). If we accept the second alternative, which seems to be more consistent with the data now available (discussions are given in Geochronology of the U.S.S.R., Anonymous, 1973b, pp. 317-319), we are inevitably led to contradictions regarding the position of the Middle and Upper Riphean boundary. In the Ural this boundary is traditionally placed at the unconformity at the base of the Zilmerdak suite (age about 1,100 m.y.), but in Siberia the most probable time-equivalent of this suite is the lower part of the Lakhanda Group (Khomentovskii et al., 1972a; Semikhatov, 1974), which contains a stromatolite complex of the same name usually included in the Middle Ripheak

Another criterion for placing boundaries - the first appearance of new groups - is adopted by Khomentovskii et al. (1972a, b). According to this, the Middle/Upper Riphean boundary is lowered by them down to the base of the Lakhandian horizon or to the underlaying Tsipandian horizon, as it is

352 M.A. SEMIKHATOV

in these horizons that the first rare minjarias, gymnosolens and inzerias are encountered among the baicalias, jacutophytons and conophytons typical for this part of the Riphean (Fig. 1). Such a lowering of the Middle and Upper Riphean boundary in Siberia has been supported by Keller (1973a).

The following argument can be used against this viewpoint. Firstly, the Lakhandian And Tsipandian minjarias, gymnosolens and inzerias are known only in two areas of Siberia, are encountered within the Middle Riphean complex of stromatolites, and are represented as a rule by endemic forms having a microstructure common with baicalias dominant in the same beds ‘and bioherm. Secondly, adherence to this viewpoint may prevent subdivision of the Precambrian using stromatolites since there are stromatolites in the Aphebian of Canada, Karelia and Australia (together with stromatolites unknown in the Riphean) which are close if not identical to Gymnosolen, Minjaria, Katavia and Tungussia by a number of criteria (Krylov, 1966a; Hofmann, 1969a, b; Walter, 1972a).

The idea of lowering the Middle and Upper Riphean boundary in Siberia was supported by Komar (1973) from a different point of view. Defining the boundary not on the change of stromatolite complexes, but as the boundary of the Yurmatian and Karatavian Series of the Uralian section, he suggested a new correlation of the latter section with the Siberian sections. In this connection, two aspects must be emphasized. First, this correlation is based on comparison of stromatolite microstructures essentially without taking into consideration the morphology of the structures and sometimes in contradiction of it. Considering the microstructures as the most important diagnostic feature, Komar in a number of cases uses for his correlation microstructures belonging to different morphological groups of stromatolites (e.g., Baicalia and Svetliella, Inzeria and Appia) or those supposed to be at the transition between stromatolitic and abiogenic structures (Malginella malgica). Secondly, the correlation is contradicted by: (1) results of tracing the Kata- vian stromatolitic association (Krylov and Shapovalova, 1970b; Krylov, 1972); (2) some of the geochronologic data (in particular, the dating of the pre-Zilmerdak events in the Ural and sediments of the Turukhan section: Ivanovskaya et al., 1973; see Fig. l), and (3) that the certain correlatives of the upper three suites in the Upper Riphean of Turukhan, referred by Komar to the upper part of this phytem, are intruded by granites on the Yenisey Ridge which have U-Pb and Rb-Sr ages of 850 k 50m.y. (Semikhatov et al., 1973).

The “level of development of characters” for the columnar structures, such as the occurrence of active branching in the Middle Riphean and of walls in the Upper or Middle Riphean, was considered several years ago to be a valid criterion for establishing stratigraphic boundaries according to stromatolites (e.g., Komar et al., 1964; Krylov, 1966a; Raaben, 1969b). This scheme constructed on the example of Riphean stromatolites in the U.S.S.R. has not been confirmed by the studies of pre-Riphean forms; some of these


proved to possess active branching and/or walls (Cloud and Semikhatov, 1969b; Hofmann, 1969a, b; Walter, 1972a).

However, it should be stated that the data now available have not dis- proved the conclusions on the qualitative differences in Riphean stromatolitic complexes of different ages, though they have made us revise the simplified concepts about these differences. So, for example, the Middle Riphean complex differs from its precursor not only in the unique combination of groups and forms of stromatolites but also in the domination of actively branching unwalled structures. The latter, though they are encountered also in older and younger sequences and do not completely characterize this complex, nowhere else reach such an abundance and variety. In the same way, the presence of actively branching and/or walled forms in the Aphebian cannot disprove the fact that such stromatolites attain abundance and, to a considerable extent, replace other groups in the Late Riphean.

ENVIRONMENTAL VARIABILITY IN STROMATOLITE MORPHOLOGY

During the period of confirming the stratigraphic significance of stromatolites (1959-1964) attention was concentrated on the problem of determining the parallelism in the changes of complexes in remote sections. The main objects of investigations were, therefore, the columnar varieties which are encountered more often than others and which possess the clearest diagnostic features. The variability of stromatolites under the effect of the environment, which had been recorded by Fenton and Fenton (1933), Cloud (1942), Fenton (1943), Maslov (1959,1960) and other investigators, was beyond the attention of the new generation of stromatolite workers during these years. The successful use of stromatolites in stratigraphy corroborated the presence of biotic control of their morphology. All that resulted in the concept that ecological factors could affect the morphology of the structures only in an indirect way, defining the location of algal coenoses (Krylov, 1963; Nuzhnov, 1967; Raaben, 1969a), and that the stromatolites with stable diagnostic features were widely developed in the Riphean and Cambrian (Korolyuk, 1963; Krylov, 1963; Korolyuk and Sidorov, 1965; Keller and Semikhatov, 1968).

The morphologic variability of stromatolites attracted attention again after Vlasov (1965, 1970) and Krylov (1965,1967a) began detailed studies of the structure of bioherms. According to Vlasov and Krylov, the morphologic features of stromatolites change in a regular manner from the centres of bioherms to their margins, and there is a definite vertical and less apparent horizontal zonation in completely developed bioherms reflecting certain stages of their development. Conclusions on the universality of such regularities were criticized (Raaben, 1969b), but the fact of the variability of stromatolite morphology within a bioherm has received new corroboration (see Ch. 6.4).

354 M.A. SEMIKHATOV

The growth stages of bioherms were attributed by Vlasov (1970) to biological regularities in the development of algal colonies. However, the ecological interpretation of the stromatolite variability in bioherms has become more popular. It has been consolidated by the correlation of changes in stromatolite morphology with changes in the composition of the enclosing rocks which reflect fluctuations of the environment (Miroshnikov, 1965; Shapoval- ova, 1965, 1968; Shenfil, 196513; Serebryakov, 1971; Serebryakov et al., 1972). However, it still remains to be determined what the environmental factors are which have created different changes in the morphology. In addition, there is a definite contradiction with the examples of the morphological variability of stromatolites which are not accompanied by changes in the composition, texture and structure of enclosing rocks, and numerous examples of morphologically identical stromatolites encountered in deposits of different facies (see Ch. 6.4).

The results of studies of stromatolite microstructure were of great importance in the interpretation of their morphological variability. The microstructures were sometimes identified as algae (Krasnopeeva, 1946; Korde, 1950, 195313; Vologdin, 195513,1962). However, the majority of Russian stromatolite workers consider them to be the products of the interaction of various carbonate-depositing processes (including biogenic processes) and subsequent modification of sediments; they regard them as one of the diagnostic features of stromatolites (usually as the main distinguishing feature of forms). Norm- ally it is considered that the peculiarities of the microstructures are, in a certain way, connected with the systematic composition of stromatolite- building algae. This viewpoint, formulated by Walcott, the Fentons, Maslov and Korolyuk, was strengthened especially when the usual consistency of microstructures in bioherms, the stratigraphic limitation of a number of microstructures, and the increase in their variety from the Lower Riphean to the Cambrian were demonstrated (Komar, 1964, 1966; Komar and Semik- hatov, 1965, 1968b; Semikhatov and Komar, 1965; Krylov, 1969; Raaben, 196913; Golovanov, 1970; Korolyuk and Sidorov, 1971). The dependence of stromatolite microstructures on the composition of the algae which formed them was strongly supported by studies of Recent stromatolites (Monty, 1967; Gebelein, 1969; Hoffman et al., 1969; Logan et al., 1974).

The weak point in using microstructure as a diagnostic feature is that we can distinguish only some of their secondary transformations, namely, those which arise at relatively late stages (Kopeliovich and Krylov, 1960; Krylov, 1963; Komar, 1966; Raaben, 1969a; Semikhatov et al., 1970). In each particular case we cannot determine the significance of diagenesis in the alteration of the primary stromatolite microstructures. That is why all the stromatolite microstructures except those of definitely secondary origin are considered to be of diagnostic significance. Such an assumption is corroborated by the frequent recurrence of identical microstructures in rocks of different composition (including terrigenous) in various areas, and also by their survival


in the rocks found at different stages of alteration (up to regional metamorphism at lower greenschist facies).

All the evidence taken together strongly suggests that the consistency of microstructure within bioherms reflects the unity of the algal communities which formed them. If so, examples of the persistence of microstructures through changes in gross morphology show (in full compliance with data on Recent phytogenic structures) the effect of environmental factors on the morphology of stromatolites (Shapovalova, 1965, 1968; Shenfil, 1965b; Komar, 1966; Krylov, 1967a, 1972; Komar and Semikhatov, 1968b; Sere- bryakov, 1971; Serebryakov et al., 1972; and others).

This conclusion, on the face of it, cannot be reconciled with conclusions on the stratigraphic significance of stromatolite morphological groups. How- ever, it has been found that the morphological variability resulting from changes in environmental conditions is not unlimited, and does have definite ranges probably determined by biotic factors. This was proved in three ways that are closely interconnected.

Firstly, in studying Riphean and Cambrian stromatolites, Krylov (1967a) came to the conclusion that the morphological varieties of certain stromatolites occur in regular combinations in bioherms; these same combinations are encountered in contemporaneous beds in numerous sections. Krylov (1972) later advanced the hypothesis of the bioherm series. These series are the combinations of morphological modifications present in a single bioherm or in several bioherms of the same kind in a single bed; each series may have been formed by one community (or species) of algae under various conditions. Each of the bioherm series as well as structures of a single characteristic group forming the central zone of the bioherm contains less-developed structures of some other groups forming the margins of the bioherm. Those groups can be common to two or more series (see Ch. 2.4).

Secondly, analysis of stromatolite cycles in stratigraphic sections (see Ch. 6.4 and 10.8) has shown that the cyclic alternation of groups depends upon the fluctuation of abiotic factors; but the systematic composition of the assemblages of groups present in such cycles, and their microstructures, are peculiar to certain subdivisions of the Riphean, and are therefore time- dependent (Serebryakov, 1971; Serebryakov et d . , 1972).

Finally, the existence of biotic control of the gross morphology of stromatolites was strongly supported by the intimate knowledge of the presence of successive complexes of stromatolites, generally of constant composition, on an inter-regional and even inter-continental scale.

The relationships between microstructure and gross morphology led several authors to the conclusion that microstructure must be taken into account not only in defining forms but also higher categories. Microstructural unity has been regarded as a decisive argument for referring all the morphologies found in a single bioherm to one and the same form (Komar, 1966; Semikhatov et al., 1967a, 1970; Komar and Semikhatov, 1968b). The morphology that is

356 M.A. SEMIKHATOV

observed in the central part of a bioherm where the stromatolites were less subjected to the effects of the environment was considered to be typical for the form, and the appearance of peripheral morphologies was regarded as ecological variability of the one form. Thus, the position of Fenton and Fenton ( 1 9 3 9 ~ ) and Fenton (1943) was actually repeated regarding the classification of ecologic varieties (ecads). In practice it leads to an understanding that the groups are not morphological categories but the sum of the variable morphological forms of a certain microstructure. Hence, the principle adopted in modem taxonomy for the determination of stromatolite groups, i.e. the use of a set of morphological features, is violated, and the taxa selected acquire new meaning, though they formally remain members of the old classification (see Ch. 2.4).

That is why it should be admitted after Krylov (1967a, 1972) that the various morphologies within bioherms and biostromes may and must have different group names in accordance with their morphological features and irrespective of the microstructure. Paleobotanical classification provides a precedent for such a procedure. This viewpoint is now approved by the majority of the Soviet stromatolite workers, including the author of this paper who previously took a different position.

Unfortunately, there are now only a few examples of descriptions of different stromatolite groups derived from single bioherms and possessing a single microstructure (Krylov, 1967a, 1975). It was suggested that all stromatolites with particular microstructures be considered and all the morphological modifications (groups) of these stromatolites represented in available material enumerated (Serebryakov et al., 1972; Komar et al., 1973; Semik- hatov, 1974). The absence of a widely adopted microstructural classification has not made it possible to use definite terms in this case and the investigators had to apply complicated definitions as “the microstructure described in distinguishing such and such species”. I t is clear that this solution cannot replace paleontological descriptions of stromatolites within the classification now in use. On the other hand, the relations considered urgently require that an independent stromatolite classification, based entirely on microstructures, be elaborated.

C 0 N C L U S I 0 N S

Thus, the understanding of Precambriarl stromatolites in the U.S.S.R. has developed from the assertion of their stratigraphic significance (Maslov) to its denial unde‘r the influence of ideas on the complete dependence of morphological features upon the abiotic factors, and then in a new phase to the corroboration of the stratigraphic significance of certain morphological groups on the basis of their empirically determined vertical distribution.

It is the empiric approach to the solution of the problems regarding


stromatolites which is a distinguishing feature of the Russian school of investigators. The analogies with Recent stromatolites, which dictated for a long period of time the genetic and ecological interpretation of stromatolites and determined a priori the negative estimation of their stratigraphic potentialities in the western literature, were always treated with great care in the Russian literature. Moreover, beginning from Maslov’s works there have developed in the Russian literature ideas on substantial differences in the ecology, mechanisms of formation and factors of morphogenesis of ancient and the most widely distributed Recent stromatolites. At present, these ideas are acquiring greater and greater popularity.

The influence of ecologic factors on the morphology of ancient stromatolites was either not accepted, or was denied, at the beginning of the new phase of studies in the U.S.S.R. The conclusion of the reality of such an influence - although its development was within certain limits probably determined by biological causes - was gradually accepted later.

At present, a four-fold scale for subdivision of the Riphean in northern Eurasia has been established on the basis of changes in the complexes of certain groups and forms of stromatolites (mainly columnar). Recently it has been indicated that these complexes may be applied to the distinction of similar subdivisions on the other continents, and therefore the Russian stratigraphic scale of the Riphean based on stromatolites can be duplicated elsewhere. Subordinate subdivisions of more or less regional significance have been recognized in the three upper phytems of the Riphean in a number of sections in Siberia and the Ural. Some of them can be traced far beyond northern Eurasia.


Chapter 7.2

INTERCONTINENTAL CORRELATIONS

W.V. Preiss

INTRODUCTION

Precambrian geologists have for many years entertained the possibility of using palaeontology for interregional and intercontinental correlations. Stromatolites are by far the most abundant Precambrian fossils, but their potential as index fossils was questioned as soon as their mode of origin became understood (e.g. Cloud, 1942). Logan et al. (1964) discovered a unique modern analogue of Precambrian columnar stromatolites at Shark Bay, Western Australia, but even here, there is nothing comparable to the enormous diversity of complexly branching forms known to be very abundant in the Precambrian.

Realization of the potential value of stromatolites as index fossils can only follow from detailed taxonomic investigations. As many features of stromatolites as possible must be described and ranges of variation within single occurrences noted, so that form taxa can be erected and discriminated by serial sectioning, reconstruction and thin-section techniques. In the classification of stromatolites, it is inadmissible to refer to “genera” and “species” since stromatolites are not fossil organisms. Nor are “form genus” and “form species” as used by palaeobotanists for parts of plants appropriate. The terms c‘grou”’ and “form” used by the Russians who revived interest in stromatolite taxonomy over fifteen years ago, have no strict biologic con- notations and are therefore most suitable for the classification of these biogenic sedimentary structures. Much of the argument that has ensued arises from the semantics of classification. The taxonomic approach has a different aim from the descriptive classifications such as that of Logan et al., (1964). We are not concerned here with the geometry of an individual lamina or column, but with an entire structure passing through a number of growth stages and having a definite range of variability, whose mode and total extent must be determined in each case. For example, a progression from laterally linked to columnar to laterally linked forms (LLH + SH + LLH in descriptive formulae) is almost universal in columnar stromatolites and hence is of little taxonomic interest. A complex of characters is thus required for the discrimination of the various taxa.

360 W.V. PREISS

Both approaches are in themselves equally valid, but the advantage of the taxonomic approach is that it leads to the definition of a number of discrete taxa whose time-ranges can be investigated empirically. If these individual time-ranges turn out to be short, then the value of stromatolites for Precam- brian biostratigraphy will have been demonstrated. If, on the other hand, the defined taxa span all or most of geological time, we will know with certainty the limitations of the method. Until this is achieved, there can be no agreement, for no amount of theoretical argument about modem analogues can prove or disprove that ancient stromatolites are useful index fossils. But even within the group of workers who accept the taxonomic approach, there are basic differences in philosophy which pose a problem to any attempt to synthesize world-wide data on stromatolite time-ranges.

THE EXTENT OF THE DATA

Despite these difficulties, there are sufficient data from the various continents to justify a preliminary appraisal of stromatolite time-ranges on a world-wide basis. The following discussion is far from complete but suggests the extent of our present knowledge. The subject is dealt with comprehensively by Semikhatov (1974).

U.S.S. R.

The first preliminary results of the Russian authors were summarized by Keller et al. (1960). Since then, the detailed descriptions of stromatolites from the southern Urals (Krylov, 1963), the Yenisey Mountains and Turukhan area (Semikhatov, 1962), northem Siberia (Komar, 1966), Tien- Shan and Karatau (Krylov, 1967a), the Uchur-Maya region (Nuzhnov, 1967) and the Polyudov Mountains (Raaben, 1964b) have laid the foun- dations for modern stromatolite taxonomy and biostratigraphy . Within the U.S.S.R., the Late Precambrian has been subdivided into four chronostratigraphic intervals, each characterized by a specific stromatolite assemblage: the Early, Middle and Late Riphean, and the Vendian. But the considerable overlap in the ranges of some of the groups often necessitates identification to form level in order to characterize each assemblage. No group is restricted to the Early Riphean, for example, but the assemblage Kussiella kussiensis Krylov, with forms of Omachtenia Nuzhnov, Gongylina differenciata Komar and Nucleella figurata Komar, characterizes the Uchur Series, generally correlated with the type Lower Riphean in the Anabar region (Keller, 1973). Komar (1966) described the columnar forms Kussiella kussiensis Krylov, K. uittata Komar, and Microstylus perplexus, the stratiform stromatolites Stratifera flexurata, S. undata and the cumulate Gongylina differenciata and Nucleella figumta from the Early Riphean Kotuikan suite of the type section


and Kyutingdin suite of the northern Siberian platform. The Burzyan Series (southern Urals), with, e.g., Kussiella kussiensis, Conophyton cylindricum has also been considered as Early Riphean, but Keller (1973) has expressed the view that it may be partly older. Conophyton is widespread in the Early Riphean, but none of its forms is restricted to it (Komar et al., 1965a). The Early Riphean can be characterized only by form-level identifications and by the absence of stromatolites first recorded from younger units.

The Middle Riphean is characterized by the presence of new groups (Ana- baria Komar, Baicalia Krylov, Suetliella Shapovalova) with abundant Cono- phyton (C. cylindricum, C. lituum, C. metulum, C. garganicum Korolyuk) and Jacutophyton Shapovalova. Of these Anabaria has been found only in the Anabar vicinity (Komar, 1966) and Svetliella in the Uchur-Maya region (Krylov et al., 1968) and the Urals (Keller, 1973), but Baicalia is very widespread throughout the U.S.S.R. (e.g. southern and central Urals, Tien Shan, Turukhan region, River Maya, Anabar Massif, Olenek Uplift and Kharaulakh Highlands). Jacutophyton is known from the Aldan Anteclise, Turukhan area, Yenisey Mountains, southern Ural and other areas (Krylov and Shapovalova, 1970b). Tungussia Semikhatov is common in the Yenisey Mountains and Turukhan area (Semikhatov, 1962) and also in younger units elsewhere.

The Late Riphean is marked in the U.S.S.R. by a significant increase in taxonomic diversity of stromatolites, a number of groups making their first appearance and/or expansion here. The most widespread groups are Gym- nosolen Steinmann, Minjaria Krylov, Inzeria Krylov, and Jurusania Krylov, known from the Urals, Tien Shan and Siberian Platform. Kotuikania Komar, Boxonia Korolyuk, Katauia Krylov, Turuchania Semikhatov and Conophy- ton miloradouici Raaben are more restricted geographically. Rare Baicalia persists into the lower part of the Upper Riphean in the Turukhan and Uchur-Maya regions, together with typical Late Riphean stromatolites. The position of the base of the Upper Riphean in Siberia is currently under discussion (Keller, 1973; Komar, 1973).

The stromatolites of Vendian or Yudomian sequences have been described by Krylov (1967a), Semikhatov et al. (1967a, 1970), Korolyuk and Sidorov (1969,1971) and Krylov et al. (in Rozanov et al., 1969). The Uk suite of the southern Urals, correlated with the Vendian, contains Linella ukka Krylov and L. simica Krylov. L. auis Krylov occurs in uppermost Precambrian beds of Tien Shan and Lesser Karatau together with Patomia ossica Krylov. Cono- phyton gaubitza Krylov also occurs in Vendian equivalents in these regions. Semikhatov et al. (1970) recognized L. simica also from the Yudoma suite of the Aldan Shield, together with Aldania sibirica Krylov. Boxonia, Patomia, Jurusania and certain cumulate stromatolites are widespread in the Yudoma suite and its Siberian equivalents.

The chronostratigraphic ranges of stromatolites described from the U.S.S.R. have been summarized by Walter (1972a, table 6), Hofmann (1973,

362 W.V. PREISS

fig. 9 ) and Semikhatov (1974, figs. 13, 14). Kussiella and Omachtenia are characteristic of, but not restricted to, the Early Riphean. The association of Baicalia, Anabaria and Svetliella with Conophyton and Jacutophyton is typical of the Middle Riphean; Baicalia, Conophyton and Jacutophyton are the most abundant but also extend into the Late Riphean, which is characterized by a divexse assemblage of Gymnosolen, Minjaria, Boxonia, Znzeria, Katavia and others. The Vendian is characterized by Linella, Patomia, Aldania with Boxonia and Jurusania. Different forms of Conophyton and Jacutophyton characterize each subdivision. The data for the U.S.S.R. are promising and indicate the potential for worldwide correlation if the time-ranges of the defined taxa can be duplicated outside the U.S.S.R.

Australia

Taxonomic stromatolite studies were aimed at testing the Russian biostratigraphy. Preliminary results were published by Glaessner et al. (1969), amplified by Preiss (197213, 1973b) and Walter (1972a) and summarized by Walter and Preiss (1972).

In summary: (a) The Lower Proterozoic Mt. Bruce Supergroup of Western Australia contains Alcheringa narrina Walter, Pilbaria perplexa Walter, Patomia f. indet. and Gruneria f. nov. (b)The unconformably overlying Bangemall Group contains Baicalia capricornia Walter and Conophy ton garganicum australe Walter, an assemblage closely resembling that of the Middle Riphean of the U.S.S.R. (c) The Burra Group of the Adelaide Geo- syncline contains Baicalia burra Preiss closely resembling Russian forms from the upper Middle Riphean, and Tungussia, and Conophyton garganicum occurs out of stratigraphic context in a diapir. (d) The Bitter Springs Form- ation of the Amadeus Basin contains an assemblage including groups typical of the Late Riphean in the U.S.S.R. (Boxonia, Minjaria, Znzeria, Jurusania, Kotuihania), Linella avis Krylov found in the Vendian of Tien Shan and new groups Kulparia Preiss and Walter and Basisphaera Walter. Very widespread is Acaciella australica Walter, type form of a group as yet undescribed from outside Australia. (e) The Umberatana Group of South Australia, with basal glacial sediments unconformably overlying the Burra Group is also rich in stromatolites with Late Riphean and Vendian affinities - Inzeria, Boxonia, Jurusania, Katavia, Tungussia, Linella, Acaciella and Kulparia. (f) The Ring- wood Member of the Pertatataka Formation, between two glacials in the Amadeus Basin, contains Tungussia inna Walter, now also identified from South Australia, just below the Ediacara fauna-bearing Pound Quartzite (Preiss, 1971a, unpublished). (g) Cloud and Semikhatov (1969b) described Kussiella in the Elgee Siltstone, Western Australia, Conophy ton garganicum from the Amelia Dolomite, Northern Territory (described in more detail by Walter, 1972a), Eucapsiphora paradisa from the Paradise Creek Formation, Queensland, and Znzeria cf. tjomusi from the Hinde Dolomite and,


doubtfully, Dook Creek Dolomite, Northern Territory. The latter was disputed by Walter and Preiss (1972). Glaessner et al. (1969), Walter (1972a) and Preiss (1971a) proposed certain correlations with the U.S.S.R. assuming that the stromatolites have similar time-ranges in the two continents. The Bangemall Group and the Burra Group were correlated with the Middle Riphean; the Bitter Springs Formation (below an older tillite) was considered as Late Riphean; the Umberatana Group and the Ringwood Member, between two tillites, were considered to be close to the Late Riphean-Ven- dian boundary in age. The correlation of the early Adelaidean Burra Group with the Middle Riphean requires further comment: it was based on the presence of Baicalia burra, similar to 3. maica Nuzhnov from the Lakhandin suite, thought to be uppermost Middle Riphean (Nuzhnov, 1967, Krylov et al., 1968), and on the absence of typical Late Riphean forms. Russian data show that Baicalia extends into the Late Riphean in the U.S.S.R., and that the Lakhandin suite also contains some typical Late Riphean groups (Komar, 1973). (This is why the Middle/Upper Riphean boundary in Siberia is contentious.) New finds by Preiss (unpubl.) of Acaciella cf. austmlica and Gymnosolen cf. rammyi, from basal Adelaidean beds in South Australia (several thousand metres stratigraphically below Baicalia burra), indicate an overlap in the time-range of Baicalia with those of Acaciella and Gym- nosolen. Correlation of the Skillogalee Dolomite with the Middle Riphean is therefore not justified, and a Late Riphean age is preferred for the basal beds and for the whole early Adelaidean. But the absence of Late Riphean forms from the Burra Group is a point in favor of some environmental control determining which stromatolites grew at these two stratigraphic levels.

North Africa

Most notable among the few other available taxonomic treatments is the work on the stromatolites of northern Africa (Bertrand, 1968a; Bertrand- Sarfati, 1972c, d). In the Taoudenni Basin the Late Precambrian sequence contains five distinct stromatolitic horizons. In summary, the oldest assemblage has Conophyton, Baicalia, Tungussia, Tilemsina Bertrand-Sarfati, Par- mites Raaben and Jacutophyton. The second is characterized by Gym- nosolen. The third has Serizia Bertrand-Sarfati and Tarioufetia Bertrand- Sarfati with Tungussia. The fourth with Jurusania, Nouatila Bertrand-Sarfati, Gymnosolen, Tungussia and Inzeria occurs only locally. The fifth, with Tifounkeia Bertrand-Sarfati and microproblematica, may have equivalents in the Schisto-Calcaire with Linella, in the West Congo. North African stromatolite assemblages were compared with those of the U.S.S.R. by Bertrand- Sarfati and Raaben (1970).

364 W.V. PREISS

India

The great potential for detailed studies in India is shown by references to stromatolites by numerous authors. In particular, Valdiya (1969a) has described and illustrated assemblages of forms from the Lesser Himalayas and the Vindhyachal Hills. His unfortunate decision to reject the modern nomenclature (he retains the Group name Collenia for many columnar stromatolites) renders comparison difficult and limits the precision of his identifications. The following taxonomically treated Indian stromatolites occur: (a) The Fawn Limestone (Lower Vindhyan) contains Conophyton cylindricum, Collenia columnaris. (b) The Bhander Limestone (Upper Vindhyan) contains Collenia baicalica (i.e. Baicalia in the modem nomenclature). (c) This form also dominates in the Gangolihat Dolomites of the Calc Zone of Pithoragarh together with Collenia columnaris, C. kussiensis (i.e. Kussiella), C. symmetrica, C. buriatica (i.e. Minjaria). (d) The Thalkedar Limestone, two formations above, contains Jurusania and Collenia sym- rnetrica. (e) The Deoban Limestone of Chakrata and the Shali Series of Himachal Pradesh have been correlated with the Calc Zone and contain C. baicalica, C. symmetrica, C. buriatica and C. columnaris (Lower Shali) and ?Jurusania (Upper Shali). Valdiya concluded an Early Riphean age for the Fawn Limestone, and Middle Riphean for the Calc Zone, Lower Deoban, Lower Shali and Bhander Limestone. The Gangolihat and Lower Shali could be transitional between the Middle and Late Riphean. The younger assemblage of the Thalkedar Limestone, Upper Deoban and Upper Shali compares with the Late Riphean.

Other areas

Here the information is as yet scanty. Cloud and Semikhatov (1969b) have given the only modem taxonomic account of stromatolites from southern Africa: (a) Archaean stromatolites from the Bulawayo Dolomite, Rhodesia; (b) Katernia africana Cloud and Semikhatov from the Dolomite Series, South Africa (a similar form appears in the Nash Formation, Wyoming, U S A ) ; (c) Baicalia aff. rara Semikhatov from the lower Abenab Formation, Tsumeb, S.W. Africa; (d) Conophyton ressoti Menchikov from the upper Abenab Formation. These authors also described the following forms from North America: (e) the Nash Formation forms mentioned above; (f) Gruneria biwabikia Cloud and Semikhatov from the Biwabik and Gun- flint Iron Formations; Gruneria also occurs in the Ventersdorp System, South Africa and in the Fortescue Group, Western Australia (Walter, 1972a); (g) Conophyton cylindricum from the Siyeh Limestone, Belt Supergroup and Mescal Limestone, Apache Group; (h) Tungussia in the Mescal Lime- stone; (i) Linella aff. ukka and Boxonia cf. gracilis in the Johnnie Formation, California. The Crystal Springs Formation, unconformable below the


possibly glacigenic Kingston Peak Formation, contains Baicalia and Jacu to- phyton (Howell, 1971). White (1970) recorded Baicalia from the Belt Super- group, whose possible correlatives in Alaska apparently also contain Middle Riphean stromatolites (M.A. Semikhatov, pers. comm. in Churkin, 1973). The equivalent Purcell Supergroup of Canada contains Baicalia and Cono- phyton and the Helikian Little Dal Formation of the Mackenzie Mountains contains Gymnosolen (M.A. Semikhatov, pers. comm.). Jacutophyton and Conophy ton are reported from the Athapuscow Aulacogen (Early Proterozoic) of Canada (M.R. Walter, pers. comm. in Hoffman, 1973). Hofmann (1969) discussed Early Proterozoic stromatolites from the Animikie Group, Canada, making kntative comparisons with previously named forms rather than firm identifications. These included Gymnosolen, restricted in the U.S.S.R. to the Late Riphean, which is a major anomaly requiring further investigation. However, Walter (1972b) has shown that at least some of these structures are of inorganic origin.

Moeri (1972) and Cloud and Dardenne (1973) described Conophyton from the Bambui Group of central Brazil and Cowie (1961) noted Conophy- ton in the Fyn S6 Dolomite (Late Precambrian) of Greenland. M.R. Walter (pers. comm., 1973) has identified the latter as Conophyton cf. cylindricum. These finds all give encouragement for further searching in other continents.

THE DATING OF STROMATOLITIC SEQUENCES

Unfortunately, reliable, precise radiometric ages for Precambrian sedimentary rocks are extremely scarce. What data are available are frequently controversial, yet they provide the only objective method of determining stromatolite time-ranges, and the .only ultimate test of intercontinental correlations. What is needed but cannot be attempted here is an independent evaluation by geochronologists of the varied and scattered data available from the various countries and rocks, which were based on different analytical methods and decay constants. Until this is achieved, the stratigrapher can only cursorily examine the published data and accept them at face value.

Piper’s (1973) analysis of palaeomagnetic and geochronological data shows that many of the Late Precambrian tillites formed in low latitudes, and therefore represent global climatic events. In addition, the ages for an “upper” glaciation cluster in the 650- 700-m.y. interval, e.g., in Australia (Compston and Arriens, 1968; Dunn et al., 1971), the U.S.S.R. (Chumakov, 1971), Norway (Pringle, 1973), and Africa (Cahen, 1970). These data give encouragement that tillites may provide useful time markers in Late Precam- brian stratigraphy, and therefore help define the time-ranges of stromatolites.

Fig. 1 summarizes the relevant published radiometric ages of the sequences discussed above. It should be pointed out that the majority of Russian dates are single K-Ar determinations on glauconite. .

W.V. PREISS

UCHURO-MAYA

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366

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Fig. 1. For caption see p. 368.


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3 68 W.V. PREISS

TIME-RANGES OF THE BEST KNOWN TAXA

A brief summary of the ranges of some important taxa discussed is as follows:

(1) Pre-Riphean stromatolites: Of these only Gruneria and Katernia africana have so far been identified in more than one continent, and their ranges may be of the order of 1,900 to 2,300 and 1,700 to 2,000 m.y., respectively. The occurrence of Gymnosolen-like stromatolites and Patomia of this age is suggested by Hofmann (1969) and Walter (1972a), and requires further investigation.

(2) Late Precambrian stromatolites: The following important taxa appear at the present time to have restricted time-ranges:

3 Conophyton cylindricum, C. garganicum: Early and Middle Riphean (1,600 to about 1,000 m.y.)

Conophyton metula, Svetliella, Anabaria: Middle Riphean (1,350 to about 1,000 m.y.)

Tungussia: Middle Riphean to Vendian (1,350 to 570 m.y.). Baicalia: Middle and early Late Riphean (1,350 to about ?800 m.y.). Inzeria, Minjaria, Katauia, perhaps Gymnosolen : Late Riphean (about

Boxonia, Jurusania, Linella: Late Riphean to Vendian (about 1,000 m.y. 1,000 m.y. to 680 m.y.).

to 570 m.y.).

Fig. 1 (pp. 366-367). Geochronology and sequences of stromatolite assemblages. A. Type sections of the U.S.S.R. B. Sequences on the other continents.

Notes: The age column in A applies to both figures, and indicates absolute ages in m.y. The black areas at the boundaries of the Russian chronostratigraphic intervals indicate the estimated uncertainty of their ages. Two alternative positions for the Middle Riphean/Late Riphean boundary are presently in use. The lower case names refer to suites in A and to formations in B. Numbers 1-21 refer t o sources of age data. indicates a single age determination quoted; $ indicates a range of ages quoted; and $ indicates a single age determination with known limits of error. The rock units are placed in the figure wherever possible according to their absolute ages.- indicates unconformity. Where a sequence is possibly older than indicated by its position, this extension is shown by an arrow. Roman numerals refer t o stromatolite assemblages: I = assemblages with Kussiella, Omachtenia, Conophyton cylindricum and others; I1 = assemblages characterized by Baicalia, Anabaria and Suetliella; 111 = assemblages with the first appearance of Gym- nosolen, Minjaria, Boxonia, Inzeria, Jurusania and others; IV = assemblages with Linella, Boxonia, Patomia and Aldania. References: 1 = Krylov (1963); 2 = Komar (1973); 3 = Keller (1973); 4 = Semikhatov et al. (1970); 5 = Nuzhnov (1967); 6 = Komar (1966); 7 = Krylov, in Rozanov et al. (1969); 8 = Semikhatov (1962); 9 = Khomentovskii e t al. (1972a); 10 = Compston and Arriens (1968); 11 = Cooper and Compston (1971); 12 = Compston et al. (1966); 13 = Clauer (1973); 1 4 = Cahen (1970); 15 = Bertrand-Sarfati (1972d); 16 = Cloud and Semikhatov (1969b); 1 7 = Tugarinov et al. (1965); 18 = Obradovich and Peterman (1968); 19 = Shride (1967); 20 = Compston and Taylor (1969); 21 = Majoribanks and Black (1974);22 = Krylov and Shapovalova (1970b).


Linella ukka, L. simica, Aldania: Vendian (680 to 570 m.y.). In Australia, the geochronological ranges of most forms are known with

even less precision owing to the poor radiometric control on the ages of the major Late Precambrian sedimentary basins. The following groups are so far known only from Australia:

Acaciella is most widespread, and probably ranges from less than 1,000 m.y. (Late Riphean) to Early Cambrian.

Basisphaera, known only from the Bitter Springs Formation, is thus probably Late Riphean.

Georginia, closely allied to Jacutophyton, is either Early Cambrian or latest Precambrian.

Kulpariu is known only from Late Riphean or Vendian equivalents. Madiganites is Middle and Late Cambrian. Eucapsiphora occurs in rocks considered to be about 1,600 m.y. old. Certain sedimentary rocks from North Africa are now dated and suggest

the best known stromatolitic sequence in the. Taoudenni Basin is entirely less than 1,000 m.y. old. The oldest assemblage (Conophyton, Baicalia, Tun- gussia, Tilemsina, Parmites, Jacutophyton) is therefore already of Late Riphean age. The youngest assemblage with Linella may well be Vendian. These correlations, based on geochronology, would fit the general time- ranges known for the Riphean stromatolites, and confirm that Baicalia occurs in the Late Riphean together with other forms appearing here for the first time.

CONCLUSIONS

The biostratigraphic study of stromatolites is still in its infancy. In the last fifteen years of active publishing, a biostratigraphic scheme has been erected which appears to be valid throughout the Soviet Union. Preliminary indications from the other continents, notably Australia and North Africa, indicate that many of the Russian taxa (many of the groups and several forms) have widespread, perhaps global distribution. The determination of their time-ranges, however, depends on geochronology and direct datings of sedimentary rocks are scarce. However, what data we do have suggest that at least certain groups and forms have restricted time-ranges, and that where assemblages of these are found in Precambrian rocks, tentative correlations can be proposed with known sequences, especially in the U.S.S.R. However, these correlations must be regarded with caution, since they are liable to change as more geochronologic data become available, thus changing the known time-ranges of the stromatolites. A major problem to be resolved is that of pre-Riphean stromatolites with a young aspect. These forms need to be carefully identified, and rocks intermediate in age need to be searched. This will indicate whether these stromatolites are the same as the Late

370 W.V. PREISS

Riphean forms, and hopefully, whether they represent convergent evolution or simply much longer time-ranges than previously suspected.

ACKNOWLEDGEMENTS

Many of the ideas expressed arose from a study of South Australian stromatolites, commenced under the guidance of Prof. M.F. Glaessner, who first introduced me to the subject of stromatolites. I am indebted to Drs M.R. Walter, M.A. Semikhatov and I.N. Krylov for valuable discussions and correspondence. Drafting by the Drafting Branch, Department of Mines, Adelaide, S.A., is gratefully acknowledged. The paper is published with the permission of the Director of Mines.


Chapter 7.3

APHEBIAN STROMATOLITES IN CANADA: IMPLICATIONS FOR STROMATOLITE ZONATION

J.A. Donaldson

INTRODUCTION

Meticulous studies of stromatolites by Soviet workers have contributed significantly to subdivision of the Proterozoic rock record in the U.S.S.R. (Raaben, 1969a, and earlier references). Apparent restriction in the time- ranges of distinctive “groups” and “forms” of stromatolites from widely separated regions within the Soviet Union has not only aided correlation within the U.S.S.R., but has also provided support for the possibility of using stromatolites as a basis for world-wide zonation of Proterozoic strata. This possibility has been enhanced by discoveries of comparable distinctive stromatolites in Proterozoic strata of other continents (Cloud and Semikhatov, 1969b; Glaessner et al., 1969; Valdiya, 1969a; Bertrand-Sarfati, 1 9 7 2 ~ ) . Particularly impressive has been the detailed and systematic classification of numerous Proterozoic stromatolites in Australia by Preiss (1972) and Walter (1972a), following the Soviet methods of reconstruction and identification. In all these studies, stromatolites of Riphean age (time-range approximately 1,650- 570 m.y.) have received most attention, although Cloud and Semikhatov (196913) and Walter (1972a) also described some pre-Riphean forms. Most of the distinctive groups of stromatolites are reported to have ranges that fall entirely within the Riphean, and therefore an obvious method of testing the validity of presently designated ranges of Riphean stromatolites is to compare them with stromatolites from significantly older and younger strata. Because of .the greatly reduced abundance of stromatolites in the Phanerozoic, attributed by Garrett (1970b) to the evolution of grazing and burrowing organisms, strata of pre-Riphean age appear to offer the most promise for comparisons.

372 J.A. DONALDSON

Fig. 1. Stromatolite-bearing sequences of Aphebian strata in Canada: 1 = Epworth Group; 2 = Snare Group; 3 = Goulburn Group; 4 = Great Slave Supergroup; 5 = Hurwitz Group; 6 = Gunflint Group; 7 = Sutton Lake Group; 8 = Belcher Group; 9 = Manitounuk Group; 10 = Mistassini Group; 11 = Kaniapiskau Supergroup.

AF'HEBIAN (LOWER PROTEROZOIC) SEDIMENTARY ROCKS OF CANA,DA

Stromatolite-bearing Proterozoic rocks of Aphebian age (time-range approximately 2,500-1,750 m.y.) are well exposed in several regions of Canada (Fig. 1). The minimum age for Aphebian rocks has been established by numerous radiometric age determinations of events that clearly postdate Aphebian sedimentation, and thus, even allowing wide margins for the precision and accuracy of radiometric data (many Riphean ages in the U.S.S.R. derive from K-Ar measurements on glauconite; most Aphebian ages in Canada are based on K- Ar and Rb-Sr determinations of igneous and metamorphic rocks), there is little doubt that rocks classified in Canada as Aphebian are also pre-Riphean.

Several studies to provide data for comparisons of Aphebian and Riphean stromatolites have been initiated recently in Canada. Although considerable work remains, sufficient information has been accumulated to indicate that features characteristic of Riphean stromatolites are also displayed by some Aphebian stromatolites. Photographs of stromatolites from the Aphebian Manitounuk Group (Hofmann, 1969a) and a subsequently published graphic reconstruction of the same stromatolites (Hofmann, 1973) show

STROMATOLITE BIOSTRATIGRAF'HY 313

I

Fig. 2. Longitudinal sections of Aphebian stromatolites, Belcher Group: a. Furcate branching characteristic of kussiellid stromatolites such as Kussiella, Omachtenia and Boxonia *. b. Digitate branching characteristic of gymnosolenid stromatolites; most have smooth margins cf. Gymnosolen; the multibranched cluster resembles Anabaria. c. Ragged subcylindrical columns with numerous linkages, cf. Jurusania. d, e. Stubby tuberous stromatolites cf. Baicalia. f . Stromatolite in which part of one branch shows conical laminations and an axial zone, cf. the branched, conically 'laminated stromatolite, Jacutophy- ton. g, h. Bifid branching and development of wall structure, cf. Linella. i. Branching from niches in parent column, cf. Znzeria. j, k, 1. Markedly divergent branching characteristic of the tungussid stromatolite Tungussia.

* Although Raaben (1969a) classified Boxonia as a member of the gymnosolenid supergroup, it is logically a member of the kussiellid supergroup, according to its furcate style of branching.

374 J.A. DONALDSON

characteristics of the Upper Riphean group Katauia. Stromatolites that resemble Gymnosolen, Tungussia, Baicalia and Jacutophyton occur in Aphebian strata of the Epworth Group and Great Slave Supergroup (see Hoffman, Ch. 10.7). Additional examples of stromatolites showing sets of morphological features characteristic of Riphean stromatolites occur in Aphebian strata of the Belcher Group; examples of these are illustrated herein.

STROMATOLITES OF THE BELCHER GROUP

The Belcher Group (Fig. 1) contains numerous units of limestone and dolostone that are richly stromatolitic (Jackson, 1960). The stromatolites display considerable diversity, ranging from low domal forms to complexly branched columns. The branched columnar stromatolites are of principal interest here, because most of the presently designated stromatolite ranges apply to branched columnar forms that are characterized by distinctive sets of morphological features.

Fig. 2 illustrates several stromatolite varieties common in the Belcher Group that resemble Riphean and Vendian stromatolites. Diameters of the Belcher columnar stromatolites commonly exceed 20 cm, and therefore satisfactory reconstruction by the method of serial sectioning is generally impractical. However, numerous sections parallel and perpendicular to bedding are provided by excellent exposures in the field, and many details of branching, column ornamentation and configuration of laminae can be established by a study of such exposures, without recourse to systematic sectioning. All of the varieties shown in Fig. 2 occur in laterally extensive units, some of which can be matched in measured sections as much as 50 km apart. It is important t o add, however, that the lithologies and primary sedimentary structures of some of these units are equally persistent, a relationship that has been established for Aphebian strata of other regions, such as the Knob Lake Group of the Kaniapiskau Supergroup, Labrador Geosyncline (Donaldson, 1963), and the Great Slave Supergroup of the Coronation Geosyncline (Hoffman, 1968 b).

Representative longitudinal sections of columnar stromatolites shown in Fig. 2 are grouped to show various styles of branching: furcate and digitate branching of subparallel columns (Figs. 2a, 2b, 2c), slightly divergent branching of stubby columns (Figs. 2d, 2e, 2f, 2g, 2h, 2i) and markedly divergent branching of dendroid stromatolites (Figs. 2j, 2k, 21). As a basis for comparison with Riphean stromatolites, the summary descriptions in Raaben (1969a), Cloud and Semikhatov (1969b) and Hofmann (1969a) have provided the primary references. Furcate branching of subparallel columns, in which division into smaller columns occurs without increase in total width of the stromatolite, can be widely recognised in the Belcher Group (Fig. 2a).


Fig. 3. Longitudinal section of Aphebian columnar stromatolites, Belcher Group. Furcate or kussiellid branching is apparent and the smooth-margined columns show some envelopment of successive laminations (wall structure).

Some have ragged margins and overhanging laminae suggestive of the named stromatolites Kussiella and Omachtenia, but more commonly they have smooth margins, locally enveloped by either selvages or wall structures (Fig. 3). Some of these display columns of very uniform diameter (left illustration of Fig. 2a), and thus are close in general appearance to Boxonia.

Digitate branching of subparallel columns, in which there is a distinct increase in total width of the stromatolite at points of division, gives rise to abundant smooth-walled forms in the Belcher Group that closely resemble Gymnosolen; a less common style of ramification into numerous subparallel branches is displayed by bouquet-like stromatolites similar in appearance to Anabaria, but generally lacking the slightly divergent branching of this group (Fig. 2b). Some parallel columns display ragged margins, coalescence, and sporadic linkage (Fig. 2c), features that are common in published illustrations of Jurusania and Omachtenia.

Stubby columnar stromatolites of the Belcher Group show a wide diversity of shapes and surface ornamentation, but most are characterized by dendroid branching, and resemble Baicalia (and to a lesser extent, Tungussia) in longitudinal section (Figs. 2d, 2e, 4). Some Baicalia-like stromatolites display sharply convex laminations, with lamination thickening in the axial zone (Figs. 2f, 5). Such features in unbranched columnar stromatolites are

376 J.A. DONALDSON

Fig. 4. Longitudinal section of Aphebian stromatolites, Belcher Group, showing irregular Boicolia-like branching. Note distinct envelopment (wall structure) along margins of several stromatolites. Markings on white scale are at 5-cm intervals.

diagnostic of Conophy ton; their occurrence in branched dendroid stromatolites invites comparison with Jucutophy ton. Another common feature of the stubby columnar stromatolites is development of a “wall structure”, in which the marginal parts of laminae successively envelop the lateral surfaces of the columns. These stromatolites generally possess sharply defined smooth margins, and thus are somewhat similar to some forms of Linella (Figs. 2g, 2h, 6). Some of the stubby columnar stromatolites display niches in parent columns where branching occurs, a feature characteristic of Inzeria (Figs. 2i, 7 ) .

Most distinctive of all are stromatolites showing markedly divergent branching, which is characteristic of Tungussiu (Figs. 2j, 2k, 21, 8). Margins of the columns are typically smooth and sharply defined, and individual branches range from uniform in diameter to strongly flared. Niches occur at the base of some columns, giving profiles that can also be compared to, Znzeria (Fig. 2k). Segments of some columns are apparently walled, but these may merely represent oblique sections through columns that are steeply inclined to surfaces of exposure in outcrop.


Fig. 5. Longitudinal section of Aphebian stromatolite, Belcher Group, showing Conophy- ton-like segment in one branch, cf. Jacutophyton.

SUMMARY

Although extensive serial sectioning and study of microstructures will be required to fully document and identify Aphebian stromatolites for accurate comparisons with Riphean stromatolites, some generalizations are warranted on the basis of present information.

(1) Columnar stromatolites of pre-Riphean age display considerable diversity; the number of distinctive pre-Riphean stromatolite morphologies is clearly greater than indicated by Awramik (1971a) in his graph based on named columnar stromatolites.

(2) Many Aphebian stromatolites display styles of branching characteristic of Riphean stromatolites. All four supergroups of Raaben (1969a) are represented (conophytonids, kussiellids, gymnosolenids and tungussids).

(3) Specific characteristics used for classification of Riphean stromatolites, notably wall structures and niches, are common in Aphebian stromatolites.

(4) Although comparisons presented here are preliminary, sufficient basic similarities are apparent, without detailed reconstructions, to suggest that some named Riphean stromatolites also occur in pre-Riphean strata.

378 J.A. DONALDSON

Fig. 6. Longitudinal section of smooth-margined Aphebian stromatolite, Belcher Group, showing well-developed wall structure, cf. Linella.

Fig. 7 . Longitudinal section of Aphebian stromatolites, Belcher Group, showing niches, a characteristic of Inzeria


Fig. 8. Divergent branching in Aphebian stromatolite, Belcher Group, cf. Tungussia. Width of view: 70 cm. Linear surface markings are glacial striae.

CONCLUSIONS

The utility of stromatolites for intrabasinal correlation is well established (lithostratigraphy), and certain stromatolites have been used with considerable success for interbasinal correlation of Precambrian strata in the Soviet Union. Recently, hopes for extension of stromatolite correlation to the intercontinental level (biostratigraphy) have been raised by discoveries of distinctive stromatolites of similar Precambrian ages in Australia, India, Africa and North America. Serious problems for such world-wide correlation are posed, however, by discoveries of specific stromatolites outside their currently accepted time-ranges. An important example of this is the occurrence of Aphebian stromatolites that closely resemble some of the distinctive “index” stromatolites currently considered to have first appeared in Riphean (and hence post-Aphebian) time. Similarities of this type have been previously reported by Walter (1972a, p. 84), who suggested that certain stromatolites might have “disjunct distributions”, appearing briefly in pre-Riphean strata, and then again in Riphean strata.

Considerable work is needed to firmly establish whether some pre- Riphean stromatolites, such as those briefly described and illustrated here, match in all respects (including microstructure) the Riphean stromatolites they resemble in longitudinal sections. Nevertherless, field and laboratory

380 J.A. DONALDSON

studies have progressed sufficiently to indicate the need for great caution in any attempt to assign geologic age on the basis of stromatolites alone.

ACKNOWLEDGEMENTS

Field work on the Belcher Islands was supported by the National Research Council of Canada. Subsequent laboratory study at Carleton University has been advanced with the assistance of Peter Hews. The manuscript was prepared during sabbatical leave spent in the Department of Geology and Mineralogy, University of Adelaide, South Australia. I am indebted to Wolfgang Preiss for his helpful comments on the manuscript.


Chapter 8.1

OPEN MARINE SUBTIDAL AND INTERTIDAL STROMATOLITES (FLORIDA, THE BAHAMAS AND BERMUDA)*

Conrad D. Gebelein

INTRODUCTION

The most important factors controlling the distribution and morphology of Recent subtidal and intertidal stromatolites in open marine waters are: (1) the widespread distribution and high abundance of grazing and burrowing invertebrates; (2) the lack of rapid cementation of organic and organosedimentary structures.

OPEN MARINE, SHALLOW SUBTIDAL ALGAL MATS AND STROMATOLITES

Several types of algal sediments occur in the shallow subtidal'zone. Their form, distribution and vertical or lateral sequences provide detailed information on depositional environments.

Algal mats

Surficial mats composed of organic materials (blue-greens, greens, diatoms, red algae, invertebrate mucilage) and bound sediment grains (Bathurst, 1967b; Neumann et al., 1970) are absent only in shallow subtidal environments where: (1) rates of surface sediment turnover by grazers and especially by burrowers arevery high; and (2) rates of sediment movement over the surface preclude any binding of grains by organic mucilage (Gebelein, 1969). In very clear waters, as around Bermuda, algal mats have been observed to depths of 50m, on the forereef slope. In general, however, growth is most rapid in depths of less than 10 m. These surficial mats represent a dynamic interface at which sediments below and above the mat are constantly being reworked, deposited on the surface, reincorporated into the mat, etc. - i.e. the mat layer is always the surface layer. Hence, laminites are not produced. These

* Contribution number 617, Bermuda Biological Station for Research.

382 C.D. GEBELEIN

mats greatly retard sediment erosion (Bathurst, 1967b; Scoffin, 1970; Neumann et al., 1970). Erosion of the mats characteristically produces flat pebbles or aggregates of bound sediment. On portions of the Great Bahama Bank, these pebbles are slightly lithified (Purdy, 1963; Bathurst, 1971; Winland and Matthews, 1974).

Algal stromatolites

(1) Distribution. Stromatolites occur in open marine subtidal environments in which, for various reasons, invertebrate grazers and burrowers are greatly reduced in numbers (Gebelein, 1969). Within these narrow environmental limits, stromatolite growth may begin on any surface irregularity (stabilized cipple, flat pebble of algally bound sediment, Thulussia rhizome, elevated reef or rock substrate) which elevates the surface above the surrounding mat- bound bottom (Fig. 1A). Maximum observed depths for subtidal stromatolite formation are 4-5m (Gebelein, 1969, and pers. obs.).

(2) Morphology and internal structure. Subtidal stromatolites range in dimension from tiny bumps, only a few millimeters in diameter and height and containing 1 or 2 convex upward laminae, to large domes, up to lOcm high and 30 cm wide (Fig. 1B). Structures are symmetrical in areas with a symmetrical (or no) sediment supply, and are markedly asymmetrical in areas with asymmetrical sediment supply. Stromatolites are elongate parallel to the direction of predominant sediment supply. Where sediment supply is from one direction only, laminae are thickened in that direction.

The lamination fabric consists of alternate thick layers (in which algal filaments are largely vertical) that form by growth and phototactic movement during the day, and thin layers (in which filaments are largely horizontal) that form during the night (Monty, 1965a, 1967; Gebelein, 1969). In areas of abundant sediment, sediment-grain density is highest in the thick layer, due to enhanced permanent binding by vertically oriented filaments. These stromatolites accrete at the rate of up to 1 mm (one lamination couplet) per day. These forms are composed of thin-sheathed filamentous blue-green algae; Schizothrix and Oscillatoria are the most common genera.

Algal oncolites

(1) Distribution. The environment of formation of oncolites at Joulters Cays, Bahamas is similar to that for attached stromatolites, except that periodic turbulence, causing initial detachment of the forms, is higher. In the Joulters, this increase in turbulence corresponds to a decrease in depth. Maximum formation occurs in depths less than 1m. Concentrations of 10-100 oncolites per m2 are common in the shoal areas. Lags of oncolites carpet the floor of some tidal channels which cut through the shoals.

ENVIRONMENTAL MODELS 383

Fig. 1. A. Subtidal stromatolites in Tholassia bed; depth 2 m, Castle Roads, Bermuda. Black lens cap is 16 cm diameter. B. Subtidal stromatolites, Bermuda: 1 and 2, top view; 3, side view, showing bound sediment beneath, stromatolite; 4, cross section showing laminae. Scale for 1 , 2 and 3: bar = 1.5 cm; scale for 4: bar = 1.5 cm. C. Section through drape-over stromatolite which formed on the margin of a channel on intertidal mud flats, Cape Sable, Florida. Scale in cm. D. Thin section of high intertidal, domal stromatolite, Cape Sable. Arrows point to thin algal laminae between coarse sediment laminae. Bar = 0.2 mm. E. Close-up of upper right-hand corner of D, showing detail of algal lamina and sediment laminae. Bar = 0.1 mm.

384 C.D. GEBELEIN

( 2 ) Morphology and internal structure. Oncolites range in size from 1 mm to a maximum of about 10cm. Lamination structure is very similar to that in subtidal stromatolites. In the Joulters, tiny stromatolites become detached by wave or current turbulence, and form the cores of the oncolites. Once detached, the forms are kept in motion, which prevents reincorporation into the mat. Id very shallow (less than 15 cm) water, the stromatolitic core may have only one or two convex upward laminae: in some cases, no core at all is detectable. In deeper water, attached stromatolite growth may produce 10-20 laminae before detachment. Thus, as suggested by Logan et al. (1964), the degree of symmetry of oncolitic laminae provides a relative estimate of the turbulence of the environment.

Oncolites from the Frazer’s Hog Cay area, Bahamas, contain an algal- foraminifera1 consortium, with a characteristic anastornosing cellular pattern imposed on the algal laminae by the growth within the oncolite of an irregular opthalmid foraminifer (Buchanan et al., 1972). These structures, some of which are lithified, are very similar to the “Osagia” oncolites of the Paleozoic.

MODEL FOR SUBTIDAL STROMATOLITES AND ONCOLITES

The distribution of subtidal forms described above fits a simple model (cf. Gebelein, 1969), in which the controls on distribution and morphology are rate of sediment movement (as it effects stability of the mat and the distribution of invertebrates), rate of sediment accumulation, and symmetry of sediment supply. Oncolites add the dimension to this previous model of increased periodic turbulence, usually corresponding to decreased depth.

Exceptions to the rule. Completely organic stromatolites have been found on hard substrates in reef, backreef and lagoon areas (Monty, 1967; Gebelein, 1969). These structures have approximately the same size and shape as the subtidal stromatolites described above. The completely organic forms have a very tough, leathery surface texture. These forms are quite rare (although a seasonal “bloom” was observed in Bermuda in the early fall of 1974). Their distribution appears to be random within the following limits: (1) they always form on hard substrates; (2) they occur in areas where the abundance of surface grazers is low. The occurrence of these Recent forms, combined with the discovery of abundant subtidal stromatolites lacking detrital textures, in the forereef sediments of the Devonian Canning Basin (Playford and Cockbain, 1969), leads to the speculation that such deposits could occur in any reef or forereef environment which: (1) lacks abundant grazers (as most reefs do); or (2) has the potential for rapid cementation (a condition found most commonly in reefs).

ENVIRONMENTAL MODELS

INTERTIDAL LAMINITES AND STROMATOLITES

385

In high-energy sandy shore zones with continuous scour and sediment movement, stromatolites and laminites are completely lacking, due to the absence of a stable substrate (contrast this situation to that in Shark Bay). As scour decreases, stromatolites and laminites develop in the least energetic, upper portions of the tidal zone. However, development is limited to the upper half of the tidal zone, regardless of energy for scour, by reworking and inges- tion of the surface sediment layer by grazing and burrowing invertebrates.

The nature of this restriction has been documented by Gebelein and Hoff- man (1968), Garrett (1970b) and Gebelein (1971): Grazing and burrowing invertebrates are restricted to the lower portions of the intertidal zone by factors related to desiccation (drying of the animals, temperature, salinity and gas composition of the interstitial fluids) which adversely affect invertebrate viability and activity. In low intertidal areas, surficial algal mats cannot develop, as removal and reworking is more rapid than mat development. Above a certain flooding-frequency isograd, invertebrate activity and viability decrease sufficiently for surficial mats to develop. At these latter elevations, laminated sediments rapidly form by the trapping and binding of sediments onto the mat.

The position of the boundary between unlaminated and laminated sediments varies from place to place, and depends on the types of algae and invertebrates present. At Cape Sable, Florida, the boundary occurs at the 25% flooding-frequency isograd - i.e. the surface flooded 25% of the time (Gebelein, 1971). On Andros Island, Bahamas, the boundary occurs at the 50% flooding-frequency isograd (Ginsburg et al., 1970).

Within the laminated portion of the tidal zone, a further zonati0.n of algal mat types may occur. This zonation is related to the distribution of various associations of blue-green and green algae according to flooding frequency. Each association generates a specific fabric, or microstructure (Gebelein, 1974). Hence, the upper intertidal area may contain laterally and vertically zoned laminites and stromatolites with different microstructures.

Open tidal flats

(1) Morphology. In areas which are alternately flooded and uncovered on each tide, algal laminites are the characteristic form. Relief is limited to that generated by the growth mode of the algal association, and is usually less than 2-3cm (see below for one exception). Various types of domes may occur with this small amount of surface relief.

“Drape-over” stromatolites, in which the laminae follow the sediment surface curvature, may occur on the margin of tidal channels, and thus generate an asymmetrical form with several centimeters of total relief on the channel-facing side (Gebelein, 1971) (Fig. 1C).

386 C.D. GEBELEIN

( 2 ) Internal structure. Most intertidal algal associations contain a predominance of filamentous forms (Monty, 1965a; Gebelein, 1971; Golubic and Park, 1973) and hence generate laminated fabrics (Gebelein, 1974). The exact fabric varies from association to association. In general, the laminae are less than 5 mm thick and consist of an algal- and sediment-rich couplet (Fig. lD, E). Laminations are smooth, lack desiccation crinkling and crenulation, and drape over minor relief features. Lamination thickness is determined by the rate at which the dominant algal species can grow and move through the sediment layer, re-establish a surface mat, and hence permanently bind the sediment deposit. At Cape Sable, about a millimeter of sediment is deposited onto the mat surface during each tide. The blue-green algae are able to grow and move upward through this layer and permanently bind it before it is washed away by succeeding tides. On Andros levees, sedimentation occurs only during storms, when over a centimeter of sediment may be deposited (Hardie and Ginsburg, 1971). However, the rate of upward movement of the algae is such that all but about a millimeter or two of the sediment is washed away on the succeeding tides.

Algae with vertical or radial growth modes form layered or domal structures either by the “sieve” effect causing the capture of sediment (cf. Hom- meril and Rioult, 1965; Gebelein, 1971), or by cementation which generates lithified carbonate molds of the filaments and of the structure (as in the lithified Scytonema domes on the Andros tidal flats, where crusts are formed at the level of maximum “capillarity” on the flats). (3) Associated sedimentary structures. On the open tidal-flats a particular

suite of sedimentary structures occur associated with stromatolites and laminites: prism cracks, ranging in depth from a few millimeters to several centimeters; flat pebbles of laminated sediment; current crescents, developed in the lee (flood tide direction) of surface irregularities; rare birdseye fenestrae and rare open burrows (Gebelein, 1971).

Ponded tidal flats (Fig. 2 )

On flats with slight surface irregularities, ponds may form. On portions of these flats, the flooding frequency is less than 5096, taken on a yearly average, but the duration of individual flooding and exposure events is quite long - days or weeks. Increase in flooding duration is accompanied by the following phenomena: (1) prism cracks become better defined (often by drape-over laminae) and deeper; (2) individual sediment and algal laminae increase in maximum thickness; (3) lamination curling and crenulation increase. In- creased time for algal growth leads to thicker (often several millimeter) organic laminae, which characteristically contain little sediment. Thick, storm-derived sediment layers are deposited on the flats. Lacking subsequent tidal floodings, these layers are not easily dispersed or washed away, and account for the thicker sediment layers. Intense desiccation of the surface


Fig. 2. A. Domal stromatolites growing over curled algal-sediment laminae, ponded area, Cape Sable. Note well-developed prism cracks between stromatolites. Rule = 1 2 in. (30cm). B. Cross-section, high portion of ponded area, Cape Sable. Note thick algal laminae, storm deposits, crenulation and curling of laminated algal muds. Bar = 10 cm. C. Thin section of algal-laminated sediments, ponded area, Cape Sable. Note thick, multiple, crenulated algal laminae, and thick, discontinuous sediment laminae. Bar = 0.1 mm.

organic algal mats during the long periods of dryness leads to lamination crenulation, and often to the curling-up of a thick section of laminated sediment. Such curling can occur only where the organic content, and hence the cohesiveness of the deposit, is very high.

Supratidal deposits on tidal flats

Large expanses on tidal flats lie at elevations flooded only by major storms (less than 5% flooding frequency at Cape Sable; Gebelein, 1971). Sediments deposited in this environment are thin-bedded (1-3 cm thick). Each bed is the product of a single storm. Lack of tidal reworking leaves the beds intact. A single algal mat layer may develop immediately after storm flooding, but

388 C.D. GEBELEIN

usually is desiccated to the extent that it dries up and blows away. Thus, truly supratidal environments are not characterized by laminated deposits. The most common internal structure of the thin beds is the birdseye fenestra (Gebelein, 1971).

INTERTIDAL SEDIMENTATION MODEL (BAHAMAS, SOUTH FLORIDA)

Desiccation-related factors provide the basic division of intertidal flats into distinct invertebrate and algal associations. However, the position of boundaries between associations is largely the result of interaction between the associations. Such interactions produce sharp lateral changes in associ- @ion composition, and hence in sediment fabric. These changes, seen in vertical section following progradation, are usually abrupt and very likely will generate bedding planes in analogous ancient carbonates. The general sequence found in intertidal sediments is: (1) a lower intertidal burrow- mottled unit; (2) an upper intertidal algal unit (mostly thinly laminated), with the possibility for several distinct sub-units with different fabrics; (3) a supratidal thin-bedded unit, lacking or with rare algal lamination. Ponded tidal flats show the same sequence. However, in ponded tidal flats the algal section is characterized by thicker lamination, increased prismcrack formation, and crenulation and curling of the laminated deposits.


Chapter 8.2

MODERN ALGAL STROMATOLITES AT HAMELIN POOL, A HYPERSALINE BARRED BASIN IN SHARK BAY, WESTERN AUSTRALIA

P.E. Playford and A.E. Cockbain

INTRODUCTION

Algal stromatolites are forming today over wide areas of the sublittoral platform and adjacent intertidal zone fringing Hamelin Pool, a hyperialine barred basin at the southeastern extremity of Shark Bay in Western Australia (Fig. 1). They are the most diverse and abundant shallow-water stromatolites known from modem seas.

The Hamelin Pool stromatolites were found in 1954 by D. Johnstone, P.E. Playford and R.L. Chase during geological investigations of the area for West Australian Petroleum Pty. Ltd. (Wapet). In the following year Chase, with B.W. Logan, studied the Proterozoic Moora Group, which contains welldeveloped stromatolites (Logan and Chase, 1961). Chase recognized that the Hamelin Pool stromatolites were modern analogues of these Proterozoic forms, and at his request Playford collected some stromatolite heads from Hamelin Pool for further study. Algae in these samples were examined by D.M. Churchill.

B.W. Logan commenced a Ph.D. project at Shark Bay in 1956, and included the Hamelin Pool stromatolites in this study. They were described in his thesis (Logan, 1959) and in a subsequent publication (Logan, 1961). More comprehensive studies of the stromatolites have since been carried out by P. Hoffman, B.W. Logan, and C.D. Gebelein, and their results are summarized by Hoffman (Ch. 6.1). Details have.been published by Logan et al. (1974).

The Quaternary and modem environments of Shark Bay are described in the memoir by Logan et al. (1970a). This includes details of some aspects of the Hamelin Pool environment and sedimentation, and has comprehensive papers by Davies (1970% b) on seagrass banks and algal-laminated sediments.

Study of the Hamelin Pool stromatolites by the Geological Survey of Western Australia began in 1968, and is still in progress. The occurrence of subtidal stromatolites was briefly reported by Playford (1973). Future

390 P.E. PLAYFORD AND A.E. COCKBAIN

I

Bernier Island

S H A R K BAY

250

Dorre Island

Fig. 1. Locality map, Shark Bay and Hamelin Pool.


detailed work will utilize low-level vertical colour air photos of important areas, which were taken for the Geological Survey in 1974. Regional geological mapping of the whole Shark Bay area will also be carried out. S.M. Awramik and S. Golubic are studying the microbiology of Hamelin

Pool algal mats and stromatolites, with the cooperation of the Geological Survey. Their field work was carried out in 1973, and systematic work on the algae by Golubic is continuing.

A meeting of stromatolite specialists was convened by the Environmental Protection Authority of Western Australia in 1973 in order to make recommendations on the preservation of Hamelin Pool and its unique stromatolites (Western Australia, Department of Environmental Protection, 1975).

THE HAMELIN POOL ENVIRONMENT

Hamelin Pool consists of a broad central basin 5-10m deep surrounded by a sublittoral platform up to 5km wide and intertidal-supratidal platforms from a few metres to 3.5 km wide. Hutchison and Nilemah Embay- ments (Fig. 2) are subdivisions of Hamelin Pool characterized by very broad intertidal - supratidal platforms.

The Hamelin Pool basin is barred to the north by the Faure Sill (Figs. 2 and 4), a seagrass bank (Davies, 1970a) extending from Kopke Point to Petit Point. Most of the sill is less than a metre below sea level, and it is cut by three main tidal-exchange channels, the Herald Loop, Faure, and Petit Channels (Figs. 2 and 4). Because of the restriction to inflow of oceanic waters caused by the Faure Sill, combined with the low annual precipitation (average 210 mm) and high evaporation (average about 2,200 mm), hypersaline conditions prevail in Hamelin Pool. Salinities range from 55'& to 70°/m throughout the year, the highest values being in the southern part. However, the restriction to circulation in the basin is insufficient to cause oxygen depletion in the bottom waters. Minor amounts of gypsum are precipitated on the intertidal- supratidal platforms.

The prevailing wind in Hamelin Pool is overwhelmingly from the south. Winds from this direction are especially strong and persistent during the summer.

There are no detailed tidal records for Hamelin Pool, but data from Logan et al. (1970a), combined with our own observations, indicate that the normal spring tidal amplitude ranges from about 1 m in the central-northem part to 0.6m in the south, and the mean *diurnal amplitude ranges from about 0.5 to 0.3 m. However, the tides are very much influenced by the wind. When a northerly wind is blowing, water banks up in Hamelin Pool and high tides are commonly as much as 0.3m above normal levels. Conversely, persistent southerly winds can lower the water level by similar amounts.


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..:.. .:. . supratidal platform . . . .

Fig. 2. Mappf Hamelin Pool showing the main bathymetric features and localities mentioned in the text.


Fig. 3. Diagrammatic section illustrating distribution of living algal stromatolites in Hamelin Pool. Abbreviations are: M.S.L. = mean sea level; H.W.S. = high-water-spring tide level; L.W.S. = low-water-spring tide level.

The fauna of Hamelin Pool has not been studied in detail, but it is clearly very restricted in species numbers compared with the oceanic and metahdine parts of Shark Bay. However, individual species can be very abundant, especially the small bivalve Fragum hamelini Iredale. “Population explosions” of this species apparently occur in certain years. Other prominent elements of the fauna are peneroplid and miliolid foraminifers, a few species of fish, and sea snakes.

The most remarkable elements of the Hamelin Pool biota are the stromatolites and associated flat algal mats, to be discussed later in this study.

GEOLOGY

The oldest rock unit exposed in the Hamelin Pool area is the Toolonga Calcilutite, of Late Cretaceous age. It consists mainly of chalk (in part with flints) and is dolomitized at some localities.

Eocene and Miocene limestones overlie the Toolonga Calcilutite in the Shark Bay area, but are not exposed around the margins of Hamelin Pool. An orthoquartzite which crops out prominently on headlands on the east side of Hamelin Pool is probably Tertiary in age.

The main exposures in the area are Quate’mary, and these are discussed by Logan et al. (1970b). The oldest unit is the Peron Sandstone (possibly Early Pleistocene to Pliocene), which occurs on Peron and Nanga Peninsulas west of Hamelin Pool. It may be an eolian deposit. The Peron Sandstone is suc- ceeded in coastal sections by thin marine Pleistocene units (Dampier Form- ation and Carbla Oolite) laid down during the Dampier Marine Phase (Logan et al., 1970b). The youngest marine Pleistocene deposit recognized in the area is the Bibra Formation, which contains a rich coral and molluscan fauna.


Fig. 4. Earth Resources Technology Satellite image of Hamelin Pool and adjoining areas. The shallow Faure Sill and its associated tidalexchange channels are clearly shown extending from Petit Point to Kopke Point along the northern margin of Hamelin Pool (see Fig. 2 for localities).


Hamelin Pool is believed to have had a configuration comparable with that of today during the Pleistocene marine transgressions, but circulation was less restricted than at present because the Faure Sill seagrass bank had not then developed. Stromatolites are not known from the Pleistocene deposits.

Red sand dunes (the Nilemah Sands) of possible Late Pleistocene to Holocene age overlie older Pleistocene deposits on Peron and Nanga Penin- sulas. The dunes are now fixed by vegetation.

The Hamelin Pool area was completely emergent during the period of low sea level associated with the Wurm Glaciation. The Pleistocene sequences were eroded during this period, and calcrete crusts developed in some places.

Playford and Chase (1955) suggestedethat coastal landforms in the Shark Bay area may have an underlying structural control. They pointed out that there is a strong north-northwest parallelism between Peron and Edel Land Peninsulas, Dirk Hartog Island, and Kopke Point and Yaringa Point headlands, and suggested that each of these features is probably controlled by folding or faulting below the Quaternary deposits. Wapet tested this hypothesis by drilling seventeen holes on Dirk Hartog Island in 1955-56. This showed that the Pleistocene dune limestone (Tamala Eolianite) forming the island has accumulated on an anticline in Tertiary rocks, having its axis parallel to the length of the island and its culmination in the centre. Similar drilling has not been carried out on the peninsulas and headlands of Shark Bay, but as they are similar in configuration and trend to Dirk Hartog Island it Seems likely that they too overlie and are localized by Tertiary anticlines. It is possible that folding has continued into the Quaternary, as Pleistocene rocks are involved in folding of other anticlines in the coastal part of the Carnarvon Basin (Playford et al., 1975). If this hypothesis is correct, Hamelin Pool marks a synclinal downwarp, while the Hutchison and Gladstone Embayments overlie smaller synclines.

HOLOCENE SEDIMENTARY FACIES

The transgression following the Wurm Glaciation occurred rapidly, and the sea reached a height of some 1.5-2.5m above present sea level in Hamelin Pool about 4,000-5,000 years B.P. (Logan et al., 1970a). It is not certain whether the subsequent emergence that has occurred is eustatic or tectonic in origin, but we believe that it is probably tectonic.

The Faure Sill seagrass bank formed across the northern margin of Hamelin Pool during the Holocene, restricting circulation in the basin, and leading to hypersaline conditions and the appearance of stromatolites.

There are few published details of the sedimentary regime in the subtidal areas of Hamelin Pool (the basin floor and sublittoral platform). Nearly all the detailed work to date has been concentrated on the intertidal and supratidal areas. However, Logan et al. (1970a) state that sediments of the


hypersaline basins of Shark Bay are dominated by coquinas of the bivalve Fragum hamelini, with species of peneroplid and miliolid foraminifers, associated with the green alga Acetabularia. Fragum coquinas and skeletal- fragment sands are widespread on the sublittoral platform, and there are also ooid shoals.

Logan and Cebulski (1970) record that in the hypersaline sublittoral sand- flat environment “at certain times an incoherent layer of blue-green algal material coats all surfaces, but at other times it is absent”. However, our observations show that well-developed algal mats and stromatolites persistently cover large areas of the sublittoral platform in Hamelin Pool.

The seagrass banks of Shark Bay are comprehensively described by Davies (1970a). The seagrasses Posidonia and Cymodocea trap and stabilize sedimentary material in these banks and provide the habitat for rich sediment- forming communities, dominated by foraminifers, algae, and molluscs.

The intertidal and supratidal platforms of Hamelin Pool are generally wide in bays and narrow on headlands. The largest expanses are in the Hutchison and Nilemah Embayments, where broad supratidal areas have resulted from the Late Holocene regression. The intertidal zone is characterized everywhere by flat algal mats and stromatolites. Algal mats also extend over wide areas of the lower supratidal platforms.

White sand and coquina beaches occur almost continuously around the shores of Hamelin Pool. The principal constituents of the sand are ooids, pellets, foraminifers, and shell particles, while the coquinas are composed almost entirely of Fragum hamelini shells. The beach deposits are cross- bedded, and strongly cemented beach rock occurs at some localities. The beaches are actively prograding over wide areas, leaving successive beach ridges extending inland for tens or hundreds of metres.

ALGAL STROMATOLITES

Introduction

The stromatolites in Hamelin Pool are the best examples known in modem seas of algal stromatolites having morphological diversity and abundance comparable to that of Proterozoic and Early Palaeozoic stromatolites. The stromatolites were first described by Logan (1961), who reached the important conclusions that they are restricted to the intertidal zone and lithify as a result of subaerial exposure. As a corollary he deduced that their maximum height approximates the tidal range. Logan suggested that his conclusions on the origin of Hamelin Pool stromatolites may also be applied to ancient stromatolites, and this view was further expanded by Logan et al. (1964).

For some years most stromatolite workers accepted that marine


stromatolites are strictly intertidal phenomena. However, in recent years there has been increasing evidence presented that many ancient and modem stromatolites are subtidal, and Monty (1973a) has shown that bacterial stromatolites are forming today in abyssal water depths (see also discussion in Playford et al., Ch. 10.4). Even so, as pointed out recently by Serebryakov and Semikhatov (1974), many authorities have continued to believe that stromatolites are indicative of intertidal conditions.

Our work has now shown that the original basis for the Hamelin Pool intertidal model of stromatolite genesis is inaccurate, as living stromatolites at Hamelin Pool are widely developed in subtidal as well as intertidal environments.

The main areas we have studied at Hamelin Pool are in the vicinity of Flagpole Landing, Carbla Point, and Booldah Well (Fig. 2). Vertical colour air photos have been taken of each of these areas at a scale of 1:5,000 and will be used in our future studies. Different types of stromatolites and algal mats can be readily distinguished and mapped on these photos, and the wide distribution of the subtidal forms is clearly. shown. Some striking lineations and patterns can also be seen in the subtidal stromatolites.

Distribution

Algal stromatolites and associated flat algal-mat sheets are widespread in the intertidal and inner sublittoral platform areas of Hamelin Pool. They are best developed in the central and southern (more saline) areas, south of Yaringa Point.

In some areas (such as around Carbla Point) the principal intertidal stromatolite occurrences are on rocky headlands formed of Pleistocene or older rocks, while extensive flat algal mats occur mainly in the intervening bays. However, in other areas there are well-developed intertidal stromatolites along straight stretches of coastline lacking prominent headlands (such as the area south of Flagpole Landing).

Subtidal stromatolites and flat algal mats occur over wide areas in front of both bays and headlands.

Living intertidal stromatolites extend to about high-water-spring tide level (Fig. 3 on p. 393). However, in most areas the living stromatolites are backed by dead forms in the supratidal zone, and these commonly reach 0.5-1 m above mean sea level. The dead stromatolites are in various stages of disintegration. At some localities (such as 2.5 km south of Booldah Well) the dead forms occur in distinct levels, suggesting periodic emergence, but in others they simply slope gradually downwards to join living intertidal stromatolites (Fig. 5A). In some stromatolites that are only partly emergent, the exposed crests may be hard, cracked, and exfoliating, while the sides (which are still regularly submerged) have actively accreting living mats (Fig. 5B). The colour of the dead stromatolite surfaces changes progressively from grey


where barely emergent, to orange, to pale greyish-yellow in the highest examples. The dead forms are coated with a thin crystalline aragonite crust which is apparently an inorganic precipitate from sea spray. The grey crusts in the lower emergent forms are covered with films of coccoid epilithic and endolithic blue-green algae (film mat of Hoffman, Ch. 6.1) which is not participating in further accretion and is apparently aiding destruction of the stromatolites (S. Golubic, written communication, 1974).

We believe that the emergence of these older stromatolites is most likely to be a result of periodic uplift in Holocene times.

In some areas there are also inactive intertidal stromatolites, even though they are regularly submerged and appear to have grown when the sea was about at its present level. These forms are generally well lithified and are covered with films of epilithic and endolithic algae. It is not always clear why growth of these stromatolites has ceased, but in many cases it can be shown that they have been “smothered” by beach deposits, mud, or sand shoals, which have later been swept away. Some are subsequently reactivated as they become progressively covered with accretional algal mat.

Most subtidal stromatolites are active today, being covered with soft sediment bound by algal mat. However, some are hard and are no longer growing. These have commonly been covered and later exhumed by moving sand or coquina. Underwater photos of subtidal stromatolites taken in different years show that the amount of individual columns protruding above the bottom sediment can vary considerably (Fig. 8). In such cases active accretion is continuing on the tops of the columns, while the lower parts (which are periodically buried) are inactive.

The idealized form zonation of stromatolites shown by Logan (1961) does not represent the typical situation at Hamelin Pool. In most areas there is a general seaward slope of the upper surfaces of living stromatolites, and they pass below sea level to join the subtidal forms (Fig. 3, 5A). In many places there is an area of mobile sand with only scattered subtidal columnar stromatolites between the main intertidal and subtidal occurrences. The living subtidal forms reach depths of at least 3.5 m below sea level, and i t is possible that more detailed examination of the sublittoral platform will show that there are some even deeper examples, conceivably reaching as deep as 5m.

Fig. 5. A. Columnar stromatolites at Carbla Point photographed a t low tide. Those in the foreground are now dead and emergent and are in an advanced stage of disintegration. The elevation of the tops of the stromatolites falls progressively seawards from the dead stromatolites through forms with dead crests and actively accreting sides to fully active intertidal forms (where the man is standing) and subtidal forms (which extend several hundred metres out to sea). B. Confluent club-shaped stromatolites (at low tide) 2.8 km southwest of Flagpole Landing. The crests of these stromatolites are now slightly above normal high-water-spring tide level (owing to recent emergence) and are hard, cracked, and exfoliating whereas the sides are still covered with actively accreting mat. The beach on the left is a Fragum hamelini coquina, which is advancing progressively over the stromatolites.


Subtidal stromatolites commonly extend in patches for several hundred metres offshore, and they can be clearly seen from the beach as darker parts of the sea floor (as for example at Carbla Point).

Stromatolites in Hamelin Pool commonly coalesce to form linear rigid bodies which are strongly wave-resistant (Fig. 7B). They were referred to by Wapet geologists and by Logan (1961) as stromatolite reefs. Davies (1970a) discussed reef and bank terminology in relation to his studies at Shark Bay. He followed the usage of Nelson and others (1962) and many other modem authors in restricting the term reef to structures having rigid skeletal frameworks. He did not discuss the problem of the Hamelin Pool stromatolite masses, but referred to them simply as “stromatolite ‘reefs”’, even though they lack skeletal frameworks.

We believe that much of the nomenclatural argument about “when is a ‘reef’ really a reef?” has resulted from a lack of appreciation of the fact that stromatolites and other cryptalgal bodies, when subject to early cementation, can form raised, rigid wave-resistant bodies, as is the case at Hamelin Pool in both intertidal and subtidal environments. There seems to be no good reason why these bodies should not be termed reefs. Many Proterozoic and Palaeo- zoic reefs were apparently built of stromatolites in this way.

Forms represented

A wide variety of stromatolite morphologies is represented at Hamelin Pool (see Hoffman, Ch. 6.1 herein). The most characteristic intertidal forms are the club-shaped to cylindrical columns described by Logan (1961). They are well developed in the Flagpole Landing and Carbla Point areas (Fig. 5). Some columns are as tall as 0.75 m, but most are less than 0.5 m tall.

In the Booldah Well area, on the west side of Hamelin Pool, stromatolites are developed as elongate flat-topped ridges 1-3 m wide, about 0.3 m high, and up to 400m long, separated by bare sandy areas (Fig. 6A). They are elongated nearly north-south at an angle of about 30” to the shoreline in this area but parallel to the strong prevailing wind direction. It seems clear that their orientation must be controlled by the wind, but the actual mechanism involved is unknown.

Some of the ridges in the Booldah Well area have subsidiary longitudinal stromatolite structures that extend in the direction of wave translation, i.e.

Fig. 6. A. Stromatolite ridges built by pustular mat near Booldah Well (at low tide). The ridges extend north-south, parallel to the prevailing wind, and at about 30° to the shoreline. The ridges slope seaward, passing below low-water-spring tide level. B. Longitudinal stromatolites near Booldah Well (at low tide). These forms are associated with the system of stromatolite ridges in this area (see A above) which are crudely developed in this view extending away from the observer (from right to left). The longitudinal forms are at an angle of about 50’ to the ridges and they are elongate in the direction of wave translation, perpendicular to the shoreline.


at right angles to the shore (Fig. 6B). Such elongation also occurs elsewhere in the Hamelin Pool area among many intertidal stromatolites. Small columnar stromatolites 1 km south of Carbla Point form lines perpendicular to the shore, in the direction of wave movement, but the stromatolites themselves are inclined south into the prevailing wind (Fig. 7A). Again the means whereby the wind seems to control stromatolite growth is unknown.

Many different stromatolite forms occur in the subtidal environment at Hamelin Pool, from tiny bulbous bodies developed on otherwise flat algal mats, to large complex mounds. Columnar forms are most common, and these range from conical to club-shaped (Fig. 9A). Narrow parallel-sided columns occur in some areas (Fig. 8). They have grown vertically and are apparently geotropic. The sun is to the north throughout the year (Hamelin Pool is situated at about latitude 26OS), but none of the columns have been seen to have phototropic inclinations to the north. Some subtidal stromatolites in water up to 2 m deep are elongate perpendicular to the shoreline in the direction of wave movement (Fig. 9B), while other large masses are elongate parallel to the shore (perhaps controlled by sand megaripples).

The colloform mat which forms most subtidal stromatolites is soft.but coherent at the surface, and lithification of the structures begins a few milli- metkes or centimetres below the surface, increasing progressively downwards. The lower parts of the columns are generally well lithified, and commonly have a “beard” of Acetabularia, which varies in abundance at different times (Fig. 8). Masses of tiny serpulids encrust cavities in the sides of the subtidal stromatolites, and they play a significant role in the construction of the margins. The small bivalve Irus irus (Linnd) also lives in the sides of many subtidal stromatolites.

Hamelin Pool stromatolites range from unlaminated (thrombolites) to finely laminated; most show only crudely developed lamination (Fig. 10). Fenestral fabrics are characteristic in lithified forms (Figs. 10, 11). The fenestrae have apparently developed largely through the decomposition of enclosed algal material, and as a result of algal mat bridging indentations on the stromatolite surfaces.

As pointed out by Hoffman (Ch. 6.1) there are some significant differences between the internal fabrics of stromatolites produced by the various mat types. The most distinct lamination occurs in stromatolites built by smooth and colloform mats, while those built by pustular mat are generally only weakly laminated or unlaminated. The largest fenestrae form from pustular mat, the smallest from smooth mat.

Fig. 7 . A. Stromatolites exposed at low tide 1 km southeast of Carbla Point showing pronounced elongation in the direction of wave translation (bearing 40-50 , perpendicular t o the shoreline). Individual columns are inclined in the direction 180’ (towards the prevailing wind). B. Stromatolite reef 1.7 km southwest of Flagpole Landing. The reef is now largely emergent and the upper stromatolite surfaces are disintegrating.


The Hamelin Pool stromatolites are constructed of sedimentary material that has been trapped and bound by algae and has been cemented by aragonite to form solid structures. The nature of the cementing process has not been studied; it may be wholly inorganic, but it is possible that the algae, or the products of algal decomposition, could also have played a significant role in lithification.

The sedimentary material making up the stromatolites typically consists of pellets of problematical origin (Fig. l l B ) , ooids, foraminifers, Fragum hamelini, quartz sand, and bioclastic sand. Small dolomite rhombs occur in some stromatolites; these may be entirely detrital (derived from soft Cretace- ous dolomite).

Algae

The algae forming the Hamelin Pool stromatolites have been studied recently by C.D. Gebelein, and further detailed work by S. Golubic is in progress.

The stromatolite-forming algal mats are broadly of three types: pustular and smooth mat (mainly intertidal), and colloform mat (subtidal). The dominant algal species in these mats (identified by Gebelein, quoted by Hoffman, Ch. 6.1) are Entophysalis major in pustular mat, Schizothrix helua in smooth mat, and Microcoleus tennerrimus in colloform mat. A number of other algal assemblages form flat algal-mat sheets; these are termed blister, tufted, and gelatinous mats, but they do not form stromatolites. They are discussed further by Hoffman (Ch. 6.1).

The subtidal mats are dominated by eucaryotic algae, and the microflora has a much higher diversity than that in the intertidal mats (S. Golubic, written communication, 1975).

The Hamelin Pool stromatolites can form an excellent basis for studying principles of stromatolite classification, as the morphology and biology of diverse living types can be studied through a range of environments. From work carried out to date it is clear that morphological characteristics of these stromatolites are controlled partly by the constituent algae and partly by environmental factors. I t would be useful to apply defined names to the various stromatolite types represented at Hamelin Pool, just as it is useful to give names to stromatolites in the fossil record. However, further detailed study

Fig. 8 . A. Subtidal columnar stromatolites 180 m offshore, 2.8 km southwest of Flagpole Landing, photographed at low tide in February 1972 in water about 1 m deep. The maximum height of columns in the foreground is about 0 .4m. The stromatolite crests are covered with living colloform mat, while the sides are coated with a dense “beard” of Ace tabu laria. B. The same group of subtidal stromatolites photographed in November 1974. Note that many of the stromatolites visible in the background in the 1972 photograph were partly or wholly covered in 1974, owing to movement of sand megaripples. In addition, the Acetabularia “beard” was considerably less dense than in 1972.


and description of the stromatolites and stromatolite-building algae represented at Hamelin Pool (especially those in the subtidal environment) will be required before a satisfactory nomenclatural system can be developed, whether it be Linnean or purely descriptive. The descriptive classification proposed by Logan et al. (1964) is inadequate to describe the wide range of morphological characteristics present in both modem and ancient stromatolites; moreover, it does not take account of internal fabric or of the algae in modem forms.

Growth rates

The growth rates of various intertidal and subtidal stromatolites have been studied in our investigation, by placing non-corrosive nails in them and examining them periodically over a five-year period. Photographs taken of particular stromatolites over an interval of seventeen years were also compared, and tracks cut many years ago through stromatolites (by camel- and horsedrawn wagons that were once used to .load wool and sandalwood onto lighters) were examined.

The conclusion reached is that growth of the Hamelin Pool stromatolites is very slow, and many of the living stromatolites are likely to be hundreds of years old. The maximum observed rate of growth amounted to about 1 mm per year (over a period of five years) on one of the intertidal stromatolites, but in most cases there has been almost no observable growth, and net erosion has occurred in some. Many stromatolites seem to have essentially reached a state of equilibrium, with growth approximately balanced by erosion.

Little or no regrowth has occurred over the paths cleared by wagons, even though they have not been used for decades. One of the most conspicuous of these paths cuts through the living stromatolite ridges opposite Booldah Well. It was last used during the early 1930’s to load sandalwood, but the path is still sharply defined and is clear of stromatolites.

Stroma tolites and Un i f 0 rmi tarian ism

The modem stromatolites at Hamelin Pool are important as guides to some aspects of the origin and environmental significance of ancient stromatolites. However, it is clearly false to assume that all stromatolites in the

Fig. 9 . A. Club-shaped, cylindrical, and conical subtidal stromatolites 180 m offshore, 2.8 km southwest of Flagpole Landing, photographed at low tide in November 1974, in water about l m deep. Note the living colloform mat on the crowns of the stromatolites and the sparse “beard” Acetabularia on the sides. B. Low mound-shaped subtidal stromatolites, 200 m offshore in water 2 m deep off Carbla Point, showing elongation of mounds in the direction of wave movement. The mounds are built by colloform mat and are surrounded by sparse Acetabularia.


Fig. 10. A. Section through an intertidal stromatolite built by pustular mat, from Flagpole Landing, showing crudely developed lamination and coarse fenestrae. B. Section through subtidal stromatolite built by colloform mat collected from water about 1 m deep, 190 m offshore, 2.8 km southwest of Flagpole Landing. The stromatolite has grown on a cobble of indurated Cretaceous chalk. Note that lamination is better developed and fenestae are generally smaller than in the stromatolite shown in photo A.

past have formed under similar environmental conditions to those in Hamelin Pool.

Stromatolites were widely distributed in Proterozoic seas, and were still abundant in the Early Palaeozoic, but since then they have steadily declined, until today Hamelin Pool is apparently the only place where diverse and abundant marine algal stromatolites are still forming. The reason for this is


t o be found primarily in the unusual conditions prevailing in Hamelin Pool which have led to development of hypersaline conditions with a restricted metazoan fauna. Although the fauna has yet to be studied in detail, it seems that organisms that could extensively graze and thereby destroy stromatolites and algal mats are rare. In particular, living gastropods are uncommon in Hamelin Pool, but are abundant in the oceanic and metahaline parts of Shark Bay. As suggested by Garrett (1970b), the decline of stromatolites through the Phanerozoic appears to be linked, at least partly, to the progressive rise of algae-consuming metazoans. Hamelin Pool, because of its hypersalinity, is now one of the few remaining places in world seas where metazoan grazers are sufficiently rare to allow stromatolites to flourish.

Hypersalinity thus appears to be the factor of prime importance to development of stromatolites at Hamelin Pool, and although such abnormally high salinities can be expected to have been associated with some stromatolites in the past, there is no reason to believe that such an association was necessary for all stromatolites, especially during the Palaeozoic and Protero- zoic. Sedimentary environments closely comparable with that of Hamelin Pool are probably only rare features of the ancient record, just as they are today.

There are lessons to be'learned from the Hamelin Pool stromatolite story regarding the use of modem models to interpret ancient rocks. It shows the dangers of accepting incompletely known modern environments as definitive guides to the past. The nature of ancient environments should be deduced from evidence in the ancient rocks themselves, without the mandatory need for modem analogues, especially as ancient environments need have no modem counterparts. Moreover, it demonstrates that the past can be as much a key to the present as the present is to the past; recognition of subtidal stromatolites in ancient rocks led to reevaluation of the picture at Hamelin Pool and to the realization that subtidal stromatolites are widely distributed there also.


Hamelin Pool is a hypersaline marine barred basin, which is probably localized by a Cainozoic synclinal downwarp. I t is barred to the north by the Faure Sill, a seagrass bank which developed in Holocene times. The fauna of Hamelin Pool is very restricted in species diversity because of the hypersaline conditions, and algae-consuming organisms are largely absent. As a result, algal stromatolites and flat algal-mat sheets flourish on the sublittoral and intertidal platforms.

Living stromatolites in Hamelin Pool extend from depths of at least 3.5 m below sea level to about high-water-spring tide level. Older dead stromatolites in varying states of disintegration occur above this level, and these have


probably emerged as a result of Holocene uplift, which may still be continu-

Lithification of stromatolites is taking place in both intertidal and subtidal environments, and coalescent stromatolites form raised wave-resistant reefs in many areas. Lamination is only crudely developed or absent in many Hamelin Pool stromatolites, and fenestral fabrics are characteristic.

Growth of stromatolites at Hamelin Pool is very slow, and many living forms are probably hundreds of years old. The maximum observed growth rate of an individual stromatolite is about 1 mmlyear, but the average seems to be much less. Many forms appear to have almost reached an equilibrium, with upward growth balanced by erosion of the living mat.

A considerable amount of additional work is necessary before the Hamelin Pool stromatolites can be fully described and classified. It is desirable that a comprehensive system of nomenclature be developed for them which will also have application in the classification of ancient stromatolites.

ing.

ACKNOWLEDGEMENTS

We wish to thank the following geologists for their assistance with our field work at Hamelin Pool: R.L Chase, R.W.A. Crowe, W.J.E. van de Graaff, D. Johnstone, and D.C. Lowry. S.M. Awramik and S . Golubic gave valuable advice and provided preliminary results of their investigations of the stromatolites, and useful discussions were held with P. Hoffman on aspects of his work. G.W. Kendrick kindly identified molluscs from Hamelin Pool.

Fig. 11. A. Thin section of a well-lithified intertidal stromatolite from Flagpole Landing showing characteristic fenestral fabric in weakly laminated pellet wackestone. Note scattered shelly fragments. B. Thin section of the same stromatolite shown in photomicrograph A. This is a fenestral pellet wackestone, with scattered ooids, shell fragments, foraminifers, quartz sand, and dolomite rhombs (?detrital). Fenestrae are lined with acicular aragonite.


Chapter 8.3

STRATIGRAPHY OF STROMATOLITE OCCURRENCES IN CARBONATE LAKES OF THE COORONG LAGOON AREA, SOUTH AUSTRALIA

Christopher C. uon der Borch

INTRODUCTION

Geologically Recent stromatolites have been described from several areas of modern carbonate sedimentation, the most notable of which is Shark Bay in Western Australia (Logan, 1961; Davies, 1970b). Other important areas include the Persian Gulf (Kendall and Skipwith, 1968; Kinsman et al., 1971), the Bahamas (Monty, 1967,1972), Bermuda (Gebelein, 1969) and the Great Salt Lake, Utah (Carozzi, 1962). In addition to the above, a somewhat restricted but nevertheless interesting occurrence of stromatolites has been described from two ephemeral carbonate lakes associated with the Coorong Lagoon in South Australia (Walter et al., 1973). These stromatolites occur in carbonate muds composed of the minerals hydromagnesite, aragonite, dolomite and calcite (Alderman and Von der Borch, 1960, 1961; Skinner, 1963; Peterson and Von der Borch, 1965; Von der Borch, 1965). Small amounts of amorphous silica are associated with these carbonates. This assemblage has its analogues in the geologic record, particularly in the Precambrian (e.g. the Skillogalee Dolomite Formation of South Australia, Preiss, 1973a) where silicified dolomitic stromatolite-bearing rocks are relatively common. Because of this, as an aid in recognizing equivalent depositional environments in the ancient rock record, the sedimentological history and stratigraphy of the Coorong stromatolite association will be described in some detail.

GEOLOGICAL SETTING

The Coorong Lagoon (Fig. 1) is the dominant physiographic feature in the area of the stromatolite lakes in question. I t is located landward of a modern calcareous barrier island known as Younghusband Peninsula and is approximately 100 km in length and 3 km in width. There is a single marine pass at the northwestern extremity, where water depth in the lagoon reaches 10 m,

414 C.C. VON DER BORCH

Fig. 1. Locality map, showing Coorong Lagoon and ephemeral carbonate lakes. The localities of the detailed maps in Fig. 2 are indicated.

however water depth throughout most of the area averages only 2-3 m. Water salinity is generally higher than that of normal seawater, reaching values of 60 parts per thousand in southeastern areas.

The Coorong Lagoon is the most recent of several that formed across a broad coastal plain during the Pleistocene. Several stranded barrier islands, now represented by calcreted calcareous eolianite ridges, occur subparallel to the present coastline and extend inland up to 65 km from the coast. These and their associated lagoonal deposits have been progressively stranded by a combination of glacially induced sea-level oscillation and gentle regional upwarping (Hossfeld, 1950; Sprigg, 1952). The landward shoreline of the Coorong Lagoon is formed in most areas by one of these Pleistocene calcreted barrier-beach-ridge complexes.

Occasional inliers of a Pleistocene eolianite ridge occur within the present barrier of Ypunghusband Peninsula, particularly in northern areas (Brown, 1965). The Coorong Lagoon itself occupies one and sometimes two inter- dunal depressions between these ridges.

Towards the southeastern extremity of the Coorong Lagoon, in the vicinity of the two stromatolite-bearing lakes (Fig. 2), a rather complex calcreted


I COORONG LAGOON I

A

L

M

E

IY J B

C CALCAREOUS SMDS OF mDEWI BARRIER MD BEACH RIOGES.

$$$$ MICROCRYSTALLINE WLORITE. CALCITE. ARAGONITE AND HYDROMGNESITE SEDIIILNTS OF LAKES AND ELEVATED FLATS. (LAYTS SHOW I N WHITE)

SILICEOUS SAND RlOGE NEAR SALT CREEK,

CALCAREOUS SANDS OF TIDAL DELTA FACIES. 0 S T R W l O L I T E OCCURKNCES

CALCRETED PLEISTOCENE EOLlANlTE ROADS

Fig. 2. Details of the two ephemeral carbonate lakes which contain stromatolites. A. Northern stromatolite lake area. B. Southern stromatolite lake area. C. Generalized cross section X-Y, showing typical stratigraphic relationships. The 120,000 and 80,000 year old barriers are indicated.


eolianite topography exists, due possibly to a combination of Pleistocene dune blow-outs superimposed on the normal beach-ridge-barrier lineation. Low-lying portions of these areas were inundated between 3,000 and 5,000 years ago when the Holocene transgression reached its maximum.

GEOLOGICAL HISTORY

The sedimentological evolution of the Coorong Lagoon and marginal ephemeral lakes is most clearly understood by reference to the sea-level curve of the last 200,000 years of Veeh and Chappell (1970) constructed from a study of Late Pleistocene coral reef terraces in New Guinea (Fig. 3). Assum- ing that the degree of upwarping in the Coorong region has been essentially negligible for that time, it is evident that the last time the sea was near its present level was 120,000 years ago during the Sangamon interglacial. During this stage the presently calcreted beach-ridge-barrier system that forms the landward shoreline of the Coorong Lagoon was probably built (D. Schwebel, pers. comm., 1975). This system forms the basement of the stromatolite lakes shown in Fig. 2. A subsequent sea level approached to within several metres of the present sea level about 80,000 years ago and this may have been responsible for the barrier island which now forms the eolianite inlier within the Younghusband Peninsula barrier. Next in sequence followed the major sea-level retreat of the Late Wisconsin Glaciation after which the Holocene transgression returned the sea to its present level, partly inundating the older barriers to form the initial stages of the modern Coorong Lagoon.

mi I I I I 0 40 80 120 MO 200

YEARS 10’

Fig. 3. Sea-level curve of the last 200,000 years, based on New Guinea data of Veeh and Chappell(l970). Relevant dates are as follows: A = high sea level of 120,000 years ago corresponding to Sangamon interglacial: formation of older barrier that now forms most of the Pleistocene basement in the area: B = relatively high sea level of 80,000 years ago: formation of eolianite inlier within modern barrier. C = Holocene transgression: formation of modern barrier.

At this early stage the lagoon water was about l m above its present level as evidenced by widespread stranded lagoonal sediments (Brown, 1965). It is uncertain whether this higher level was due to a sea level slightly higher than at present or was related to unrestricted access of the ocean through numerous passes. However, this was the time during which low-lying areas of the


Pleistocene topography marginal to the lagoon, including the stromatolite lakes, were flooded by waters with oceanic affinities.

As sea level stabilized in the Late Holocene, the Younghusband Peninsula barrier built over the older ridge causing the lagoon to become increasingly restricted until ultimately its southern waters reached their present hypersaline state. Lagoon waters at that time, possibly 3,000 years ago, receded l m to their present level which averages that of mean sea level, causing the stranding of the marginal stromatolite lakes. These lakes now have floors about 1 m above present high lagoon level and receive their annual charge of water during winter months by local and regional groundwater seepage from surrounding calcareous eolianite ridges and deeper Tertiary aquifers. This water reaches a maximum depth of about 0.5 m in late winter and evaporates to dryness in early summer. The salts in the lake waters are derived in part by leaching from marginal flats, possibly augmented by magnesium ions from breakdown of high-magnesian calcite in the calcareous aquifers. It is in this environment that various carbonate minerals and the stromatolites are forming at the present day.

STROMATOLITE DESCRIPTIONS AND STRATIGRAPHY

The Coorong stromatolites have been described in detail by Walter et al. (1973). Three morphological forms were described: globular, stratiform and crenulate (Fig. 4). They occur in marginal areas of the two shallow ephemeral lakes shown in Fig. 2 and are associated with a carbonate mineralogy of hydromagnesite, aragonite, dolomite and calcite. During winter and early spring months pH of the associated lake waters varies from 8.2 to 9.9 (Von der Borch, 1965) with salinities reaching a minimum of about loo/, . At this stage a sparse growth of the aquatic grass Ruppia maritima occurs on lake floors. Complete desiccation generally occurs during ensuing arid summer months. The actual stromatolite structures persist throughout the dry period with varying degrees of preservation, however during the wet season small myriapod arthropods burrow into them and destroy some of their features.

Sediment cores taken through the stromatolite lakes to depths of several metres (Fig. 5) show a stratigraphy that reflects the evolutionary history of the Coorong Lagoon and its associated marginal lakes. In both cases basement is Pleistocene calcrete, which is developed on calcareous eolianite. Overlying this is a shallow-marine sedimentary unit (Protected Marine Phase, Fig. 5), consisting of skeletal grainstones and packstones containing a variety of bivalves such as Katelysh spp. and Venerupis sp. This was formed during and immediately following the Holocene transgression, before significant build-up occurred of the Y ounghusband Peninsula barrier. This grades upwards into pelletized organic-rich aragonite and magnesian calcite muds rich in a lagoonal fauna comprising the small gastropod Coxiella confusa, the


Fig. 4. Coorong stromatolite morphologies. A. Globular stromatolites, southern stromatolite lake (coin 23 mm wide). B. Desiccation polygons in stratiform stromatolites, southern stromatolite lake (scale 30 cm long). C. Crenulate stromatolites on the margins of desiccation polygons of stratiform stromatolites, southern hydromagnesite lake (pencil 18 cm long).

foraminifera Ammonia beccarii and a variety of ostracods (Lagoonal Phase, Fig. 5). This change in sediment-type reflects a period of increasing restriction due to barrier development. Upper portions of cores, finally, are composed of white, microcrystalline carbonate muds and pellet packstones composed of the minerals hydromagnesite, calcite, aragonite and dolomite (Ephemeral Lake Phase, Fig. 5). This unit, which developed in response to the l m drop in lagoon level, contains stromatolite structures both at the surface and occasionally at depth.

GEOLOGICAL SIGNIFICANCE

Stromatolites are rare in modem carbonate sediments, largely because of the browsing on algal mats by organisms such as crustacea. The fact that stromatolites occur in two lakes in the Coorong area is therefore somewhat sur-


NORTHERN STROM. LAKE -

0

1 -

2 - vl Y

t

2

3 -

419

SOUTHERN STROM. LAKE

NANTLY HYDROMAGNESITE, ARAGONITE,

SKELETAL GRAINSTONE AND PACKSTONE W I T H B I V A L V E S K A T E L Y S I A AND V E N E R u P I S .

W I T H

.......J

PLEISTOCENE CALCRETE

PLEISTOCENE CALCAREOUS E O L l A N l T E

Fig. 5. Generalized stratigraphic columns from northern and southern stromatolite lakes. Terminology partly adapted from Brown (1965).

prising, particularly in view of their general absence in apparently comparable lakes nearby. One contributing factor may be that the two stromatolite lakes, due to greater localized ground-water influx, do not become as thoroughly desiccated as other lakes. This higher ground-water discharge in turn may lead t o the formation of the carbonate mineral hydromagnesite which is associated with the stromatolites.

It would be quite possible for sediments formed under conditions similar to the above to be preserved in the geologic record, particularly in structurally negative areas. Such occurrences have been described in the literature (Peterson, 1962). The deposits would be typified by an association of carbonate cycles similar to the one described from the Coorong sediment cores. A single upward-fining cycle would ideally measure a few metres in thickness, beginning at the base with a shallow marine carbonate unit. This would grade upwards into an organic-rich pelletized lagoonal carbonate unit and would culminate in a siliceous microcrystalline carbonate with associated stromatolite structures. Mineralogy of the upper unit would be variable in both the lateral and vertical sense and could comprise combinations of the minerals magnesite, dolomite or calcite. The carbonate units themselves would be lenticular and laterally discontinuous and would be closely associated with shoestrings of calcareous barrier sands.


ACKNOWLEDGEMENTS

Sediment coring in the Coorong area has been supported by grants from the Australian Research Grants Committee and Flinders University. The manuscript was read by D. Schwebel. Photographs of the Coorong stromatolites were provided by M.R. Walter.


Chapter 8.4

ALGAL BELT AND COASTAL SABKHA EVOLUTION, TRUCIAL COAST, PERSIAN GULF

David J.J. Kinsman and Robert K . Park

INTRODUCTION

The distribution of sabhkas in the Persian Gulf is shown in Fig. 1 and the area of detailed study in Fig. 2. The first studies of the Trucial Coast stromatolitic sediments were made by Kinsman (1964) and these were later expanded by Kendall and Skipwith (1968) who recognized a zonation of forms throughout the intertidal zone. Our very detailed work on the algal belt near Abu Dhabi (Park, 1973) has added considerably to these earlier studies.

ALGAL-BELT ENVIRONMENT

A review of sedimentation along the Trucial Coast has been presented by Purser and Evans (1973). A general view of the lagoon immediately to the northeast of Abu Dhabi Island is shown as Fig. 3. The lagoon barriers comprise oolite and reef facies sediments; within the lagoons sediments are mainly pelleted aragonite muds but shoals downwind of extensive open lagoon areas may comprise winnowed accumulations of skeletal sands and gravels (predominantly molluscan debris).

Algal mats colonise relatively stable sediment substrates in the intertidal zone and the algal belt may be as much as 1-2 km in width. High-energy intertidal zones where the substrate is mobile are not colonised to the extent that a permanent mat is established (Fig. 4). Low-energy intertidal zones along protected lagoon margins are typically colonised by algal mats; surface slopes are very low, and spring tidal range usually close to 120cm (profiles Area I and IV in Fig. 4).

The physical, environmental variables can be summarised as follows: (a) predominant wind and wave direction from the NW; (b) tidal range diminishes from 210cm at the lagoon barrier to 120cm in the innermost lagoon (tidal range is one of the important controls on the thickness of stromatolitic

422 D.J.J. KINSMAN AND R.K. PARK

A R A B I A

OCEAN

Fig. 1. Arabian Peninsula showing extent of areas < 200 m above sea level and distribution of coastal and continental sabkhas. Main study area is near Abu Dhabi.

sections which are developed); (c) air temperatures range seasonally 15-47"C, lagoon waters 15-4OoC but exposed algal-mat surfaces 15 to approximately 50°C; (d) lagoon water salinities in areas of algal-mat development are 42-70°/oo but pools within the algal belt may reach halite saturation for short periods; (e) annual evaporation rates are approximately 150cm per year; (f ) rainfall is sporadic but averages about 4-6 cm per year.

ALGAL-MAT MORPHOLOGIES

In the subtidal environments to depths of 3 m occur occasional gelatinous, domal forms, up to 6 cm high and 18 cm in diameter. They display millimeter- scale layering and are dominated by Phormidium hendersonii. Also in some very shallow subtidal and lower intertidal areas are gelatinous coccoid algal patches which are not cohesive, not layered and are destroyed at the merest handling; these form ephemeral mats up to 20 cm in diameter and are dominated by Aphanothece sp.

Our studies have shown that several different algal-mat morphologies can be recognised and that these occupy certain positions within the subtidal and intertidal zones (Fig. 5). A major control on the distribution of the different forms is the frequency of wetting (flooding' frequency) by tidal or storm- driven lagoon waters (Fig. 6).

M z

U n

Fig. 2. Physiographic and sedimentary environments extending from the Oman Mountains in the east to the Trucial Coast of the Persian Gulf, near the town of Abu Dhabi, in the west. Infilling of coastal lagoons behind a complex lagoon barrier has developed extensive supratidal coastal s a b e a s over the past 5,000 years. Note the extensive development of a stromatolitic algal belt (black) dong many lagoonal intertidal areas.

k 0


Fig. 3. Oblique aerial view looking NW, seaward along axis of detailed study area outlined in Fig. 2. Sabkha in foreground; intertidal algal belt is dark zone; seaward again is the broad lagoon, with lagoonal barrier islands and open Persian Gulf shelf environment in far distance. Lagoon approximately 20 km broad; algal belt in foreground 1-2 km wide.

The major mat morphologies which comprise the algal belt proper are as follows : (1) Pustular Mat. This form is equivalent to the Pustular Mat of Logan et

al. (1974) from Shark Bay and to the Cinder Mat of Kendall and Skipwith (1968). It is commonly the initial coloniser in mid and lower intertidal areas between + 15 and +70 cm above our datum of low-water-spring tide level (Fig. 5); in the GD area (located in Fig. 9), high creek-water salinities lead to absence of grazing gastropods and Pustular Mat here occurs at - 65 cm. This form is dominated by the coccoid alga Entophysalis major. Sediment fabrics resulting from this form are clotted mud with little “memory” of the algal precursor; little organic peaty fraction is preserved. Thus, this form has very low preservation potential.

(2) Smooth Mat. This form is equivalent to the Smooth Mat of Logan et al. (1974) and to most of the Polygon Mat of Kendall and Skipwith (1968). It extends generally from + 15 to + 90 cm but in high-salinity areas occurs subtidally down to -40 cm. Smooth Mat is best developed in pool and channel sites and covers 30-60% of the typical algal-belt surface. The mat community is dominated by the filamentous alga Microcoleus chthonoplastes. This mat form is commonly rich in carbonate sediment grains, which are present either


Fig. 4. Tidal range and slope of intertidal zones of various degrees of exposure. Oolite beach of barrier island is steep, the sediment very mobile and tide range over 200cm. Downwind lagoon coast (Area 111) is moderately steep; sediments are sandy and too mobile for algal colonization. Algal colonization of low-energy intertidal areas with low slopes and more muddy sediments is illustrated by Areas I and IV, (tidal range 120 cm). Profile IV runs north-south across the broadest part of the algal-belt embayment in the center of Fig. 3.; station G lies seaward of the algal belt and station F lies o n the sabkha to landward of the algal belt.

as sediment-dominated layers or as dispersed grains in an organic-rich matrix (Fig. 7). In the seaward Smooth Mat, the stromatolitic section is sediment- dominated with thick sediment layers (up to l c m ) and thin organic layers (<0.5mm); as upward growth continues this ratio changes and the uppermost parts of some Smooth Mat sections are algal peats with minor sediment content.

Smooth Mat is commonly broken into polygons giving rise to structures shown in Fig. 8. Smooth Mat has a high preservation potential and is the most important mat form in terms of producing recognisable stromatolitic sediments. (3) Pinnacle Mat. This form is equivalent to the Tufted Mat of Logan et al.

(1974). It occurs over a 40 cm range from + 60 to + 100 cm. This form covers a fairly large fraction of most algal areas (30-40%) and the pinnacles themselves are predominantly formed by the large filamentous alga Lyngbya aestuarii. Occasional large pinnacles become sediment-layered and preservable as sedimentological entities, however, most of the Pinnacle Mat gives rise to no recognisable sediment fabric. This form commonly grows on higher, well- drained areas between pools and the sediment section is of clotted mud, well oxidised and burrowed, in contrast to the well-layered anoxic sections of Smooth Mat, beneath the floors of adjacent pools.

'zol .

M SL -

FD - AREA L f V E U l b V m n l D U E O

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4,. . . ..

.. DATA COLLECTED FROM 3 AREAS OF ALGAL MATS

WHERE DETAILED LEMLLING CARRIED OUT

1.j - +-- I .A

F I '

. . Fig. 5. Range of algal-mat forms in Trucial Coast with respect t o mean sea level (M.S.L.): tidal range 120 cm. Areas T, C and GD located in Fig. 9. Abbreviations: Sa = sabkha; F = flat mat; W = wrinkled mat; B = blister mat; Pi = pinnacle mat; P = pustular 2 s mat; S = smooth mat; Cps = cerithid pellet sand; Po = polygons. Areas T and C have algal mat confined to intertidal zone. The 7;1 more saline GD area has subtidal algal mats, due to lack of normal burrowing and grazing biota. In the right hand part of the figure the total vertical range of each algal-mat form is indicated by the solid line to the left; the solid part of the right-hand line for each form indicates the vertical range over which the form covers more than 30% of the substrate surface.

.


Fig. 6. Sea-water flooding frequency of intertidal zone and seaward sabkha areas. Tidal flooding represents flooding by pure tidal component; onshore winds increase the flooding frequency for any location, as also does the presence of depressions within the intertidal zone. Flooding frequency of seaward sabkha zone indicated as number of flood episodes per year. Note relationship of algal forms and evaporite mineral distribution to flooding frequency. Profile extends from station G, seaward of the algal belt (see Fig. 4 caption), via A to F which is located on the sabkha landward of the algal belt. Stations B, R99 and LK6 lie progressively further landward on the sabkha. Algal-mat forms abbreviated as in Fig. 5.

(4) Blister M a t . This form is analogous to the Blister Mat of Logan et al. (1974) and to the Crinkle Zone of Kendall and Skipwith (1968). It occurs at +80 to + 100 cm, has a leathery, domed and blistered appearance and has usually a low carbonate sediment content. It is often broken into polygons and the blisters are usually infilled with discoidal gypsum crystals. In terms of preservation and giving rise to an identifiable stromatolitic structure, this form is of minor importance.

(5) Wrinkle M a t . This form occurs between +90 and + 120 cm, is typically a thin, surficial layer only and forms no recognisable stromatolitic structure which will be preserved. It is usually underlain by a mush of gypsum crystals.

( 6 ) Flat M a t . This poorly developed algal form extends from +110 to + 140 cm and is only flooded at the highest-or storm tides. As a mat form it gives rise to no recognisable structures.


Fig. 7 . Vertical section through Smooth Mat in pool area; pale layers are carbonate sediment; dark layers are organic algal-rich laminae. Scale: white bar is 10 cm.

Fig. 8 . Desiccated, polygonal Smooth Mat; contrast with Fig. 7 . Scale: black bar is 10 cm.


PRESERVATION OF STROMATOLITIC STRUCTURE

429

In spite of a wide variety of surface mat forms, generated by a series of coccoid and filamentous algal communities, in response to flooding frequency, metazoan grazing and other variables, the stromatolitic fabrics and textures preserved for periods longer than 102-103 years are few. In fact, the only frequently preserved form is the Smooth Mat in its simple or polygon varieties (Figs. 7 and 8). The great variety of pinnacles, pimples, blisters, wrinkles, domes, etc., which are such a striking feature of the surface algal belt, are essentially rarely or never preserved as sedimentary structures. The precipitation of gypsum destroys all .algal structures in the upper 20-30 cm of the intertidal zone. Combined with the metazoan grazing control in the lower intertidal zone, this limits most stromatolite sections to about 40- 50 cm in thickness in this area. Only where the lower intertidal and shallow subtidal areas are flooded by very saline (>70°/oo) lagoon waters, do stromatolitic sections occur which are thicker than 50 cm.

Our studies of the algal filaments, cells and pigments have also indicated that very little identifiable material is preserved beyond about 100 years. Occasional sheaths of Lyngbya can be found in 3,000-4,000 ye& B.P. Smooth Mat sediments and even in 7,000-8,000 year B.P. material from the transgressive stromatolite layers (Fig. 10; 14C dating of algal peats and carbonate sediments are described in Kinsman et al., 1976).

SABKHA MORPHOLOGY AND STRATIGRAPHY

The post-Wisconsin transgression flooded across an earlier continental sabkha surface with a seaward slope of about 1:2,000 and around 5,000 years ago (Kinsman et al., 1976) reached a position indicated on Fig. 9 as the Highest Holocene Shoreline. Since that time, carbonate sediment accumulation has resulted in the infilling of the lagoons and the lateral migration of the shoreline. A subsidiary factor in the migration of the shoreline has been a relative fall in sea level of 120 f 10 cm over the past 5,000 years (determined by detailed levelling of sedimentary facies boundaries in core and trench sections). This sedimentary offlap in places has proceeded at an average rate of 2 m per year. A broad supratidal surface has been developed which is underlain by a transgressive-regressive sediment wedge, indicated in Fig. 10. Intertidal sedimentation rates were too slow to keep pace with the rate of transgression and only an attenuated, 10 cm thick transgressive intertidal unit was formed, even though the tidal range was probably 100-200 cm. Subsequent regression has been caused mainly by sediment offlap and under these conditions a regressive upper and lower intertidal unit has been formed, the 120cm thickness of which is equal to the tidal amplitude (see Fig. 10 caption). An isopach map of stromatolitic sediment thickness in the upper, regressive intertidal unit comprises Fig. 9.

4 30 D.J.J. KINSMAN AND R.K. PARK

Fig. 9. Isopach map of Holocene (< 5,000 years B.P.) stromatolitic sediments of the upper regressive intertidal unit beneath the coastal sabkha (in cm), in detailed study area outlined in Fig. 2. The two areas of thickest (> 50 cm) stromatolite facies sediments represent earlier sheltered embayments in which very fine-grained sediments accumulated. High- energy intertidal zone, skeletal sands and gravels accumulated as a split or barrier between these areas and in a northern area (stromatolite facies absent from these areas). Dotted line indicates position of the 5,000-year B.P. shoreline.

Above the upper intertidal facies sediments, whether stromatolitic or otherwise, occurs an aeolian evaporite-bearing unit, shown in Figs. 10 and 11. The detrital fraction of this unit accumulates by adhesion of windblown grains to the moist sabkha surface. Within both intertidal and supratidal units, large-scale, interstitial precipitation of gypsum and anhydrite occurs, from brines which become progressively evaporated within the sediment pore spaces (Kinsman, 1966, 1974). In the sabkha environment dolomitisation of some of the marine subtidal and intertidal sediments, including the stromatolitic facies, also occurs on a fairly large scale (Patterson, 1972; Patterson and Kinsman, 1976; Kinsman and Patterson, 1976).

The intertidal algal-belt sediments can thus be seen to represent the final stage of infilling of the marine-hypersaline lagoons. Sabkha evaporites form in the upper parts of the algal belt and continue to form as the stromatolitic

+LOWER INTERTIDAL-UPPER INTERTIDAL + - - - ~ - -- - - - - COASTAL SABKHA - - - - - - - - - - - - - - + CONTINENTAL SABMA +ALLWIAL+HINTERL CERITHID + PELLETSAND FLAT ALGAL FLAT FAN

Fig. 10. Composite NW-SE stratigraphic section across the Abu Dhabi coastal-continental sabkha belt in the area of detailed study outlined in Fig. 2. (HWST and LWST: high- and low-water-spring tide levels, respectively). The stratigraphic section, from the surface downward, is as follows: 1. Supratidal Facies (0-100 cm): a wedge of detrital sediments thickening inland from the shoreline. Dominated by anhydrite in inland areas: may be magnesite rich at base. 2. Upper Intertidal Facies (60 cm): commonly an algal facies but coarse skeletal sands or gravels are high-energy equivalent; anhydrite may be present in upper part; gypsum often abundant; intensely dolomitized in some areas. Full regressive development. 3. Lower Intertidal Facies (60 cm): cream, muddy, pellet sands with abundant cerithids; large gypsum crystals common; fairly intensively dolomitized in some areas. Full regressive development. 4. Subtidal Facies (0-300 cm): grey-brown, peneroplid foraminifera and lamellibranch-rich muddy sands; thickens seaward. Minor dolomitization. 5. Lower Intertidal Facies (2-5 cm): cream, pellet sands with abundant cerithids. Attenuated transgressive development. 6 . Upper Intertidal Facies (2-5 cm): algal facies in some places, otherwise mixed pellet and detrital sands. Attenuated transgressive development. Unconforrnity : generally fairly sharp: sometimes burrowed. 7 . Grey Aeolian Sands (50-150 cm): grey, cross-bedded aeolian sands: original iron oxide films around grains now present as iron sulphide minerals; some gypsum. 8. Brown Aeolian Sands (thickness unknown, probably about 5-10 m): brown, cross-bedded aeolian sands with some gypsum. Unconforrnity . 9. Miocene Rocks.


Fig. 11. Trench in central sabkha (shovel for scale). Lower intertidal cerithid pellet sands near bailed out water level in pit, overlain by 50cm of Smooth Mat stromatolite-facies sediments containing discoidal gypsum crystals in upper. part; overlain by nodular anhydrite and contorted anhydrite in aeolian supratidal unit. Ruler marks top of section.


sediments become progressively buried beneath the aeolian supratidal unit. The great sheet of stromatolitic sediments is obviously diachronous and records the progressively changing location of the intertidal zone over the past 5,000 years.

ACKNOWLEDGEMENTS

The Trucial Coast research programme, of which these stromatolite studies form one part, was directed by Kinsman and funded by the following agencies to whom grateful acknowledgement is made: American Association of Petroleum Geologists, American Petroleum Institute (API Project 99), National Science Foundation (GA-1170, GA-13489) and the Petroleum Research Fund of the American Chemical Society (PRF 3009-A2). Park completed a Ph.D. thesis on these studies at the University of Reading. Additional fieldwork was carried out by R.J. Patterson, who completed a thesis at Princeton University on aspects of .the dolomitisation process in the sabkha. Detailed study and identification of the algal flora has been completed by S. Golubic; reference should be made to his contributions in the present volume for information on these aspects of the Persian Gulf algal mats. A.G. Fischer reviewed a first draft of this paper.


Chapter 8.5

TEXTURAL VARIATION WITHIN GREAT SALT LAKE ALGAL MOUNDS

Robert B. Halley

INTRODUCTION

Aragonite algal mounds cover over 100km2 of the nearshore areas of Great Salt Lake, Utah. The internal fabric of these mounds is formed by a framework of aragonite precipitates, presumably induced by algae. and characterized by various morphologies, internal sediments, and inorganic cements. Internal fabric variation is believed to reflect the local and periodic variation of brine concentration. The lake, with no outlet, is an evaporative remnant of its giant precursor, Lake Bonneville, and is maintained by three rivers flowing from mountain ranges to the east (Fig. 1). Evaporation coupled with mineral inflow has concentrated the lake water to a brine containing about 275,000 ppm dissolved solids. The proportions of dissolved solids in the brine are similar to those in sea water and hypersaline sea water (Table I).

The post-glacial history of Great Salt Lake is one of continual temporal and areal fluctuations of brine salinity. Natural changes in salinity coincide more or less with fluctuations of lake stage, because most increase in volume is caused by fresh water. Depth varies seasonally by about 0.5 m. During the last one hundred years, the lake level has fluctuated over 6 m with corresponding salinity changes ranging from 14% to 30% (Handy and Hahl, 1966). During the last fifteen years, major variations have been induced by man’s interference with the lake’s natural circulation. Construction of a railroad causeway from Lakeside to Promontory Point (Fig. 1) has greatly restricted the flow of brine between the northwestern arm and the remainder of the lake. Since completion of the causeway in 1957, salinity of the northwestern arm has increased from a “normal” of about 25% to 32.5% (Greer, 1971). Construction of a passenger car causeway between Farmington Bay and the remainder of the lake isolated the southeastern arm into which the Jordan River flows. Greer (1971) reported the salinity of this arm reduced to 8.2%.

436 R.B. HALLEY

Fig. 1. The Great Salt Lake, Utah (after Eardley, 1938).

TABLE I

Proportion of major dissolved solids in sea water, Great Salt Lake brine* and hypersaline sea water from Boca de Virrila (Peru)**

Sea water Great Salt Lake brine Hypersaline sea water

K+ 1.1 1.8-2.6 1.9

Ca2+ 1.2 0.07-0.10 0.10 Na+ 30.6 31.2-26.9 27.2

Mg2+ 3.7 3.5-4.2 6.5

c1- 55.3

HC 0; 0.40 so',- 7.7

52.6-49.7 7.5 -8.8 0.20-0.17

55.1 9 .o 0.16

PH 8.2 7.9-7.6 7.3 Dissolved solids 34,400 190,000-275,000 3 4 5,5 0 0 (PPm)

*After Handy and Hahl (1966); **after Morris and Dickey (1957).


TEXTURAL VARIATION OF GREAT SALT LAKE ALGAL MOUNDS

Within this setting of salinity fluctuations, algal mounds have developed. These were mapped and described by Eardley in 1938 (Fig. 1); most mounds are shown to be circular or oval in plan outline, ranging up to 1.5 m in diameter and 0.5 m high, and are spaced up to 10 m apart. Eardley (1938) and Carozzi (1962) have described various morphologies of these algal structures, including isolated mounds, flat-topped mounds, composite rings, and fes- tooned ridges. The mounds, which extend from the shoreline to a depth of 4m, are normally surrounded by rippled ooid sand. During the summers of 1971 and 1972, samples of algal crust were collected in order to determine the relationships between the living surface and internal structure of those algal mounds to the southeast of Stansbury Island and to the north of Antelope Island.

Fig. 2 . Vertical section through a Great Salt Lake algal precipitate illustrating change in texture (at B ) from unlaminated to laminated. Note the dark laminations (at A ) which appear to be the residue of an older living surface film within the structure.

Variability is the outstanding feature of internal textures of the Great Salt Lake algal structures. End members of this variation are illustrated in Figs. 2 and 3 by laminated and unlaminated textures, respectively. Between the two extremes, there exists a great variety of mixed laminated, unlaminated and poorly laminated textures.

All algal structures are lithified throughout and must be hammered or sawed apart to examine the internal struci$re. Laminated forms are relatively smooth on their upper surface, with laminations convex upward to form a stromatolitic structure. Internal texture consists of laminated micritic aragonite. Laminations are formed by millimeter-scale variation in grain size and impurities causing color variations. Small spherical voids in the micrite

438 R.B. HALLEY

Fig. 3. Unlaminated algal precipitate consisting of light, clustered aragonite grains (A and Fig. 8) and internal sediment. Fine internal sediment typically layers void floors ( B ) .

Fig. 4. Irregular, “tuffaceous” outer surface of an unlaminated algal mound.


possibly result from the removal of brine shrimp eggs, whereas other more asymmetrical voids appear to have developed as surface irregularities incorporated into the growing structure. There is no indication of algal precipitation within the homogeneous inner micrite of the stromatolite. Although the outer surface is coated with an algal film (see below), it is not known if algae play a major role in the formation of these laminated structures.

Unlaminated algal mounds are characterized by a highly irregular “tuffaceous” outer surface (Fig. 4). On the interior, the structures reflect the surface irregularity by high porosity (Fig. 3). Three distinct elements produce a clotted appearance typical of unlaminated algal structures: a framework; internal sediment; and inorganic cement.

Framework

The framework of unlaminated mounds in Great Salt Lake consists of a suite of aragonite precipitates, presumably algally induced. These precipitates are not skeletal elements but incidental precipitates resulting from chemical changes induced in the brine by algal physiology, presumably COz utilization. Some precipitates are molds and casts of filaments (Figs. 5,6 , and 7). These take the form of tubules preserved in micrite and are similar to, but larger than (about 30pm in diameter), the familiar fossil algal genus Giruanella (Fig. 5 ) . Other preserved tubules form frond-like arrays similarly preserved in micrite (Fig. 6). Filament molds are only found at the base and interior of some algal structures and are associated with gastropod and ostracod skeletal fragments (Fig. 7). Another framework texture consists of clusters of aragonite grains. Each grain, 100-300pm in diameter, is composed of slightly coarser aragonite (microspar) and appears much lighter in thin section than the surrounding micritic matrix (Fig. 8). This clustered texture is typical of the outer portion and irregular surface of the unlaminated algal structures of Great Salt Lake.

Internal sediment

Aragonite cement of various morphologies lithifies the internal sediment and framework elements of the Great Salt Lake algal mounds. Micritic filament molds act as loci for radiating aragonite needle cements (Fig. 5). Micritic cements encrust voids and cement internal sediment but may be indistinguishable from algally precipitated micrite, or in some instances, micritic internal sediment.

440 R.B. HALLEY

Fig. 5. Photomicrograph of filamentous framework precipitate. Note the needle cement radiating from micritic, filament-mold loci in the upper left field of view. Polanzers crossed .

THE LIVING SURFACE

The living surface of Great Salt Lake algal mounds consists of a film less than 1 mm thick, slippery to the touch and firmly attached to the underlying aragonite. The film can be removed by scraping or by mild acid dissolution of its aragonite substrate. The film on top of these structures is light green or brown in color and consists of simple, small (2pm in diameter), coccoid blue-green algal cells within a mucilaginous matrix (Fig. 9). In cross-section these cells are arranged in crude layers parallel to the surface (Fig. 10). This light top merges with a darker green area at the sides and bases of the structures, an area which contains a diatom (Fig. 11) in addition to the blue-green alga. These are the only two organisms recognized in the living surface of the algal structures. There is little or no detrital sediment present, although small particles of aragonite are sometimes found within the base of the film.


Fig. 6. Frond-like pattern of framework element composed of radiating filament molds. Polarizers crossed.

DISCUSSION

The mound structures in Great Salt Lake are, from all evidence, algal precipitates. Preservation of filament molds, lack of sediment in the surface film, hardness of the structure surface, and absence of a primarily clastic texture in these structures indicate they are not cemented, algally trapped and bound sediment.

However, it is apparent that the present living surface bears little or no relationship to the internal texture. The presence of filament molds within the structures is indicative of the importance of filamentous microorganisms in the development of these structures. Yet, filamentous forms are not living on the structures, nor have they been identified as living anywhere within the main body of the lake except in the southeast arm. Similarly, there is little morphologic resemblance between the surface film and aragonite grain clusters which compose the framework of the outer volume of the mounds. Thus, the organisms at the surface of the Great Salt Lake algal mounds are probably not those which are responsible for the internal structure.

The association of filament molds with gastropod and ostracod skeletal debris suggests that this occurrence is relict from times of less saline

442 R.B. HALLEY

Fig. 7 . Filament molds are associated with skeletal fragments such as the gastropod shell in the lower right. Detrital sand and micritic aragonite cement lie between the shell and the filament molds. Polarizers crossed.

conditions. Filamentous algal mats up to 2cm thick were found to cover large areas of the sediment surface in the less saline southeast arm of the lake during the summer of 1971. These mats were not reported by Eardley (1938) and probably have developed since the reduction of salinity in this

The stromatolitic structures may be produced by the films now at the surface. However, it is difficult to explain why the same film should produce a laminate structure in some areas and a non-laminate structure in the case of the unlaminated algal mounds. It appears that much of the variability of internal texture of the algal mounds might be explained in a change of biotic composition of the living surface.

Salinity is the dominant environmental variable of the Great Salt Lake and, therefore, is the most probable cause of changes in the algal mound flora. Such changes are recorded in internal texture of the mounds. The sensitivity of blue-green algae to salinity has been discussed by Desikachary (1959). One possible mechanism for this sensitivity may result from the varying ability of cyanophytes to expel sodium from their cells (Batterton and Van Baalen, 1971). Changes in algal flora brought about directly or indirectly by

arm.


Fig. 8 . Cluster of microspar aragonite grains (A), a framework element of the structure illustrated in Fig. 3 . The cluster is surrounded by cemented sand grains and pellets ( B ) , and micritic cements or micritic internal sediment (C). Polarizers crossed.

variations in salinity have not yet been well documented in the Great Salt Lake. Handy and Hahl(l966) suggest algal density is, in part, salinity-dependent and, therefore, alters the color and clarity of the brine.

Algal structures of precipitated carbonate are not known to be formed by blue-green algae in normal marine subtidal environments (Wray, 1969). The recognition of precipitated structures produced by a blue-green algal framework, their internal sediment and cement, suggests hypersaline, fresh or brackish water influence. There is no particular aspect of these algal mounds indicative of development in a saline lake. Considering the similarity between Great Salt Lake water and hypersaline water, it seems likely that structures very similar to the Great Salt Lake algal mounds might develop in hypersaline marine basins. It is conceivable that even the extreme variation in internal texture could be reproduced in a hypersaline marine basin by salinity fluctuations. Thus, the entire milieu of algal mounds should be considered when interpreting algal mounds or stromatolites. The reader is referred to the compilation of criteria for recognizing lacustrine deposits by Picard and High (1972).

444 R.B. HALLEY

Fig. 9. Coccoid blue-green alga which composes the film on the surface of the algal mounds. Sample was separated from the surface by mild acid dissolution of its aragonite substrate.

Fig. .lo. Section perpendicular to the living surface Qf an algal mound illustrating contact (A) between the algal film and aragonite substrate. Note crude lamination of cells in the algal film. Cells are stained dark and may have been disrupted during sample separation. Fig. 11. Small diatoms which populate the undersurfaces of the algal mounds causing them to be zoned by a darker green surface color on shaded surfaces.

ENVIRONMENTAL MODELS 44 5


Great Salt Lake algal mounds contain: (1) a framework of non-skeletal, algally induced aragonite precipitates; (2) internal sediment; and (3) inorganic cement. These three elements create a variety of laminated, poorly laminated, and unlaminated internal textures. Interior framework precipitates bear little resemblance to the present living film of the mound surface. Internal texture of the mounds is believed to be largely relict and to have resulted from precipitation by algae different than those presently living at the surface. The most probable cause of local extinction of the algal flora is change in brine salinity.

Precipitated blue-green algal structures in ancient rocks may indicate other than normal marine salinity and nearshore sedimentation. Extreme variation of internal texture probably reflects extreme environmental variability typical of closed basin lakes. Recognition of mounds similar to those in the Great Salt Lake can be a first step toward recognition of ancient hypersaline lake deposits, if such an interpretation is substantiated by consideration of the entire depositional milieu of precipitated algal mounds.


Chapter 8.6

THE GEOLOGICAL SIGNIFICANCE OF THE FRESHWATER BLUE-GREEN ALGAL CALCAREOUS MARSH

C.L. V . Monty and L.A. Hardie

INTRODUCTION

Much attention has been paid during the last decades to the leathery algal mats and algal-laminated carbonate sediments of the intertidal zone, while the “supratidal” freshwater algal marsh has been generally neglected by geologists. This is surprising considering that today this setting covers a large part of all the Holocene carbonate platforms of those subtropical regions enjoying a rainy climate, areas covering thousands of km2. The objective of this paper is to bring to the attention of geologists the significant features of the Holocene algal-marsh depositional environments of the Florida- Bahama region.

To document these active “supratidal” freshwater blue-green algal marshes, we have chosen to discuss and illustrate two main situations: (1) the “interior marshes” such as occur on the eastern half of Andros Island, Bahamas (Black, 1933, pp. 173-175; Monty, 1965a, pp. 163-236, 1967, pp. 68-76, 1972; Newel1 and Rigby, 1957, pp. 50,61) or in the Everglades, Florida (Dachnowsky-Stokes, 1928; Gleason, 1972, Gleason and Spachman, 1974); (2) the “coastal marshes” such as those fringing western Andros Island, Bahamas (Shinn et al., 1969; Gebelein, 1972b, 1972c; Hardie, 1975b; Hardie and Garrett, 1975).

THE INTERIOR FRESHWATER MARSHES

Setting

“Interior marshes” refers to those freshwater areas which are never (or very seldom) invaded by marine waters, even during the strongest storms. Only their very margin may be influenced temporarily by the brackish waters of the major tidal creeks. Such marshes are well developed on the eastern half of Andros Island, Bahamas, and cover much of the Everglades. Most of these

448 C.L.V. MONTY AND L.A. HARDIE


areas consist of low-lying, seasonally marshy plains covered with sawgrass (a sedge) and small Rhizophora, interspersed with tree-covered “hummocks” and shallow freshwater lakes (Fig. la); they develop over a karst cut in the Pleistocene bedrock. On Andros, this basement is deeply pitted and honey- combed, with sinkholes several meters in diameter, still active (“blue holes”) or partly filled with mud and rock debris. The ponding areas and lakes (Fig. l b ) correspond either to natural depressions in the bedrock or to former channels, creeks and ponds that have been progressively filled up, as can be seen on aerial photographs (Fig. lc). The karst is covered by a veneer of low- magnesian calcite mud (1.5-3 mole % MgC03) a few cm to 30 cm thick, on which algal mats grow.

In southern Florida, algal mats growing over a white low-magnesian calcite mud (averaging 2 mole % MgC03) up to 1.5m thick (Gleason, 1972; Idyll, 1969) cover an area of approximately 4000km2 (1 million acres, Gleason, 1972). This calcite mud is the second most important sediment after peat; it extends under the Recent marine muds of southern and southwestern Florida.

Climate

Both areas enjoy a tropical marine climate with long, warm and rainy summers followed by mild and dry winters. The showery character of the rainfall produces a succession of short wet and dry periods, which is very favorable to the growth of blue-green algae as it eliminates the aquatic plants proper which concentrate in more permanent waters. (Such plants are much scarcer on Andros than in the Everglades.)

The algal mats

All that is wetted or flooded by a few centimeters to a few tens of centimeters of freshwater during the rainy season is progressively covered with an algal mat interspersed with calcite crystals. The mat, which can reach a thickness of 2-3cm after about two months of growth, is dominated by two blue-greens : the terrestrial Scy tonema myochrous (Dillwyn) Agardh (called

Fig. 1. General setting of the “interior” blue-green algal freshwater marsh. a. Typical view of these flat, seasonally flooded, low-lying areas covered with algal mats and sedges that characterize the interior blue-green algal freshwater marsh. b. Aerial view of ponding areas along mid-eastern Andros Island showing the complex of freshwater lakes -- floored with white flocky calcitic mud - bordered by seasonally flooded algal flats passing laterally to the progressively rising and pine-covered Pleistocene bedrock. c. Aerial view of mid-eastern Andros, showing that the pattern and development of the actual freshwater lakes (white) and algal marsh stretches (various shades of grey) are geo- morphologically controlled by a former system of creeks and channels cutting through the coastal ridges.


S. hofmanni in Gleason, 1972; Gleason and Spachman, 1974) and the aquatic Schizothrix calcicola (Agardh) Gomont; they are associated with other filamentous forms (Johannesbaptistia pellucida, (Dickie) Taylor et Drouet, Microcoleus sp., etc.) as well as with unicells (Entophysalis major ErcegoviE, Gloeocapsa sp., etc.) and various amounts of diatoms. Where the mat is very spongy, it harbours a complex community of insect larvae, ostracods, s m d crustaceans, flatworms, etc. (Idyll, 1969). Such felt-like mats cover the ground, the rocks and the plant stems (reeds, sedges, rushes, etc.), and appear to be an extremely active and productive biological complex (Idyll, 1969).

The deposits

According to the relative elevation (differences measured in cm) of the substrate on which the mat grows, and the nature of this substrate (former mats, mud, rock), the mat will remain underwater, or at least wet, for longer or shorter periods. Various types of algal deposits will consequently be generated.

The higher mats, or those growing on rock, will have very short periods of growth and they will dry up almost as soon as the rain stops to form thin, earthy, strongly curled or convoluted chips. Under more permanent moisture conditions, mats will be able to reach thicknesses of 3-6cm. Three microstructural types have been described by Monty (1965a, 1967): (1) mats controlled by flabellate growths of Scytonema (Fig. 2a) (Monty, 1967, type 1, pp. 70-71, pls. 5, 6); (2)mats controlled by superposition ,of films of Schizothrix (Fig. 2b) (Monty, 1967, type 2, pp. 71-72, pl. 7); (3)mats formed by alternating layers of Scytonema and of Schizothrix (Monty, 1967, type 3, p. 72). These mats and resultant various microstructures are distributed along gradients of flooding, substratal moisture and or salinity (see Ch. 5.1, pp. 199-204). Upon drying, these mats will shrink considerably and break into “algal polygons” lying on the mud (Fig. 3a). They will not build up to form typical stromatolitic deposits but, instead, desiccation, rain impact, and other agents will contribute to their disintegration.

More interesting are the biostromal stromatolitic deposits described by Monty (1972) from the mud flats bordering freshwater lakes of Andros. Besides being flooded by as much as 30 cm of water during the rainy season, significant areas of the flats remain very wet for a long period of time because of the elevation of the water table; accordingly, algal structures suffer only short periods of drying. Here are found contiguous laminated algal heads and domes (Fig. 3d) up to 30cm in diameter and 6-10cm in thickness (see Monty, 1972). The laminated stromatolitic structure extends down to 10-15cm beneath the deposit and then passes to bioturbated calcitic mud resting on the bedrock (Fig. 3c). The domes, pancakes and lily-pad


Fig. 2. Interior blue-green algal freshwater marsh: dominant mats. a. Vertical section through a mat controlled by flabellate growths of calcified filaments of Scytonema myochrous (type 1, Monty, 1967). Scale:.l div. = 500pm. b. Vertical section through a mat controlled by the rapid superposition of thin-calcified films built by Schizothrix calcicola (type 2, Monty, 1967). Scale: 1 div. = 1000pm.

types of stromatolites may be complex and differentiate columns up to 6 cm high, made of a core of erect bundles of filaments (Fig. 3b) surrounded by concentric overgrowths of Schizothrix and Scytonema which drape around the core to produce mini-Conophyton-like structures (Monty, 1972, p. 759, pls. 14, 15).

In vertical section, the stromatolites show beige to white micritic magnesian calcite laminae alternating with greenish to brownish organic layers. The magnesian calcite averages 1.5-3 mole % MgC03. These values may be up to 8-10 mole 5% at the periphery of the marsh where the mats are contaminated by the salt water of the major creeks. In these inland marshes the mineralization as well as the stromatolitic lamination are entirely algal in origin. The lamination results from periodic algal growths and blooms whereas the magnesium calcite is organically precipitated in at least four sites: (a) calcification of the sheaths of Scytonema producing equant crystals (Fig. 4a); (b) deposition of blocky, dendritic or acicular crystals around the Schizo- thrix or in association with them (Fig. 4b, c, d); (c) deposition of rods, grains (Fig. 4e, f ) , rhombs (up to 10pm) or fans (Fig. 4g, h) in the mucilage surrounding the Scytonema; (d) micritization of algal material and colonies of unicells by bacterial action (Monty, 1965, pl. 26, 27; 1967, pl. 8). As described in Monty (1972, pp. 767, 776-778) from Andros and reported by Gleason (1972, p. 45) from the Everglades various metabolical factors may alter the original stromatolites and the deposit as a whole. Small-scale peeling of surficial laminae, due to the accumulation of photosynthetic gases, will leave scars in the stromatolites (Monty, 1972, figs, 21-22); on a larger scale,


important portions of the algal mat can peel off the substrate and drift away, leaving erosion-like scars in the overall algal deposit; whole algal domes may float to the surface to be rafted across the water and piled up over shallows (Monty, 1972, figs. 23-28). Similarly the cylindrical coatings around rushes and spikes (Monty, 1972, fig. 31; Gleason, 1972) may float away to accumulate chaotically. The whole sequence of events happens in very quiet environments. Finally, the algal deposit is bioturbated by crustaceans and worms, as well as being disturbed by the roots of reeds and other plants.

Genetic relationships between the algal mat and the underlying calcitic mud

These relationships must be discussed and understood with regard to the frequent association in the fossil record of stromatolites with fine-grained limestones, or the lateral facies change from coal-swamp deposits to aphanitic limestones.

(1) Field observations carried out during the dry season at places where algal mats grew directly on the karst surface showed them to crack into pieces that are disintegrated by the wind and the rain to yield calcitic dust and gravel that accumulate in the pits and the holes of the bedrock.

(2) Study of cores shows no sharp discontinuity between the stromatolitic deposit and the underlying mud, but rather a progressive fading away of the lamination which passes to an amorphous bioturbated mud.

(3) Rhombs and dendritic crystals such as found in the mat are also found in the mud, although the dendritic crystals seem to reorganize themselves into more compact ones deeper in the mud (Gleason, 1972; Monty, personal observation).

(4) Isolated, small clumps of calcified sheaths of Scytonema can be found in the mud (Monty, 1965a; Gleason, 1972). Such Giruanella-like bundles of twisted, compacted tubes are, however, rare, as the bacterial decomposition of the supporting sheath material enhances the disintegration of the calcified tubes into their constituent crystals. Another stage of the destruction of the calcareous tubes is their fragmentation into small pieces which appear as denser clots in the mat and in the mud (Monty, 1965a).

Fig. 3. Interior blue-green algal freshwater marsh: types of deposits. a. Blue-green algal mat cracking into “algal polygons” upon drying, a very common feature of inland Andros Island and of the Florida Everglades. b. Contiguous and interlocking algal domes overgrown by algal tufts that will differentiate into mini-Conophyton-like columns (see Monty, 1972). Wilmo Lake, eastern Andros Island. c. Core taken into algal biostromes of Fig. 3b and d: the organically controlled lamination fades progressively downward and passes to a bioturbated magnesian calcite algal mud. Scale in centimeters. d. Vertical section through a laminated blue-green algal biostromal deposit. Wilmo Lake, eastern Andros Island. Length of the knife 35 cm.


( 5 ) The mud is composed of the same low-magnesian calcite as the stromatolitic calcite.

(6) There is a seasonal variation of about 6% in the carbon isotopic ratio of the mats (Gleason, 1972; Monty, unpublished data). This ratio is lower (around - 3%) in layers precipitated during the wet season (great imput of light organic carbon due to great biological activity); it is higher in the layers formed during the dry season (highest value found by Gleason is around + 3%); as for the carbon isotope values of the calcitic mud, they are intermediate between the seasonal extremes of the mat values, suggesting that the sediment ratios are average values (Gleason, 1972).

What is really surprising and what might really be misleading in terms of paleogeochemistry is that “the carbon isotopes of the sediment and of the dry season algal mat calcite have values commonly associated with marine salinities” (Gleason, 1972, p. 211), if figures provided by Keith and Weber (1964) are taken. As the water in which the algal mats and the derived mud are found is always a typical freshwater, these stromatolitic formations and associated muds appear rather special in terms of isotopic composition and seem to constitute an exception to the general rules.

There is little doubt left that the calcitic mud which underlies the stromatolitic deposits of the freshwater marshes is algally derived and results from the complete breakdown of mineralized but non-cemented stromatolites and algal mats.

Fig. 4. Interior blue-green algal freshwater marsh: patterns and diversity of in situ mineralization. a. Calcified sheath of Scytonema myochrous made of equant crystals of low-Mg calcite. Note, underneath, an uncalcified filament of Microcoleus showing the specificity of the precipitation. S.E.M. Scale bar = 5 pm. b. S.E.M. view of a calcified film built by Schizotrhix calcicola, showing the tiny filaments embedded in an autochthonous matrix of blocky microcrystals of low-Mg calcite. Scale bar = 5pm. c. S.E.M. view of fresh interstitial mucilaginous material loaded with acicular crystals which pierce through or appear by “transparence” (see d). Scale bar = 1 pm. d. S.E.M. view of X-shaped dendritic calcitic crystal units after the elimination of the mucilaginous matter of photo c with hydrogen peroxide. Scale bar = 1 pm. e. S.E.M. view of tiny rods of low-Mg calcite associated with dense mucilage. Scale bar = 1 pm. f . S.E.M. view of closely packed anhedral grains of low-Mg calcite in deeper parts of the mat. Scale bar = 1 pm. g. Rhombs of low-Mg calcite clustering in the mucilage around a living filament of Scytonema myochrous. Transmitted light. Scale bar = 10pm. h. Fan-shaped growth of low-Mg calcite crystals on, or in between, the living filaments of Scytonema. Scale bar = 10pm.

456

Stra tigraph ical rela tionships

C.L.V. MONTY AND L.A. HARDIE

In the Everglades, the freshwater calcitic mud and algal mats may alternate with or pass laterally to freshwater peats (Leighty, 1965; Gleason, 1972). Peat mainly forms in standing waters where blue-green algae are displaced by the prolific growth of aquatic plants and where the organic matter remains unoxidized. On the other hand, algal-lime mud forms in areas where flooding by freshwater is a temporary periodical phenomenon. Peat and blue-green algal calcareous deposits are hence antipathetic. Local relief (in terms of centimeters to one meter) and depth of the freshwater table will determine sediment distribution.

The vertical alternation of peat and algal-lime deposits may be interpreted in terms of: (1) alternations of more pluvial periods with drier periods; (2) oscillations in the Recent marine transgression determining, inland, fluctuations in the altitude of the freshwater lens (further discussion in Gleason, 1972, pp. 258-259).

This freshwater complex “peat to algal mats and mud”, which is 1.5-3 m thick, passes laterally to the paralic environment composed of coastal swamps and mangrove forests (Fig. 5): here are deposited fibrous peats

Organic rich calcllutitk shell debris

0 5 10 15 - Krn.

Calcitic mud with freshwater gastropods or

titic Silt fibrous fresh-water peat

25” 81 I

Fig. 5. Succession of freshwater to marine environments and associated deposits in southwestern Florida. After Scholl and Stuiver (1967).


1 .

2.

3.

4 .

Shelly caIciIutlt#C sill lo shelly floe-grained sand

Fig. 6. Idealized cross-section through SW Florida, showing the stratigraphical relationships of the freshwater algal deposits. The transgressive sequence shows initial freshwater deposits (similar to the ones forming now in the Everglades, to the east (right)) resting uncomformably on the karstic Pleistocene bedrock; they interfinger with and, seaward, are overlain by, paralic swamp deposits that are topped by marine sediments. After Scholl and Stuiver (1967), modified.

derived from mangrove trees and/or saltwater marsh plants. In the submerged area are found detrital and shelly, organic-rich, calcilutific or quart- zose sands and silts (Scholl and Stuiver, 1967).

In the geological record, the Everglade situation (Fig. 6) would produce a transgressive interfingering among: (1) shallow marine mudstones and sandstones; (2) paralic coal beds asiociated with organic-rich mudstones and showing physiographical features such as channels, levees, bars, etc.; (3) a unit composed of fine-grained limestone with freshwater gastropods, ostracods, pollen and Giruanella-like remnants, locally stromatolitic (crenulated or showing algal heads), more generally structureless (except for mudcracks), alternating or interfingering with seams, beds or lenses of freshwater coal. Such situations must have been rather common in the Pennsylvanian of the Eastern United States for instance, and a good example has been described by Berryhill et al. (1971).

On Andros Island, there is no such intensive deposition of peat and no dense mangrove forests. The interior lacustrine flats and freshwater marshes described above pass westward to a coastal .complex made of tidal flats and freshwater blue-green algal marshes (see below). The marginal situation of these freshwater marshes endows them with very particular features which separate them clearly from the interior marshes.

THE COASTAL “FRESHWATER ALGAL MARSH-TIDAL FLATS COMPLEX”

The northwestern coastal belt of Andros Island has been investigated by Shinn et al. (1965, 1969) and by Garrett, Ginsburg and Hardie (see Hardie,


Fig. 7 . Aerial photograph of NW Andros (Three Creek Area) showing the juxtaposition of the coastal freshwater blue-green algal marsh (right) with the tidal flat complex (or “channeled belt”) seaward. The calcareous tidal flat complex replaces here the paralic organic-rich swamp and mangrove deposits of the Everglades.

1975a). We shall focus on this area as one of us (L.H.) has conducted extensive field work there whereas the other has had the opportunity of studying representative thin sections of the marsh deposits.

This belt, up to 20 km wide and 130 km long, comprises two main units (Fig. 7).

(1) Seaward: the tidal flat system, or “channeled belt” of Shinn et al. (1969), shows an association of active tidal channels, levees, ponds, algal marshes, etc.; it is about 3-4 km wide. The shores of the inter-channel ponds between mean tide level and mean high water are fringed by an “algal marsh” (Hardie and Garrett, 1975), a lush meadow of spongy Scytonema “pincushions” (they appear as dark patches in Fig. 7). Here, along the pond shores where the algae are frequently submerged in seawater, the filaments are not calcified; only moulds of them are left in a muddy matrix (see


Hardie, 1975b) as is also the case for the Scytonema of the Swash along eastern Andros (Monty, 1965, 1967). Cemented algal crusts form on the back- side of some of the levees; they have been described elsewhere (Hardie, 1969,1975b). (2) Landward: the tidal flat system passes abruptly into the freshwater

algal marsh proper (or “inland algal marsh” of Hardie and Garrett, 1975). This marsh, which is entirely supratidal, is about 4-8 km wide and runs for 50km along the island; it forms the single largest unit of the coastal belt covering an area of 200-400 km2 (Hardie and Garrett, 1975).

The landward (or inland) algal-marsh setting

The landward coastal marsh is sharply separated from the channeled belt (see Fig. 7), although the boundary is deeply indented by shallow creeks that are the headwaters of the major tidal channels which drain the rainfall runoff from the marsh. The main body of this extensive Scytonema marsh is essentially a flat featureless plain, the monotony of which is broken only by scattered growths of small Rhizophom. The living fauna is very sparse; Shinn et al. (1969) reported only the land snail Cerion and the large land crab Cardisoma. The whole surface of the marsh is covered with a continuous and rather thick (0.5-3cm) mat of tufted Scytonema locally invaded by sawgrass.

Heavy summer rains keep the marsh under water for long periods, the water depth being up to lOcm in June and July; furthermore, “runoff and ground water from the pineland to the east ensure that the marsh stays wet with freshwater through most of the showery summer months” (Hardie and Garrett, 1975). This is the period of algal growth. In the winter, the marsh is completely drained and the algal mats dry up into polygons. Seawater reaches the main body of the inland marsh only during infrequent hurricanes (once every several years) which carry sediment-laden waters from the Bahama Bank onto the marsh (see Hardie and Ginsburg, 1975).

The deposits

( I ) Algal-laminated sediments The living surface Scytonema algal mat proper, with its flabellate growth

of filaments, is essentially similar to the mats described above and elsewhere from eastern Andros Island and the Everglades. The underlying sediment shows a characteristic thin bedding and thick lamination (see Fig. 8) which extends to the bedrock over a thickness up to 1.7m (see Hardie and Ginsburg, 1975). The bedding is typically disrupted by roots of sawgrass and mangroves.

The fabrics and the microstructure of this deposit will be described in detail (see also Hardie and Ginsburg (1975). The layering is composed of


Fig. 8. Coastal blue-green algal freshwater marsh. Portion of a core taken in the algal-laminated sediment and showing the complex origin of the layering: layered structure within each major unit with Roman numerals results from: (1) a period of freshwater algal growth (basal dark brown), suddenly interrupted by (2) sediment-laden floodwaters washed onto the marsh by hurricanes (light tan to white layer), (3) pioneer colonization of surface of newly deposited sediment by Schizothrix (thin papery aragonitic film). Scale in centimeters.

3 units: (1) a tan to dark brown layer, up to 2.5 cm thick, representing several years’ growth of the Scytonerna mat (mat of type 1: Monty, 1965, 1967); (2) a white or gray layer (in some cases internally laminated or cross-laminated) of peloidal sand (0.1-10cm thick) with a few miliolids and penero- plids, representing a hurricane deposit (once every few years); and (3) a very thin (generally less than l m m thick) light aragonitic “paper crust” (most


Fig. 9. Thin-section in the coastal blue-green algal freshwater-marsh deposit showing a microsequence such as appears in Fig. 8 : the algal layer, composed here of well-preserved and calcified filaments of Scytonema, is overlain by a peloid layer containing some miliolids; the latter is capped by a thin aragonitic film on which lies the algal layer of the next microsequence. Scale bar = 500 pm.


commonly topping the peloid layer) representing the post-storm pioneering colonization of the surface by Schizothrix. A rhythmic succession of these units, algal layer + peloid layer +. aragonite “paper crust”, is clearly shown in Figs. 8 and 9. This micro-cycle is produced by several years’ growth of Scy- tonema in freshwater followed by a catastrophic flooding by sediment-laden seawater forced onshore by hurricane winds and then rapid colonization of the newly deposited surface by a Schizothrix mat. As the intensity and the regularity of these processes may strongly vary over the years, one can expect a significant range of variation in this basic scheme in which some units may be missing, reworked, more complex, etc. A brief account of the three types of layers might be most useful for the geologist and will differentiate clearly this coastal seawater-influenced algal-marsh deposit from the interior purely freshwater ones.

(a) The algal layer, where well preserved, consists entirely of thick flabellate growths of calcified branching tubes originally developed around the sheaths of Scytonema (Fig. 9). Hardie (197513) refers to these layers as algal tufa. The algal tubes are made of closely packed magnesian calcite crystals up to 16 mole 7% MgCO, (see Hardie, 197513) a few microns wide (Monty, 1965, 1967). Precipitation of a slightly coarser micrite in the interstitial mucilages tends to obliterate the voids although the porosity remains high. A probably inorganic microsparitic cement, made of interlocking crystals, may locally develop in some interstices but is not common. At any rate, the amount of interstitial micrite, hence the porosity, is highly variable: in places the calcite tubes appear self supported, in others they are almost completely embedded in micrite (Fig. 10a, b). This algal layer is very pure in case the overlying peloids do not infiltrate between the filaments; this may mean that the mat was actively growing when the s t o w occurred, the abundance of interstitial

Fig. 10. Some microstructural aspects of the algal-laminated sediments from the coastal freshwater marsh. a. Detail of the algal layer: situation showing vertical calcified and self-supported filaments of Scytonema with little interstitial micrite. Crossed polarized light. Scale bar = 100pm. b. Similar to a with well-preserved calcified filaments, but where the porosity has been considerably reduced by more intense deposition of interstitial micrite. Crossed polarized light. Scale bar = 100 pm. c. Deposit where the flabellate algal growth has been dismembered, the filaments fragmented and packed with the interstitial micrite to form a rather compact, but still identifiable algal fabric. Scale bar = 100 pm. d. Situation where fragmentation of the algal calcareous tubes into their component microcrystals led to the formation of more cryptic, less evident, algal structures. Scale bar = 100 pm. e. Detail of a peloid layer showing its normally poor cementation. Crossed polarized light. Scale bar = 100 pm. f. Contact between a poorly cemented peloid sublayer (characterizing the lower half of the peloid layer) and a cemented packstone sublayer where particles become poorly visible. Crossed polarized light. Scale bar = 100 pm.


mucilage preventing the downward movement of the peloids. In cases where the algal mats were dormant when flooded, peloids and foraminifera infiltrated in various amounts between the filaments.

In some algal layers, the calcified algal bundles are dismembered and the filaments fragmented; the fabric appears denser due to closer packing of the filaments and the interstitial micrite (Fig. 1Oc). In extreme cases of this the original flabellate growth has been completely reduced to a very dense packing of tiny fragments of calcified filaments closely pressed one against the other while the original interstitial micrite chokes all the cavities. Such layers are difficult to distinguish from inorganic muddy ones (Fig. 10d). In such algal layers, the algal mat probably was completely dried up, shrunk and the sheaths completely oxidized (hence tubes were very brittle) when flooded by marine waters.

The top of the algal layer often shows irregular contacts with the overlying peloid layer. This may mean that there has been erosion by the sediment-laden waters, or that dried-up chips of the mat surface were blown away by the wind before the phase of flooding.

( b ) The peloid layer generally overlies the algal layer abruptly. It is composed of rather well-sorted peloids ranging in size from 30-80 pm, associated with scattered foraminiferal tests, and is normally uncemented. However, some of these layers may be complex, showing a basal sublayer of non-cemented peloids, overlain by a sublayer of cemented peloids (Fig. 10e, f). The porosity is very high in the basal sublayer, where loose clumps of peloids and clumps showing incipient microcrystalline cementation (Fig. 10e) are found side by side. The incipient cement generally starts as a meniscus between two grains or shows an assymetric stalactitic fabric; locally interstitial voids are filled that way, but druses never seem to appear. The overlying cemented peloids appear to form a zone where peloids, foraminifera, etc. disappear almost completely in a yellowish micrite (Fig. 10f). Such sublayers are cut by vertical cracks, whereas horizontal dislocations may separate them from the underlying non-cemented sublayer. All this probably means that the marine flooding that brought in the peloids was followed by an intense desiccation; the uppermost part of the peloid layer cracked and shrunk while the evaporating waters deposited a compact microcrystalline cement.

(c) The aragonitic crust generally tops the peloid layer, whether the latter is cemented or not. It appears as a thin micritic film reminiscent of those deposited in felts of Schizothrix (see Ch. 5.1) although it is aragonite-rich. These lay& represent surface crusts such as those observed by Hardie and Garrett (1975) on the seawardmost edge of the marsh where “the sediment surface is covered by a white paper-thin, brittle aragonitic crust that cracks like thin ice when walked on”. This crust constitutes the substrate on which the Scytonena mat of the next sequence will settle.


Discussion. This rhythmically layered deposit characterizes all of the coastal freshwater marsh, and differs widely from the interior stromatolitic and algal- mud deposits. Instead of having a piling up of stromatolitic heads and mats associated with algally produced mud, the coastal marsh records the recumng impact of a marine inundation. Flooding of the marsh by sediment-laden sea- waters has several important effects on the overall marsh life and stratigraphy, besides its influence on the mineralogy (aragonite + magnesian calcites) and the diagenesis of the deposit. First of all, it brings in from the Great Bahama Banks sediment particles which are mechanically deposited on the mat; the algae have practically no control on the sites of sediment accumulation (elsewhere the algae differentially bind and trap sediment). This detrital sediment, which is deposited catastrophically on the algal mat, produces a characteristic and striking succession of marine and freshwater layers. Although the coastal marsh is essentially a freshwater setting, the overall deposit shows as many marine as freshwater phases, the catastrophic marine event contributing as much, if not sometimes more, in the stratigraphical record as the yearlong prevailing freshwater conditions.

Another important consequence of the aperiodic marine invasion is that the resulting sedimentary layer “fossilizes” the algal mat; the latter, being suddenly buried, will escape further destruction by desiccation, wind, rain, etc.; accordingly, there will be no significant production of algal mud as occurs inland. Furthermore, the sudden burial will also prevent the mat from various seasonal and complex differentiations (as occur inland), so that it will have a rather simple internal structure. I t is controlled by Scytonema myochrous and matches the simplest mat of the inland marshes (type 1, p. 450). However, if the initial microstructures are almost identical the diagenetic evolution of both mats will differ as the inland ones are subjected only to meteoric water and precipitate low Mg-calcite, whereas the coastal ones are aperiodically contaminated by salt water and precipite low and high Mg-calcite. Finally, the algal laminated sediments of the freshwater coastal marsh differ also considerably from the algal laminated sediment described from the intertidal flats of the Persian Gulf, of Western Australia, etc. In these areas, the algal layer is generally represented by a soft or leathery film that will progressively be reduced to organic streaks, whereas in the freshwater marsh, the algal layer is con- structional and adds autochthonous calcareous layers to the deposit.

(11) Stromatolites In the depressions where freshwater can become ponded for several weeks,

blue-green algae form domes, pancakes and large fleshy “lily pads” similar to the ones growing in the interior marshes and described by Monty (1972). However, in the coastal marshes the algal domes may be partially buried by peloids and marine bioclasts, and so their textures and fabrics will be different from the stromatolites of the interior marshes. The stromatoids dominant in the coastal marsh setting are produced by draping of Scytonema mats


Fig. 11. a. General view of a calcified algal head built by Scytonerna on slightly elevated ground within the coastal freshwater blue-green algal marsh, northwestern Andros. b. Verti- cal cross-section through the algal head shown in a. Note the radial flabellate growth of the filaments forming, in this case, a pure algal colony which has not been “contaminated” by sediment.

over polygonally cracked sediment-rich mats; they are the Type C heads described by Black (1933).

(111) Lithified algal crusts and heads These occur in the coastal marsh as well as in the channeled belt. Several

types can be recognized, each differing in the internal microstructure, the degree of cementation, the overall mineralogy, etc. A detailed description has been given by Hardie (1975b). Two main types will be described here from the high freshwater coastal marsh.

(a) Lithified sediment-free algal heads (pure algal tufa of Hardie, 1975b) have been reported by Black from the interior marshes where they form on slightly more elevated grounds (Black, 1933, pp. 170, 181; his fig. 2b, pl. 22). They occur in similar situations in the coastal freshwater marsh. Typical examples are represented by low rounded grey heads up to 2-3 cm high and 4-10 cm in diameter with circular, ovoid or lobate contours (Fig. 11). They are basically


made of radial flabellate growth of calcified sheaths of Scytonerna myochrous (Fig. 12a), very reminiscent of Monty's type 1 mats (1967; this Chapter p. 450) and of the algal layers of the laminated sediments described above. In the latter cases, however, the mats remain soft, although there is about the same amount of interstitial micrite between the algal tubes; the reason is probably that the interstitial crystals, a little larger than the ones calcifying the sheaths, remain embedded and separated by a pervading mucilaginous matter. In the lithified heads, however, the interstitial crystals may reach larger sizes than in the soft mat and interlock to form a microsparitic discontinuous mosaic (crystals 3-9 pm) (Fig. 12a, b). Perhaps the higher elevation at which the lithified heads are found results in frequent desiccation of the interstitial mucilages, a process which brings the crystals in close contact with each other to form an interlocking mosaic. Finally, pore-filling druses can be found where crystals get larger toward the center af the cavities, or show a tendency to form elongated prisms in the same direction. Phreatic meteoric cementation which might be inhibited by the mucilage in low-lying wet mats of the east coast, may well be more efficient in these drier crusts, and superimpose its effects on the algally induced precipitation. Other similar heads are more complex and show the co-habitation of two different (?) algae producing different fabrics: (1) Flabellate growths of Scytonerna rnyochrous with their sheath calcified by a micritic magnesian calcite to form fine-grained tubes patchily cemented together in a microsparitic cement. These tubes persist when the algal filaments die and can still be clearly distinguished from the cementing microsparite (Fig. 12a). (2) Flabellate growth of an alga which can be related to Scytonerna stad. crustaceurn Agardh, where the sheath never calcifies to form micritic tubes (Fig. 12c, d). In this case the microsparite is deposited just on the organic algal sheath itself (Fig. 12e); when the filaments die and disappear, a dense microsparitic fabric is left, pervaded by black organic streaks (Fig. 12f, g). According to the experience of one of us (C.M.) the Scytonerna myochrous stages would record growth in freshwater, whereas the development of Scytonerna stad. crustaceurn would testify temporary contact of the heads with brackish waters (see also Hardie, 197513, for a discussion of the environmental factors influencing Scytonerna growth and morphology).

Such heads have never been covered by hurricane layers and result from algal growth alone.

( b ) Lithified sediment-bearing algal heads. Structures become more complicated at the seaward margin of the marsh, where algal heads not only are covered by brackish water but also by sediment-laden seawater during infrequent hurricane flooding. This produces:' (1) the deposition of peloid layers; (2) the frequent inhibition of the calcification of algal filaments by salt water. The mineralogy of the crusts becomes also more complex and shows, besides aragonite cements and acicular linings, the development of high magnesian calcite up to 16 mole % MgC03 (see Hardie, 197513). These heads, which are


laminated in cross-section, show a basic scheme very reminiscent of the one found in the marsh sediments i.e. the recurrence of a unit composed from base to top of an algal layer overlain by a peloid layer that is then topped by a micritic lamina (Fig. 13).

The algal layer is seldom well preserved due to the lack of general calcification of the sheaths: at places, it shows scattered calcified algal remnants, or isolated long bundles of filaments, stretching across a loose layer locally contaminated by “floating” pockets of poorly cemented peloids (Fig. 13). Frequently, however, the algal phase has almost completely disappeared as a result of the oxidation of the uncalcified sheaths. This may produce several subfabrics such as: (1) voids which are elongate parallel to the lamination, floored with a lining of aragonitic needles and overlain by a poorly cemented pellet layer (Fig. 14a); in this case, the peloids were deposited on a tough mat and acquired their incipient cementation before the total oxidation of the underlying mat; (2) layers showing a patchy distribution of peloids, irregularly scattered in juxtaposed cavities, or even layers made of vugs interspersed with groups of loosely packed peloids (Fig. 14b); in these cases, the peloids infiltrated between the algal filaments, and become poorly cemented as the algal framework progressively disappeared.

Fig. 12. Some microstructural aspects of the calcified algal heads of the coastal freshwater marsh, such as shown in Fig. 11. a. Vertical section through a growth of Scytonema myochrous typical of freshwater conditions: the algal sheaths have first been calcified by deposition of tiny crystals of low-Mg calcite to form “skeletal” calcareous tubes rooted in the sheath material. The later interstitial cementing micrite and microsparite was deposited on this initial framework. The process will lead to a very good preservation of the algal growth such as found in the Paleozoic Orthonella. Crossed polarized light. Scale : 1 div. = 100 pm. b. Cross-section through a bundle of filaments shown in a: each tube presents an inner very fine-grained micritic ring, corresponding to the calcified sheath, overgrown by coarser crystals of the interstitial matrix. Cross polarized light. Scale bar = 30pm. c. Vertical section through a growth of Scytonema that has been subjected to brackish or saline conditions during its growth (referred to S. myochrous stad. crustaceum in Monty, 1967). The inner calcareous tube fossilizing the filament as in a is here missing; the high-Mg calcite micrite cement or microsparite is deposited on the organic sheath itself. As shown in the following views, the preservation of the filaments f becomes very poor. Compare with a. Crossed polarised light. Scale: 1 div. = 100pm. d. Cross section through bundle presented in c and showing the cementing microsparite deposited on an organic sheath, in the absence of an inner micritic calcareous tube. Cross polarized light. Scale bar = 10pm. e. f, g. Series of photographs showing the obliteration of the algal structures during progressive cementation of heads with originally uncalcified filaments such as in c: In e, microsparitic crystals ( 3 - 9 ~ ) are deposited directly onto the foreign surface of the organic sheath (appearing in black). Scale 1 div. = 10pm. In f , the microsparite extends into the majority of the interstices: the uncalcified dead filaments f shrink and become less and less distinct, appearing as discontinuous black streaks. Crossed polarized light. Scale bar = 30 /.l.m. In g we reach the final stages showing subcontinuous microsparitic fabric pervaded by scattered and residual black linear patches of the original filaments (f ). Cross polarized light. Scale bar = 30pm.


Fig. 13. Detail of a vertical cross-section through a calcified head such as shown on Fig. 11 but that has been repeatedly inundated by sediment-laden marine waters due to its location near the seaward marsh edge. Such heads show similar microsequences to the ones found in the algal-laminated sediment (Figs. 8 and 9), although greater contamination by marine waters inhibits calcification of algal filaments and so the algal layer is poorly preserved (Fig. 13 base). See comments in the text.


Fig. 14. Some microstructural aspects of the lithified sediment bearing algal heads growing at the seaward edge of the marsh. a. Development of empty cavities due to complete oxidation of non-calcified filaments of the former algal layer. Cavities may be floored by irregular linings of acicular aragonite, as appears in the lower third of the photograph, and may be capped by unsupported poorly cemented pelletal layer. Cross polarized light. Scale bar = 500 pm. b. Basal algal layer secondarily transformed into a “vug and pellet clump layer” due to the infiltration of pellets in between the filaments during the marine invasion and their rapid incipient cementation, while the uncalcified algal filaments rapidly decayed, leaving vugs.

This algal phase, as complex and variable as it is, is overlain by a peloid layer, the cementation of which increases from base to top, to finally form a micritic sublayer (Fig. 13). Cementation results from at least two processes: (1) precipitation of a micritic cement around and in between the peloids, with eventual microsparitic phases; (2) precipitation of void-filling aragonitic needles. This micritic layer is overlain by a thin aragonitic film as in the marsh sediment. It is often broken and disrupted by bundles of Scytonerna that produced the next algal layer.


I I I I

I I

I I I I

HIGH ALGAL MARSH I CHANNELLED BELT I OFFSHORE MARINE

- 4 - 0 Km -1- 3 - 4 K m . I- up to 100 Km -

Thinly Ihm1n.t.d 8-n and ;Z;,;m:InEd& lhlnli

b..ch ,Ids. ndh“,.

Lbht #my bl0turb.t.d plrnlcrlt lc .dJmcenl rnwln. s.dlm.nt.

Fig. 15. Stratigraphical cross-section along NW Andros Island showing the relationships between the freshwater blue-green algal marsh deposit (to the left), the tidal flat system and the off-shore marine deposits. The transgressive situation is evident. Compare with situation in the Everglades (Fig. 6 ) (modified from Hardie and Ginsburg, 1975).

Stratigraphical relationships

Lateral The algal-laminated deposits of the coastal marsh (including lenses of domal

stromatolites and capped by lithified algal heads and crusts), form a well- defined magnesian calcite-aragonite lithosome, a little less than 2 m thick (Fig. 15). It encroaches landward on a locally outcropping karstic basement interspersed with residual calcareous lakes, and generally covered with algal mud, less than 30cm thick, supporting algal polygons and domes. I t interfingers seaward with a complex of aragonitic tidal flat deposits topped by laminated levee and beach ridge sediments (the “laminite cap”, see Hardie and Ginsburg, 1975) which locally carry thin “dolomitic” crusts (Shinn et al., 1969).

Vertical Along NW Andros Island, the freshwater algal-marsh deposits accumulated

on the Pleistocene bedrock itself to form a rather homogeneous body that seems to be progressively destroyed by the landward migrating tidal flats. In this case (as in the Everglades and eastern Andros), it appears to be the first deposit to accumulate over a low-lying karst in equilibrium with the rise of the freshwater lens.

It is interesting to oppose this transgressive situation t o a regressive one resulting from shoreline progradation over a submarine platform, as appears to be the case along southwestern Andros Island (Fig. 16); here the


progradation of a submarine platform, for the last 3,000 years, has been followed by seaward displacements of successive shorelines (Gebelein, 1970, 197213, personal communication). The sedimentological framework runs like this: the establishment of a new shoreline, ahead of the extant one, isolates inner shallow lagoons in which peloids as well as skeletal debris of the local fauna will soon accumulate; sand spits driven by the wind will tend, as they grow and join, to split the lagoon into isolated shallow bays that will be progressively filled up by sediment influx from storms and in situ production. After complete infilling, a widespread flat algal marsh develops and covers the previous deposits. At this stage, the marsh is not yet typically “supratidal” but is rather equivalent to the “upper intertidal” (Gebelein, loc. cit.) although normal marine waters never reach it. As in the north, this marsh is covered by 15cm of freshwater during the rainy season and dries up completely in winter; algal laminated sediment, similar to the ones of the NW coast, progressively accumulate; towards the slightly more elevated flanks of the marsh, the algal mats pass to several types of lithified thinly laminated crust, about 3cm thick, similar to the ones described above (pp. 466-471) and by Shinn et al. (1969, p. 1218). As the algal marsh progressively fills up, the crusts migrate toward its center and will cover its whole surface in the very final stage of the shallowing-upward sequence, when the marsh surface reaches the supratidal zone (Gebelein, 1975).

In this situation, accordingly, the algal-marsh deposit constitutes the very last unit of a shallowing-upward sequence. In many places of the SW coast of Andros, the algal marsh starts just behind the present-day beach ridge or hammock (Fig. 16), a situation somewhat different from what has been described from the NW coast. According to Gebelein’s model (1975), such a situation would imply a net stabilization of the present-day shoreline, permitting the complete infilling of the associated landward lagoon (namely by storm deposits); (this has not been the case for former shorelines inland, as many bays have not been filled yet, Fig. 16). One might expect that such littoral freshwater marsh deposits will suffer a more pronounced diagenetic evolution, due to greater proximity of seawater, than the deposits of the NW coast.

CONCLUSIONS

The present paper has attempted to describe and I iscuss various aspects o the environments- and the deposits that characterize what appears-to be a rather important, but too often neglected geological setting: the freshwater blue-green algal marsh. Its importance stems not only from its wide present- day distribution, but also from the great variety of structures and fabrics that can be found in the deposits and which are not uncommon at all in the

414 C.L.V.MONTYAND L.A.HARDIE


geological column. The topic has been approached by opposing two main settings: (1) the interior marsh, and (2) the coastal marsh, where two different stratigraphical situations have been presented. Geologically speaking the resulting deposits bear few unequivocal clues as to their prevalent geochemical environment of formation and some classical arguments are really misleading:

(1) Carbon isotopic data gathered from typically freshwater mats and associated muds, from interior as well as from coastal marshes, may often show abnormal values, typical of sediments formed in marine waters.

(2) The development of freshwater blue-green algal mats produces significant concentration of magnesium within the deposit; the magnesium is not only abnormally concentrated onto the sheath material as shown by Gebelein and Hoffman (1971), but the carbonate phase which is precipitated around the filaments and in the mucilages is a magnesian calcite ranging from 2 to about 16 mole 5% M g C 0 3 . This magnesium may later become available for an eventual dolomitization of (parts of) the deposit as postulated by Gebelein and Hoffman (1971). Such dolomitic stromatolitic deposits may be automatically interpreted as of intertidal origin. (3) The severe life conditions in the marsh prevents the development of a

rich and diversified freshwater fauna; typical freshwater fossils will be rare and limited to some gastropods and ostracods. Furthermore, the coastal freshwater marshes are periodically “contaminated” by tests of marine animals which are washed over during hurricanes; in many cases, these marine “markers” are more abundant or more conspicuous in the resulting algal-laminated sediment than the freshwater fossils.

(4) As seen in the Everglades and in the Andros interior marshes the evolution of the blue-green algal-marsh deposit will not always yield stromatolitic beds, but quite frequently produce only a fine-grained structureless limestone with scattered fragments or clumps of filaments.

The association of cracks, polygons, floated mats, bioturbation by roots, particular petrofabrics, will point to “temporarily exposed environments”, but not necessarily to an essentially freshwater terrestrial setting. In many geological situations, detailed study of the deposit joined to thorough understanding of the lateral and vertical facies successions will yield, after sound

Fig. 16. Aerial photograph of southwestern Andros Island. To the NW, intense erosion produces active shoreline retreat. The southwestern shores, on the contrary, have been prograding for the last 6,000 years, according to the following sedimentological model : establishment of a new shore ahead of the older one, on a submarine prograding platform; isolation of an inner lagoon which splits up into isolated bays if shoreline remains active (importance of the input during hurricanes and storms). When the shoreline remains stable for long periods of time (as the present one) the infilling of the bays leads to the development of a widespread freshwater marsh that tops the whole regressive sequence (as occurs today just behind the present shoreline). If on the contrary the shore becomes replaced by a new one that settles a few miles ahead, the inner bays and lagoons will starve and the development of the latest phase, i.e. the freshwater marsh, will be considerably delayed.


comparison with present-day models, the more reliable indications as to the most probable environment of formation. In the overall freshwater marsh litho-biotope itself, sedimentological gradients (thickness, frequency and composition of the detrital layers) as well as petro-geochemical (Mg/Ca ratio, diversity and types of cement, etc.), or structural gradients, are strongly marked across the strike (from coastal to interior marshes) and should help the paleo-ecologist. Also, the population gradients, going from typically marine assemblages, to the tidal flat complex where organic diversity decreases while dominance suddenly increases, to the freshwater algal-laminated sediments where skeletal remains are almost absent, might give indications as to the general ecology of the genetic environment of the latter settings; indeed, this lack of fossils cannot result from hypersaline conditions (or else sabkha features would be recorded); it cannot either result from anoxic conditions as by its sedimentological structures the deposit is typically “shallow water” and is partly built by aerobic organisms. We agree that these are just indications, not clues; that is why we need more penetrating studies of Recent freshwater blue-green algal marshes.

Identification of such deposits may yield interesting paleoclimatic data; as opposed to the sabkha sedimentary suite, which records arid climatic conditions, the freshwater algal marsh reported here is typical of wet and rainy tropical belts characterized by short showery periods alternating with seasonal mild dry periods. Flooding of the marsh by freshwater must not last too long as standing water will inhibit stromatolitic growth while favoring more competitive aquatic plants leading to peat formation. Accordingly, the development of freshwater blue-green algal calcareous marshes reflects at least two environmental factors: (1) the presence of a flat, low-lying4 locally ponded, calcareous platform; (2) climatic conditions bringing short showery periods, long enough to allow the blue-greens to bloom and contribute their seasonal deposit, but short enough to prevent standing water and colonization by higher plants.

According to the local relief (in terms of centimeters to one meter at the most), the geographical situation, the groundwater alkalinity, the characteristics of the rainy periods (duration of flooding and/or substrate moisture, frequency of wet and dry phases), etc., various types of deposits will form, among which are:

(1) Continental stromatolitic biostromes made of adjacent, calcareous but unlithified, algal heads and domes (Fig. 3b, d; Monty, 1972).

(2) Continental algal muds associated with algal mats and polygons (Fig. 3a; Black, 1933; Monty, 1967; Gleason, 1972). (3) Continental lithified laminated biscuits associated with lime muds

(Mawson, 1929). All these formations are monomineralic, i.e. made of low magnesian cal-

cites; furthermore, except for features like desiccation cracks, they are entirely algally controlled.


(4) Coastal complexes of algal-laminated sediments associated with various types of cemented crusts and heads. These deposits are polymineralic (aragonite, various low and high magnesian calcites, protodolomite, dolomite (?), etc.) and may suffer strong diagenetic alteration. The algae control only the structures and the initial mineralogy of the algal-mat phase, whereas meteorological agents control the sedimentary phase, the eventual alteration of the algal mat and the diagenetic potentialities of the groundwaters.

The three-dimensional shape of the resulting deposits will include : (1) Calcareous algal lenses scattered through various continental deposits

(2) Widespread, thin, sheet-like calcareous strata covering flat carbonate

(3) Calcareous wedges interfingering with coastal marine sediments. Notwithstanding climatic conditions nor the availability of calcium ions

and COz, freshwater blue-green algal-marsh deposits may be expected to occur :

(1) At the base, or ahead, of a transgressive sequence, when the sea slowly invades a flat and low-lying platform, producing inland a progressive and gentle rise of the freshwater lens.

(2) At the top of regressive calcareous sequence where the marsh pro- grades over retreating shores and former intertidal flat sediments.

(3) At a smaller scale, in rhythmic, cyclothemic deposits, at the top of shallowing-upward cycles.

(4) Over paleokarsts, during low sea-level stands.

(limnic coals, calcareous sandstones, etc.).

platforms.


Chapter 8.7

FRESHWATER STROMATOLITIC BIOHERMS IN GREEN LAKE, NEW YORK

Jane R . Eggleston and Walter E . Dean

INTRODUCTION

The ability to interpret stromatolites in the geologic record is dependent upon a good understanding of the various environments in which Recent stromatolites form. One such environment, Green Lake, located in Green Lakes State Park, near Fayetteville, New York (Fig. l), offers the necessary physiographic, biological, and chemical properties to encourage extensive growth of freshwater stromatolitic bioherms. These bioherms occur as solid, wave-resistant lenses protruding outward from the shore at many locations around the lake. They grow primarily by a process of trapping of carbonate sediment by blue-green algae and mosses, with subsequent cementation by precipitated CaC03.

So much previous work has been done in the limnology of this lake that we will present only a brief summary. Eggleton (1931,1956) first studied Green Lake in 1925 and described the general limnological features and meromictic nature of the lake. The bioherms were described by Bradley (1929a), when he compared them to the algal reefs of the Green River Formation in Wyoming. Howe (1932) described the lake’s littoral deposits. The theory that the lake basin originated as glacial melt-water plunge pools was suggested by Miner (1933) and later verified by Coon (1960) and Muller (1967). Chemical properties of the lake were studied by Deevey et al. (1963), Berg (1963), Turano and Rand (1967), Takahashi et al. (1968), Brunskill (1969), Brunskill and Harriss (1969), and. Brunskill and Ludlam (1969). Biological aspects have been described by Jackson and Dence (1958), Bishop (1964), Harman and Jackson (1967), and Culver and Brunskill (1969). Eggleston (1972) and Dean and Eggleston (1975) described the algal bioherms and compared them to a Bermuda dgal-gastropod cup reef.

Green Lake is approximately 1 km long and 0.3 km wide with a maximum depth of 52.5 m. The total drainage area for Green Lake and nearby Round Lake is relatively small (4.3 km2), and most of the water is supplied by ground water (Brunskill and Ludlam, 1969). Both lakes are meromictic with

480 J.R. EGGLESTON AND W.E. DEAN

LAKES

SYRACUSE

GREEN LAKE ELEV 128m

8 1 0 200 I

I METERS -

POINT

Fig. 1. Map of Green Lake, Green Lakes State Park, Fayetteville, New York.

a sharp chemocline at approximately 18 m. The mixolimnion stratifies thermally during the summer with a thermocline at approximately 5 m. The temperature of the monimolimnion remains relatively constant at about 7OC throughout the year (Brunskill and Ludlam, 1969). The dissolved-oxygen content of the mixolimnion varies seasonally, but is usually near saturation. The monimolimnion, on the other hand, is always anoxic. The entire lake is saturated or supersaturated with respect to calcite throughout the year as a result of surface and ground-water influx of calcium carbonate and calcium sulfate from limestone and gypsum in the drainage area (Brunskill, 1969). The degree of supersaturation ranges from 1 to 2 times in the monimolimnion to 8 times or more in the surface waters during the summer. The supersaturation results in massive precipitation of calcite at a rate of about 2 g per m* per day, mostly during the period between late May and August (Brunskill, 1969).

METHODS

General observations of the location and morphology of the Green Lake biohermal lenses were made and samples were collected using SCUBA. The internal structure of the samples was studied in acetate peals, polished and etched slabs, and thin sections. Broken pieces and polished sections were photographed under a scanning-electron microscope (SEM). The polished sections were etched for 30 sec with 1/3N formic acid.

Mineralogy of biohermal materials was analyzed using a General Electric XRD-5 X-Ray Diffractometer and Ni-filtered Cu-K, radiation. The mole percent MgCO, was determined using curves constructed by Goldsmith and


Graf (1958) in which mole percent MgC03 is plotted against the position of the d z l l peak.

RESULTS

Bioherm distribution, morphology, and associated sedimentary environments

Stromatolitic biohermal lenses occur commonly throughout the lake’s littoral zone except in the northern end of the panhandle. The most extensive bioherm development, however, is located at Dead Man’s Point, on the northeastern side of the lake (Fig. 2). The bioherm grows outward as lobate,

Fig. 2. Stromatolitic bioherms growing at Dead Man’s Point, Green Lake.

overhanging ledges which protrude as much as 8 m into the lake. The solid ledges extend from lake level to a depth of 10m. In the spring, the ledge is as much as l m underwater, but by late summer much of it is exposed above lake level and is only wetted when storm waves pass over it. The top of the bioherm is relatively flat with a slightly raised outer edge, and the sides are rounded, resembling a cauliflower-like growth, interrupted by vertical


Fig. 3. Photomicrograph of blue-green algal filaments with entrapped calcareous particles in a Green Lake algal mat (bar = 0.3 mm).

grooves. The underside of the well-defined lip slopes inward toward,the bank of the lake, ending at an outcrop of Silurian Vernon Shale at a depth of about 10 m. Below 10 m, a thick marl deposit slopes steeply toward the lake.

Associated with the stromatolitic biohermal lenses are two zones of deposition: the littoral and the profundal zones. The narrow littoral zone is occupied by a soft marl bank extending 2 m below water level; the bank projects 1-3m into the lake. This zone ends abruptly, dropping at a 40" angle into the profundal zone. Littoral vegetation is composed primarily of the calcareous green alga Cham sp., filamentous green and blue-green algae, and diatoms. The branches of fallen trees on the marl bank are often encrusted with a coating of calcium carbonate. The sediments comprising the littoral bank facies are primarily biogenic in origin, consisting of Cham fragments, with minor amounts of aragonitic gastropod shells and wood chips. Much of this sediment is eventually carried down the steep marl slope to the profundal zone. Ludlam (1969) found the profundal sediments to be of two types: (1) beds of alternating dark and pale laminae and (2) massive, often graded, beds. The laminated sediment is primarily chemically precipitated calcite, appearing as yearly couplets, or varves. Each couplet consists of a pale lamina which grades upward into a dark, organic-rich lamina, with sharp contacts between couplets. The massive, usually graded beds are deposited by turbidity cur-


Fig. 4. SEM micrograph of columnar pattern of calcite crystals growing perpendicular to the bioherm surface (bar = 50 pm).

rents arising from the steep slopes just below the littoral zone. These massive turbidite beds, commonly with sole markings, account for 40-65% of the total profundal sediment accumulation, and consist of littoral, terrestrial, and biohermal material.

Organ isms

Green algae, diatoms, filamentous and unicellular blue-green algae, aquatic mosses, six species of gastropods, and some small freshwater sponges have been observed living on the bioherm surfaces. The most common blue-green algae are Entophysalis rivularis (Kutzing) Drouet, Schizothrix calcicola (Agardh) Gomont, and Calothrix parietina (Naegeli) Thuret (personal communication, Dr. Herman S. Forest). These blue-green algae form a mat up to about 1 cm thick, coating much of the biohermal surface from early spring until late autumn. The gelatinous sheaths of the algae trap and bind calcareous particles (Fig. 3) such as calcified reproductive structures and cortication tubules of Chara (which commonly grows on the adjacent marl ben’ch). The carbonate sediment, generally ranging in size from 2 to 20pm, is too fine grained to permit recognition of Chara structures, but some was probably derived from comminution of Chara.


Fig. 5. SEM micrograph of radiate growth pattern of calcite crystals (bar = 20pm).

Microstructure and mineralogy

The internal structure of the Green Lake bioherm results from alternate deposition of detrital low-Mg calcite (which the algae trap), and calcite precipitated from the lake water. Precipitation occurs primarily between late May and August, when the saturation in the epilimnion increases threefold to fourfold. This increase can be attributed to the higher water temperature (Brunskill, 1969). Epiphytic algal photosynthesis, however, may be even more important than temperature in precipitating calcite near the bioherms, although this has not been documented. The crystals normally grow in concentric layers, forming botryoidal structures on the bioherm surfaces. Most botryoidal nodes are columnar (Fig. 4), perpendicular to the growth surface, reflecting crystal growth perpendicular to bioherm surfaces, but other crystal growth patterns, such as a radiate pattern (Fig. 5), were also observed. Because of the nodular surface growth pattern, numerous internal and external cavities are created as the various bioherrn projections grow together. These cavities are progressively filled with drusy calcite crystals as large as 100 pm long (Fig. 6).

Just below the actively growing blue-green algal mat, a spongy, extremely porous crust is common. The crust is 0.2-5cm thick, with linear pores oriented perpendicular to the growth surface. These pores are the molds of former algal filaments. Below the crust, the bioherm is much more dense,


Fig. 6. SEM micrograph of calcite crystals growing into a cavity in Green Lake stromatolites. Note voids (indicated by arrows) within crystals, interpreted as locations of former liquid inclusions (bar = 40 pm).

and traces of the algal mats appear as black, organic-rich laminations (Fig. 7). These laminations show up in the microstructure as linear voids parallel to the growth surface. The voids are partially or completely filled with calcareous particles, measuring 1-5 pm in diameter, with no regular shape or orientation (Fig. 8). The voids were presumably left upon decay of the algal filaments, and the small calcareous particles are the loose sediment that had been trapped and bound by the algae. These partially filled voids alternate with layers of precipitated calcite crystals.

DISCUSSION AND CONCLUSIONS

Stromatolitic bioherm growth in Green Lake is strongly related to the lake’s unique physiography, biology, and chemistry. Because the lake is supersaturated with CaC03 during most of the year, the calcareous alga Cham thrives. Sediment derived from the lake banks, including Cham fragments and chemically precipitated calcite, are washed over the bioherm, where they are trapped and bound by the mucilaginous sheaths of blue-green algae. The algae may also cause direct precipitation of CaC03 onto their filaments.


Fig. 7. Hand specimen photograph of internal structure of Green Lake bioherm material. Note black organic layers (indicated by arrows) parallel to growth surface.

Upon decay, the algae leave a spongy, porous structure. Bradley (1929a) found that many of the fossil reefs in the Green River Formation display this same spongy structure. Subsequently, cementation by precipitated calcite occurs. The laminated pattern created by these processes is obvious in hand specimens as well as under light and scanning-electron microscopes.

Bioherm growth begins on solid objects, such as tree branches, pebbles and even beer cans, in the littoral zone or on the steep shelf slopes of the lake. The maximum depth at which growth occurs is controlled by the chemocline at 18 m. At the chemocline, there is a sharp increase in water density and decrease in light transmission with depth. The lack of adequate light below 18 m prohibits algal growth. The most extensive bioherm development is on the northeastern side of the lake, presumably due to algal growth response to mid-afternoon sunlight. Because of the steep shelf slopes and extreme depth of the lake, the biohermal growth is restricted to close to shore, where the bioherms grow upward and outward into the lake, creating a lip. If the physiography of the lake were different, the bioherms might have taken a different form. For example, in the Tully Green Lake, New York, we found solid calcareous mounds growing upward from the outer edge of a shallow marl platform that extends more than 50m from shore. These mounds usually contain pebbles or other debris and often have tubular holes,


Fig. 8. SEM micrograph of linear void left by a former algal mat (indicated between arrows); void now partially filled with small calcite crystals (bar = 15 pm).

probably left by the decay of tree branches. A carbonate encrusted golf ball also was found, proving that these mounds are forming today. Given time, these mounds could grow to be as extensive as the Green Lake stromatolitic bioherms, but in the form of offshore, platform-edge mounds, rather than lens-shaped ledges attached to shore. In Onondaga Lake, Syracuse, New York, stromatolites are growing, but in yet a different form - as oncolites. The lake is much larger than either the Fayetteville Green Lake or the Tully Green Lake, and with greater exposure to winds there is consequently greater wave activity. The oncolites occur on a soft, shallow (0-2m) platform along the north shore, where enough agitation occurs to overturn and round them. These three lakes all have the necessary biological and chemical conditions for stromatolite growth. However, the form the stromatolites take is different due to different physiographic conditions (lake size, depth, and slopes) in each lake.

In conclusion, in using the Green Lake stromatolitic bioherms and associated littoral and profundal carbonate facies as a model in interpreting ancient rock sequences, the lake’s unique environmental qualities must be taken into account. Calcium-carbonate saturation of the water, conditions suitable for algal growth, and steep slopes, all tend to encourage growth of the stromatolitic biohermal lenses in Green Lake.

488

ACKNOWLEDGEMENTS

J.R. EGGLESTON AND W.E. DEAN

We greatly appreciate the assistance of Dr. James Sorauf in allowing us to use the scanning-electron microscope at the Department of Geology, SUNY Binghamton. Mr. Ronald Strong provided valuable field assistance in our study. The Green Lake algae were identified by Dr. Herman S. Forest and Mr. Tracey F. Maxwell at SUNY Geneseo and the identifications were verified by Dr. Francis Drouet of the Philadelphia Academy of Science.


Chapter 8.8

HOT-SPRING SEDIMENTS IN YELLOWSTONE NATIONAL PARK

M.R. Walter

INTRODUCTION

There is very little literature on hobspring and related sediments of Yellowstone National Park. The purpose of this chapter is to outline what is known about sediment types and distributions (to provide the sedimentary context for Chapters 3.3 and 6.2 in which geyserite and hot-spring stromatolites are described) and to provide a sketch of the “hot-spring facies” (to facilitate its recognition in the geological record). Hopefully the inadequacy of the data presented here will provoke a more intensive study of hot-spring sediments in Yellowstone or in one of the other major hot-spring areas (e.g. Iceland, New Zealand, Japan).

GEOLOGICAL SE‘ITING

The geology of Yellowstone is briefly described by Keefer (1971). The present geothermal activity is one of the last phases of a period of extrusive volcanism which began about 2 m.y. ago and continued until 70,000 years ago (Christiansen and Blank, 1972). The rocks produced are largely rhyolites and rhyolitic welded tuffs, with some basalts; they form the “Yellowstone rhyolite plateau”. Within the plateau is an enormous elliptical caldera, 70 by 45km wide, partly filled by lava flows. Pleistocene glacial sediments and Pleistocene and Recent fluviatile sediments are intercalated with and overlie the volcanics. Some of these sediments are described by Howard (1937).

DISTRIBUTION AND CLASSIFICATION OF HOT SPRINGS

The springs are scattered over an area of several thousand square kilometres; the area of occurrence is irregular in shape but has a maximum dimension of about 9Okm. There are approximately 3,000 springs in 100 clusters (Allen and Day, 1935). Many of the clusters are grouped into

490 M.R. WALTER

Fig. 1. Small mud cone with mud flows, Shoshore Geyser Basin. Scale 10 cm.

“geyser basins”, which are areas of intensive geothermal activity. These cover areas of several square kilometres. They are of three types: alkaline (carbonate- or silicadepositing) and acid. The hot-spring travertine of the Mammoth Geyser Basin (Weed, 1889a; Allen and Day, 1935) consists of aragonite (Friedmann, 1970); this basin will not be mentioned further here. Very little deposition of silica occurs from acid waters such as those of the Norris Geyser Basin (rarely, acid springs do deposit geyserite; see Walter, Ch. 3.3). Rock decomposition caused by the acid waters produces surficial deposits of clay and siliceous mud, and in extreme cases of low water flow, “mud pots” and mud cones result (Allen and Day, 1935; Fig. 1). Other minerals deposited by the acid waters include sulphur, pyrite, barite and alunite (Allen and Day, 1935). The remainder of this chapter is concerned with the silica-depositing alkaline hot springs.

DISTRIBUTION OF SEDIMENTS RESULTING FROM ALKALINE HOT-SPRING ACTIVITY

The broad features of hot-spring sediment distribution in Yellowstone National Park are shown on the U.S. Geological Survey Miscellaneous Geo- logic Investigations Map 1-710 (Richmond et al., 1972). On a larger scale, the distribution is illustrated by Fig. 2, an aerial photograph of the Fairy Creek


Fig. 2. Oblique aerial view looking SW over the Firehole River to the Fairy Creek meadows. The light grey patches in the meadows are areas of geyserite and sinter around hot springs; the darker grey areas of the meadows are marshes underlain by diatomite.

meadows in the Lower Geyser Basin. In this photograph the light grey areas in the meadows are deposits of siliceous sinter around active and inactive hot springs, and the dark grey areas in the meadows are marshes underlain by diatomite. Fig. 3 is a low-level aerial photograph of “Conophyton Pool” in the Fairy Creek meadows (see Walter et al., Ch. 6.2). The sediment types surrounding and derived from springs such as this are listed and briefly described below :

(1) Geyserite is deposited in or from water hotter than about 73°C. (2) Very high-temperature stromatolitic sinter forms in water with the

(3) High-temperature stromatolitic sinter forms in water with the approxi-

(4) Medium-tempemture stromatolitic sinter forms in water cooler than

(5) Sinter breccia forms in the well-drained interfluves between spring out-

( 6 ) Marsh deposits form in poorly drained areas; water temperatures prob-

(7) Stream sediments form in the cool streams draining the marshes and

(8) Lacustrine sediments form in cool lakes.

approximate temperature range of 59- 73°C.

mate temperature range of 30-59°C.

about 30°C. Lower temperature limit unknown.

Bow channels.

ably less than about 25°C.

hot springs.

492 M.R. WALTER

Fig. 3. Very low-level high-angle oblique aerial view of “Conophyton Pool”, Fairy Creek meadows. Two almost contiguous kidney-shaped hot pools at the source of the spring are visible to the lower left of centre. There is no geyserite in this spring but it is surrounded by a large area of stromatolitic sinter.

In addition, extensive hydrothermal alteration of the rocks through which the spring waters pass is caused by both the alkaline and the acid waters (Fenner, 1936; Honda and Muffler, 1970).

Geyserite

Geyserite consists largely of amorphous silica. Its morphology, microstructure and genesis have been described and discussed elsewhere (Walter, 1972b; Ch. 3.3 herein; see also White et al. 1964). It is deposited as a result of the cooling and evaporation of nearly sterile waters hotter than about 73’C. Consequently it forms only in and immediately adjacent to geysers, hot springs and outflow channels. Bacteria rarely accumulate in these hot waters to form stringy masses that can be silicified.

Very high-temperature stromatolitic sin ter

Alkaline waters with the approximate temperature range of 59-73’C are abundantly populated by filamentous bacteria, particularly the motile


Fig. 4. Very high-temperature stromatolites in Brock’s Pool A, near Great Fountain Geyser, Lower Geyser Basin. These stromatolites are built by the photosynthetic bacterium Chloroflexus aurantiacus and the unicellular cyanophyte Synechococcus sp. ; the stromatolites are finely laminated. Scale 10 cm.

photosynthetic bacterium Chloroflexus aurantiacus Pierson and Castenholz, and also by various species of the unicellular cyanophyte (“blue-green alga”) Synechococcus. These form laminated stratiform and, infrequently, nodular mats (Figs. 4, 5; Doemel and Brock, 1974), which trap detrital particles (these mats may also be silicified but this has not yet been established). The 59OC isotherm is usually within several metres of the spring and geyser vents so this sediment type is areally very restricted. Doemel and Brock (1974) state that these mats form at temperatures as low as 5OoC, but most, if not all, microbial mats in waters cooler than 59OC contain filamentous cyanophytes in abundance, and belong to the next category.

High-temperature stromatolitic sin ter

Filamentous cyanophytes first appear in microbial mats in Yellowstone at about 59OC. Mats in the range 59-30°C are dominated by the finely filamentous cyanophyte Phormidium (Copeland, 1936) and take on a variety of macroscopic forms (Weed, 1889a, plates 82-87, fig. 56; Allen and Day, 1935, figs, 37, 38; Walter et al., Ch. 6.2 herein). In contrast to the very high- temperature mats, and also to most of the medium-temperature mats, these

494 M.R. WALTER

Fig. 5. Close-up view of part of the area shown in Fig. 4. Each scale division 1 cm.

frequently have structures with a growth relief of several centimetres or more (including the columnar stromatolites Conophyton weedii and Vacerrilla walcotti described in Ch. 6.2). The high-temperature stromatolitic sinter is highly porous after silicification and decomposition of the organic matter.

Medium-temperature stromatolitic sinter

Microbial mats in water cooler than about 30°C in Yellowstone are dominated by coarsely filamentous cyanophytes such as Calothrix. They usually are stratiform, with less than l c m of surface relief (Figs. 6, 7), and are heavily silicified, forming firm, felt-like sheets (Fig. 8 shows silicified Calothrix filaments in one of these mats).

Sin ter breccia

Well-drained, raised spring and geyser mounds are coated with a gravel of tabular fragments of opaline silica. The fragments are of millimetre to centimetre size. They'are often cemented by opdine silica to form an indurated rock. The breccia is derived from the dehydrated sinter of abandoned outflow channels and terraces. Brecciation is due to trampling by bison and deer and in some cases possibly to freezing of interstitial water during winter. The


Fig. 6. One of the outflow channels of “Column Spouter”, Fairy Creek Meadows. The small, dark grey, anastomosing terraces are built by a mat of the cyanophyte Calothrix coriacea, in the cooler parts of the outflow. To the left of the terraces the water is hotter and here small columnar stromatolites form. (See also Ch. 6.2.)

Fig. 7. Looking down on the surface of a silicified mat of Calothrix coriacea, “Column Spouter”, Fairy Creek meadows. Scale 10 cm.

496 M.R. WALTER

Fig. 8. Thin section perpendicular to the lamination of a silicified mat of Calothrix coriacea. The coarse, silicified filaments of the cyanophyte are clearly visible. The lamination is supradiurnal, perhaps seasonal. “Column Spouter”, Fairy Creek meadows. Scale bar 2 mm.

various forms of stromatolitic sinter and sinter breccia commonly form thin beds (Fig. 9) in which lateral facies changes occur frequently.

Marsh deposits

Many Yellowstone hot springs and geysers are in marshes and meadows (e.g., Fairy Creek meadows, Sentinel Meadows). According to Weed (1889c), the marshes support an abundant diatom flora, producing a sediment which in some cases is almost pure diatomite. Weed states that the diatom beds cover many square miles and are up to 6 ft. (2 m) thick.

Stream and lacustrine sediments

I know of no study of the stream sediments in Yellowstone National Park. The sediments of Yellowstone Lake, which is fed by thermal waters, are unstratified, contain more than 80% (by weight) of diatoms, and are moderately rich in organic carbon (Goodwin and Schmit, 1974). The sediments in Heart Lake (Fig. 10) may be similar.


Fig. 9. The wall of Excelsior Geyser Crater, Midway Geyser Basin, showing thin-bedded stromatolitic sinter and sinter breccia.

Fig. 10. Hot-spring draining into Heart Lake.

498 M.R. WALTER

CHARACTERISTICS OF THE HOT-SPRING FACIES

Former areas of alkaline hot-spring activity can be expected to show the following characteristics: (1) presence of volcanic rocks; (2) extensive hydrothermal alteration of host rocks; (3) presence of geyserite (its distribution around vents may be recognizable in wellexposed sequences); (4) presence of other siliceous sediment types, particularly stromatolites and diatomite; and (5) limited areal distribution.

ACKNOWLEDGEMENTS

The U.S. National Park Service permitted field work in Yellowstone National Park. Professor T.D. Brock provided laboratory facilities in West Yellowstone. Financial support was provided by the Department of Geology and Geophysics, Yale University.

9. APPLICATION OF RECENT MODELS TO THE GEOLOGICAL RECORD

Chapter 9.1

THE EFFECTS OF THE PHYSICAL, CHEMICAL AND BIOLOGICAL EVOLUTION OF THE EARTH

Conrad D. Gebelein

INTRODUCTION

The application of Recent models of stromatolite growth and distribution to the geologic record depends in large measure on how closely we can equate modem environments to ancient ones. Where this analogy cari be made confidently, Recent models may be applied directly to the interpretation of ancient stromatolites. Many physico-chemical and biologic attributes clearly have not been the same throughout geologic history. Gradual changes in many environmental parameters have taken place since the origin of the earth. Some of these changes are related strictly to physical changes in the earth, such as crust formation and continental drift, and to chemical changes, such as weathering of rock and transport of ions and solids to the growing oceans. The vast majority of changes, however, are inextricably tied to the evolution of life and the effects of this evolution on nearly every aspect of the earth’s surface. Thus, the attempt to divide this chapter into two ‘parts - effects of physical and chemical evolution of the earth, and effects of biological evolution on the earth - is arbitrary, but hopefully will make for the clearest presentation. Where physiological evolution of the biota had important effects on the physico-chemical properties of the environment, these are discussed first with physical evolution, and then the implications of these changes on biotic evolution are considered.

EFFECTS OF PHYSICAL AND CHEMICAL EVOLUTION OF THE EARTH

Physical attributes which were distinctly different from those of the modem earth are confined largely to the earliest phases of the earth’s history, phases during which either no stromatolites were formed or none have been preserved. The period of the earth’s history prior to the great thermal episode of 3.5-3.7 b.y., termed by Cloud (1974) the Hadean, is extremely poorly represented in rocks preserved today. We know nothing about the

500 C.D. GEBELEIN

atmosphere, hydrosphere or biosphere which may have existed during that era. Whatever those systems may have been like, their influence on subsequent earth history probably was small, as it is unlikely that any hydrosphere, atmosphere or living organisms could have survived the intense heating associated with the close of the Hadean.

The oldest preserved sedimentary rocks are included by Cloud (1974) in the Archean, which lasted from 3.4 to 2.6 b.y. The physical character of the earth’s surface then was very different from any later period (Cloud, 1972, 1974). Sediments were generally immature, and consisted largely of volcanic materials. These sediments were deposited in subsiding belts between lighter plutonic rocks. The crust, where differentiated at all, was very thin and weak. Cratonal deposits, as known from later geologic periods, were very rare during the Archean, as were environments for stromatolite development during this era: shallow platform areas and even continental slope-type environments were rare or very unstable tectonically.

It is clear from several lines of reasoning, however, that photosynthesis and the ancestral blue-green algal floras evolved during the Archean. During this era, lack of suitable physical environments was the controlling factor in the scarcity of stromatolites, and probably also in the evolution of the stromatolite-forming microbiotas. Nonetheless, the oldest recorded stromatolites are known from the Late Archean: the Bulawayan forms of southern Africa (Macgregor, 1940), now dated at 3.1 lo9 years (Bond et al., 1973). These forms may have developed on one of the small shallow pinnacles of cratonal material which originated during this era.

From the beginning of the Early Proterozoic, or the Proterophytic of Cloud (1974; dated by Cloud as 2.6-2.0. b.y.), the physical character of the earth’s surface takes on a nearly “modern” aspect. Widespread cratonal sediments are known from this era, including shallow marine carbonate deposits, as well as continental deposits of fanglomerates. The presence of these deposits, although generally unoxidized, indicates the development of widespread stable crustal regions on the earth. Studies of the margins of the crustal platforms (Hoffman et al., 1974) present during this era indicate that continental-drift mechanisms similar to those operative today were present in the Proterophytic. The implications of this “modem” aspect of Protero- phytic and later eras are: (1) physical environments for development and proliferation of stromatolites and stromatolite biotas were available, both on the platform and on the slope; (2) geographic mechanisms controlling distribution of stromatolite biotas, as studied in the Recent (Gebelein, 1974), should be applicable to rocks younger than 2.6 b.y.

The development of continental and shallow marine platforms set the stage for the development of stromatolite-building communities. Indeed, all available slope, platform-margin and platform-interior environments were colonized by stromatolite-forming biotas in the Slave Province (Great Slave Lake, Canada) rocks of approximately 2 b.y. age (Hoffman, 1974a). While

APPLICATION OF RECENT MODELS 501

the physical character of the earth’s surface never again became a limitation to stromatolite development in toto, the abundance of stromatolite forms, and very likely the abundance of microbiotas controlling stromatolite structure, did change throughout the remainder of geologic time in response to the availability of shallow-platform and slope environments (i.e. in response to tectonic uplift and subsidence, or eustatic sea-level changes). This relationship between stromatolite distribution through time and the presence of platform and shallow-slope environments has been noted by Semikhatov (1966) and Gebelein (1974) for the Precambrian, and by Halley and Eby (1973) for the Cambrian and Ordovician. These changes in the availability and distribution of stromatolite environments provide a first-order control on the distribution and abundance of stromatolites from the Proterophytic onward. Superimposed on this first-order control are the variations and changes in the chemical environment and in the composition of the biosphere.

Of critical importance are the chemical changes in the hydrosphere and atmosphere which occurred during the Precambrian. Paleozoic chemical fluctuations probably also have occurred, but on a much smaller scale. These chemical changes were brought about largely by biological activity, and this activity, in turn, influenced the nature of subsequent organic evolution. In terms of the effects of chemical evolution on stromatolites and stromatolite environments, the two most important general topics are the influence of the chemistry of the hydrosphere on carbonate precipitation, and the direct effect of hydrospheric and atmospheric evolution on the development of new stromatolite-forming biotas.

It now seems likely that the earliest life, for which we might expect to find remains, evolved in a C02, N2-rich atmosphere in the near absence of free oxygen (Rubey, 1951; Holland, 1962). That an extensive hydrosphere was generated within this milieu is indicated by the widespread occurrence of water-transported sediments in the Early and Middle Archean. The earliest organisms were anaerobic. However, several lines of evidence point to a very early development of photoautotrophy in this environment: 13C/12C isotopic ratios of the Swaziland System (approximately 3.4 b.y.; Hurley et al., 1971) show very heavy values for the lower Onverwacht Group, but much lighter values for the middle and upper Onverwacht. Oehler et al. (1972) surmise that this change may represent the evolution of a photosynthetic system for partitioning of the isotopes. Cloud (1972) hypothesizes that the earliest O2 -releasing photoautotrophs would lack the advanced internal oxygen- mediating enzymes, and hence would require a local external source as an oxygen acceptor to maintain low O2 concentrations. He postulates that ferrous iron, then in large concentration due to the low oxygen values, would serve as this acceptor. Evidence for this oxygen-acceptor role would be the presence of banded iron formations (BIF). Since the earliest (although very local) BIF are older than 3.76 b.y. (Moorbath et al., 1973), Cloud

502 C.D. GEBELEIN

suggests that the origin of oxygen-releasing photosynthesis must have preceded or been coincident with these oldest BIF.

The most important consequence of the absence of free oxygen from the atmosphere and hydrosphere during Archean and Proterophytic times was the absence of the ozone shield against UV-radiation. In the absence of this shield, primitive microorganisms, both bacteria and algae, must have adapted some mechanisms or strategies for shielding the UV. Data on modern blue- green algae indicate that several intracellular and extracellular adaptations still exist to cope with UV irradiation: DNA repair mechanisms are widespread in blue-green algae (Monty, 1971; Cloud, 1972); photoreactivation mechanisms for photosynthesis are known (Van Baalen, 1968); and both intracellular and sheath pigments are known to be efficient UV-absorbents (Monty, 1971). It is certainly likely that such intracellular and extracellular shields were equally well developed in the Early Precambrian. In fact, Jackson (in Cloud, 1974) indicates that the concentration of UV-absorbing phenolic compounds in sediments increases through the Archean and Pro- terophytic until about 2 b.y. ago (initial development of the ozone screen at that time?), and then declines in younger rocks.

The importance of the efficiency of UV-shielding lies in the effect of UV- radiation on habitat availability for stromatolite-forming biotas. Water is a very efficient UV-absorbent, and below several meters it is unlikely that UV-radiation could have effected the viability of microbiotas, even in the total absence of ozone. Thus, Early Precambrian microbiotas capable of forming stromatolites could have survived in shallow marine and freshwater environments. But could stromatolites have formed in intertidal environments? Several investigators have postulated that stromatolites in, intertidal environments could form where the algae or bacteria were protected from UV by detrital or precipitated sediment (Fischer, 1967; Monty, 1971). Unfortunately this argument fails, as it is clear that at some often repeated intervals, the algae (or bacteria) must grow to the very surface of the structure to provide the preserved rhythmicity of the laminations. Thus, it is unlikely that intertidal stromatolites could form until ozone and biogenic UV-screens were sufficient to ensure cell viability. Several lines of evidence provide data on the timing for the appearance of intertidal forms. Direct evidence would appear to come from the Great Slave Lake section, where several distinctive stromatolite morphologies and associated sediments could only have formed in periodically exposed positions (Hoffman, 1974). These sediments date from 1.7 to 2.2 b.y.

An excellert study of the greater than 2.2 b.y. old Transvaal Dolomite (Eriksson and Truswell, 197413) documents the occurrence of intertidal stromatolites and compares them with modern forms in Shark Bay, Australia. Indirect evidence may be obtained from Cloud’s (1972) argument that the tremendous abundance of BIF at the end of the Proterophytic heralds the development of more efficient intracellular oxygen-mediating enzymes, and


thus the release of free oxygen to the hydrosphere and atmosphere. Within a very short interval, approximately 2.2-1.9 b.y., all of the ferrous iron was swept from the oceans, and free oxygen began to accumulate rapidly. It is unlikely that UV-radiation would have posed any restriction on the habitat (both intertidal and subtidal) availability for stromatolites in younger eras. Thus, I would conclude that all physical habitats in which stromatolites could possibly develop were present and ready for colonization by the beginning of the Proterozoic (Cloud’s 1974 definition, 2.0 b.y.). The arguments above indicate that both the potential settings for stromatolite development, and the nature of stromatolite-building microbiotas, probably underwent significant changes at the end of the Archean and at the end of the Protero- phytic. The periods before and between these episodes probably were marked by much slower change in the physical environment and in the evolution of biotas. Following the beginning of the Proterozoic, the potential physical environments for stromatolite development have a distinctly “modern” aspect. Other controls, mostly biologic, did, however, affect importantly the colonization of these potential environments.

It is suggested from the above that changes in the oxygen content of the atmosphere, per se, had little effect on stromatolite biotas and stromatolite structures in the Proterozoic and later world. One possible effect, however, relates to the apparently rapid increase in oxygen content which may have occurred near the end of the Precambrian, and which continued into the Early Paleozoic. Regardless of the reason for the oxygen increase (rapid tum- over of the seas and oceans, releasing vast quantities of nutrients and hence stimulating photosynthesis is the most frequent suggestion: Cloud, 1972; Schopf et al., 1973), oxidative reactions would become more prevalent following this increase. Surface and near-surface organic-rich sediments (such as stromatolites), which until that time would have been reduced, would now undergo much more rapid and complete oxidation. Such oxidation would have the effect of removing organic material and very likely (as noted by Golubic, 1973a, for modern environments) would lead to the destruction of stromatolites. This oxidative destruction may be an important reason for the rapid decline of stromatolites in the Early Paleozoic. Such effects would be minimal in very hypersaline environments, where aerobic bacterial action is much reduced (S. Golubic, personal communication, 1973). Changes in the ionic composition of seawater could have important consequences for the types of stromatolites which can form. Changes in parameters related to carbonate saturation, and hence precipitation of carbonate in stromatolites, would be particularly important. Most workers agree that the oceans have always been “salty” (Rubey, 1951; Holland, 1965), although the concentration limits of the individual ions are hard to estimate. Calculations of the steady-state Precambrian ocean yield a seawater composition quite similar to that of the present oceans (Lafon and Mackenzie, 1974). The work of Garrels and Mackenzie (1974) shows that the differences between modern

504 C.D. GEBELEIN

sediments and ancient rocks can be explained largely as the result of post- depositional changes. Their data suggest that the composition of seawater has been nearly constant for at least the last 2 b.y. Further, they suggest that maximum variation in the concentrations of the elements with long life- times in the ocean (Cl, S, Na, Mg) has been * 10% of their present values. Holland (1972), after considering the mineralogy and chemistry of Archean (Fig Tree) and later sediments, concludes that differences in these parameters from modem values are very small, and in particular that silicate mineralogy (and hence seawater composition and buffering) has been very similar to present seawater chemistry for most of geologic history. Examination of ancient evaporite sequences and consideration of the mineral equilibria correlated with these sequences led Holland (1971, 1974) to suggest: (1) that the possible ranges of concentration for Ca and HC03 (since the first appearance of gypsum as a common sedimentary mineral) are-Ca, 1-30 mmol/kg of seawater (present value, 10.3 mmol/kg), and HC03, 0.1-20 mmol/kg (present value 2.3 mmol/kg); (2) that the mineralogy of the evaporites associated with the Bitter Springs deposits (800-900 m.y. old) are so nearly “modem” in aspect that no profound changes in seawater chemistry could have been associated with the very profound biotic evolution after Bitter Springs time.

Other types of chemical changes through geologic time are more difficult to assess. I t is clear that the Early Precambrian ocean, at least until the end of the BIF era (about 1.9 b.y.), was saturated with respect to silicic acid. Continuous precipitation of silica is favored at relatively low pH’s, about 6.8-7.5 (Holland, 1971). Using a rather different approach to the state of the Precambrian ocean, Lafon and Mackenzie (1974) made computer calculations on the state of the inorganic chemistry of seawater during the reactions between igneous rock and the outgassed materials. This calculation (which is based on “instantaneous” outgassing of all materials) yields silica saturation at a pH of 7.66 on conclusion of the reaction. The importance of these calculations is that it is unlikely that seawater pH was ever higher than at present, and, in fact, may have been somewhat lower during most of the Precambrian.

The import of the above discussion may be summarized as follows. First, at least since 2 b.y., it is unlikely that major changes in seawater chemistry (except redox reactions) have occurred. Thus, this factor may be eliminated as a possible difference between modem and ancient environments of stromatolite formation. Secondly, it appears that carbonate equilibria have not changed markedly since at least the beginning of the Proterophytic. This is an important statement, since much emphasis has been placed on the importance of carbonate precipitation in the formation of Precambrian and Paleo- zoic stromatolites (Monty, 1973a; Serebryakov and Semikhatov, 1974). Serebryakov and Semikhatov cite several lines of evidence to indicate a difference in cementation between Precambrian and later stromatolites :


(1) they have found calcareous stromatolites in “marine” non-carbonate rocks; (2) they have observed limestone-dolomite lamination alternations in stromatolite bioherms, and not outside the bioherms; (3) they commonly find differences between the stromatolites and surrounding sediments in terms of mineralogy and insoluble residues. Alternative explanations can be offered for each of these occurrences: (I) there is seldom, if ever, unequivo- cable evidence for the marine nature of the calcareous bioherms - precipitation in freshwater is equally likely (similar structures form in modern freshwater environments); (2) the presence of limestone-dolomite alternations is evidence only for the alternation of algal-rich and carbonate (either precipitated or detrital)-rich layers, not necessarily for precipitation; (3) differences in mineralogy and insoluble residue between stromatolite bioherms and surrounding detrital sediments would be expected also in the complete absence of cementation within the bioherm.

Is there any evidence which may be used to establish the presence of precipitated stromatolites in the Precambrian? It is clear that many Precambrian and Paleozoic stromatolites lack evidence for preserved detrital carbonate grains (although detrital quartz is common in these and most other stromatolites). This fact in itself may not be conclusive. The cemented modem stromatolite columns at Shark Bay (Logan et al., 1975), as well as some of the uncemented forms, show very little or no evidence for preserved detrital grains, even though these stromatolites were formed entirely by grain trap- pings. Micritization of grains is very common in the hypersaline, “beach- rock” environment of Shark Bay. Perhaps some of the Precambrian stromatolites lacking detrital texture also originated in this way. However, in some instances it is possible to show that several Precambrian stromatolites have preserved detrital textures (e.g., Tungussia) and occur in the same section as other forms which lack detrital textures, for example Baicalia or Inzeria (Cloud and Semikhatov, 196913; Walter, 1972a). This situation appears to indicate that precipitation does occur in some stromatolite forms. Precipi- tation of crudely laminated carbonate structures in a terrigenous marine slope deposit is documented by Hoffman (1974). The origin of these presumably stromatolitic structures is not understood, but may be related to non-photosynthetic algae or bacteria. Another line of evidence for precipitation in Precambrian forms derives from the physical constraints on stromatolite height. Evidence from the Recent, and simple intuitive reasoning, argue that stromatolites with greater than about lOcm of relief must have been cemented at the time of their formation. It does not appear feasible for structures of such a size or greater to s u p p ~ r t their own bulk without the addition of precipitated cement. While this would eliminate the necessity of cementation for most Precambrian forms, it would certainly ensure the cemented nature of certain forms, such as some of the conophytons.

It is interesting to note that, with the exception of one particular environmental setting, stromatolites with high relief disappear at the end of the

506 C.D. GEBELEIN

Precambrian. This change may be related to a subtle, but important, shift in carbonate equilibria associated with the development of calcareous metaphytes and metazoans. All the chemical and mineralogic evidence indicates that carbonate equilibria in the Precambrian could not have been very different from the present. The level of supersaturation must have been somewhat higher, in order that extracellular carbonate precipitation could occur in any blue-green algae (intracellular precipitation is unknown in any procaryote). It seems extremely unlikely that any physiological change has occurred since the end of the Precambrian which would make the blue-greens less able to precipitate carbonate in the marine environment (i.e. reduced photosynthesis since the Precambrian). Lack of precipitation in Paleozoic forms must then be related to the slightly higher level of carbonate saturation which existed during the Precambrian. The ability of only some stromatolite-building biotas to precipitate must have been related to very subtle differences in photosynthetic rate, or to parameters such as sheath consistency or thickness (which would control the rate of degassing away from the cells). I postulate, then, that the absence of carbonate precipitation in Paleozoic and younger stromatolite-building biotas is related to a slight lowering of the carbonate supersaturation due to removal of bicarbonate ion by metaphytes and metazoans - not, as Monty (1973a) argues, to the evolutionary disappearance of stromatolite-building biotas with the inherent ability to precipitate carbonate (Golubic, 1973a, indicates that carbonate precipitation is not a genetic feature of any blue-green algae)*.

Thus, I would conclude that cementation of stromatolite forms was common in normal marine environments during the Precambrian, a situation very different from the present oceans. Widespread cementation probably ceased near the end of the Precambrian, and thus would have eliminated the potential to build certain stromatolite forms. High-relief stromatolites do continue into the Paleozoic, although in much reduced abundance and diversity. A comprehensive literature study by Halley and Eby (1973) indicates that Paleozoic high-relief stromatolites are invariably associated with evidence for hypersalinity. The lines of evidence they use for hypersalinity are: casts or pseudomorphs of gypsum, anhydrite or halite; nodular or bedded evaporites; length-slow chalcedony ; collapse features (especially oolites) ; early dolomite; restricted faunas. Their data, which extend to the Permian, indicate that high-relief forms were generated in environments very similar to Shark Bay, where cementation is basically an inorganic process (an additional easily recognized environment of Phanerozoic high-relief stromatolite formation is the forereef, described below). I would suggest that careful study of Paleozoic and younger high-relief stromatolites m,ay support a direct cokrelation with modem occurrences of high-relief forms (i.e. a Shark Bay setting).

* For evidence that carbonate precipitation is a characteristic of particular species see Chapter 8.6 (Editor).


EFFECTS OF BIOTIC EVOLUTION ON THE RECENT-ANCIENT COMPARISON.

Perhaps the easiest way to begin this discussion would be to very briefly summarize the modern state of stromatolite-building biotas: (1) blue-green algal stromatolites are basically an upper intertidal phenomenon in normal marine waters; subtidal blue-green algal stromatolites occur only in environments which, for various chemical (hypersalinity) or physical (sediment movement) reasons lack invertebrate grazers and burrowers; blue-green algal stromatolites are common in freshwater settings, both periodically exposed and submerged (Monty, 1972); bacterial stromatolites occur in environments which lack filamentous blue-green algae (thermal springs, deep sea) and which have very few burrowers and grazers (Walter et al., 1972; Monty, 1973a). These particular biologic and environmental groupings have not been the same throughout geologic history. The most likely parameters to affect this modern situation are the physico-chemical changes discussed above, and the evolution of various life forms which either form or destroy stromatolite structures.

It is clear that the earliest microorganisms must have been non-photosynthetic heterotrophs, similar in some respects to modern bacteria (Cloud, 1974). The revelation that heterotrophic bacteria have the potential to make stromatolites (Monty, 1973a) in the deep sea implies that stromatolites could have formed as soon as the earliest sediments and hydrosphere were formed. Such organisms would, of course, lack any phototactic response as the inherent cause for upward growth within the stromatolites. It is likely, however, that they would exhibit a chemotaxis, such that continuous or periodic growth toward the upward surface would occur. A t present, such non-photosynthetic forms are excluded from constructing stromatolites in all lighted portions of the hydrosphere by the much more efficient photosynthetic bacteria and algae. I t is certainly possible that heterotrophic forms existed in the lighted portions of the oceans during the Hadean and Early Archean. It is also possible that such forms played an important role in the unlit slope areas surrounding platforms in the Archean and younger eras. Perhaps some of the stromatolitic and other “organic” structures reported by Hoffman (1974) from the slope areas of the Great Slave Lake region were formed by such organisms. Occupation of the slope areas, first by heterotrophic bacteria and then by photosynthetic bacteria, could have continued until the colonization of these areas by invertebrate grazers and burrowers in the Paleozoic. Such deeper-water biotas and settings would probably generate microstructures and macrostructures with some characteristics unique to that environment.

Sedimentologic information on the rocks associated with the vast majority of stromatolites indicates that growth probably took place in relatively shallow water, almost always within the photic zone. For that reason, the evolutionary history of the photosynthetic biotas is of greatest importance

508 C.D. GEBELEIN

for our consideration. While this topic is considered in detail in another chapter, it needs to be reviewed here and incorporated into our theme. Several lines of evidence indicate that photosynthesis evolved during the Early to mid-Archean. The change in 13C/’*C ratios from the lower Onver- wacht to the upper Onverwacht Group (Oehler et al., 1972) is presumptive evidence for the evolution of photosynthesis prior to 3.1 b.y. The preserved microfossils of the Fig Tree Group (Barghoorn and Schopf, 1966) include Archaeosphaeroides barbertonensis, a form which has its closest affinities with modern coccoid blue-green algae. No filamentous forms are known from these rocks - an important consideration as well-laminated stromatolites apparently are made only by filamentous forms. However, presumptive evidence for filamentous forms is found in the Bulawayan stromatolites (about 3.1 b.y.), which also show evidence for photosynthetic carbon fractionation (Hoering, 1967). An important point for our Recent-ancient comparison is: are these early forms bacteria, algae, or both?

The carbon-isotope ratios cited above imply photosynthesis, but data are not available on carbon fractionation by photosynthetic bacteria. Assuming that bacteria fractionate carbon in much the same way as blue-green algae, what evidence can we use to distinguish these two groups? The very large size of the coccoid forms of the Fig Tree Group argue against bacteria, and for blue-green algae. Other forms in the Fig Tree very likely are bacteria. Other evidence is indirect. If Cloud’s hypothesis for the origin of BIF is correct, then oxygen-releasing algal photosynthesis originated as early as 3.76 b.y. However, it seems unlikely that these earliest algal forms, with their strong dependence on the availability of iron in solution, formed the stromatolites in the Early and mid-Archean. The best strategy for rapid transfer of gaseous oxygen to a local acceptor would be for the cells to be suspey ded in a medium which contained the acceptor - i.e. for the algae to have a ,lank- tonic habit. Benthic forms, incorporated into a detrital deposit or a precipitated fabric, would have far less water exchange and thus a far lower ability to transfer oxygen to the acceptor. Thus, dependence on an external oxygen acceptor may have limited the stromatolite-forming capabilities of algae in the Archean, and left the role of stromatolite-building largely to the photosynthetic bacteria. Evolution during the Archean probably was fairly slow, with numerous attempts at new physiological and biochemical pathways. However, the number of phenotypic possibilities remained small, and the number of possible photo- and chemo-tactic responses was also small. Thus, the nuniber of stromatolite microstructures and forms was limited in the Archean, and changes through time were few and far between.

This situation would change abruptly, and for all time, with the evolution of the advanced oxygen-mediating enzymes of photosystem 11, when more efficient “modern” blue-green algae would have excluded the photosynthetic bacteria from nearly all environments. The timing of this event may be related, as postulated by Cloud (197313) to the period of abundance of BIF


toward the end of the Proterophytic. Increased BIF was initiated by glacially related turnover of ocean waters, releasing large quantities of iron and nutrients to the photic zone. Increased productivity led to increased numbers of blue-green algae, and with this increase in numbers came an increase in mutation rates which led to the evolution of the advanced oxygen-mediating system. During this period, oxygen was released to the hydrosphere and atmosphere in larger quantities than ever before, sweeping the iron from solution forever. The relatively short duration of this interval probably means that: (1) evolution of blue-green algae, with nearly modern phenotypic and biochemical aspects, tQok place at a very rapid rate; (2) coincident with this rapid evolution, and in part responsible for its continuation, would be the opening of many new habitats for blue-green algal colonization as a result of the freedom from dependence on soluble iron. Blue-green algae would have become the dominant benthic form in lighted portions of the hydrosphere (with the possible exception of the very deep photic zone, where slow water turnover might limit oxygenation, and allow the photosynthetic bacteria to continue their dominance for some time).

This sudden burst of evolutionary activity led to the development and diversification of most of the important stromatolite-building blue-green algae known from modern environments. Evidence for the existence of these blue-green algae comes chiefly from the Gunflint Iron Formation (Barghoorn and Tyler, 1965; Cloud, 1965; dated 1.6-2.0 b.y.) and from the lower Belcher Group (Hofmann and Jackson, 1969; dated at 1.7 b.y). These floras contain numerous examples of coccoid, oscillatoriacean, and nostocacean blue-greens. Thus, the strong possibility exists that, whereas stromatolite- building communities and their distribution may have been different in Arche-r and Proterophytic times from now, analogies may be more confi- dentl: drawn between the stromatolite-building biotas and distributions of Proterozoic and younger times with the Recent. The following analogies should be possible: (1) the types of blue-green algae, and their associations and communities, should be similar; (2) controls on the distribution of blue- green algal species and communities should be similar; (3) phototactic and chemotactic responses should be similar. As discussed in the first section, the Proterozoic marks the onset of physical and chemical conditions quite similar to the Recent. Thus, we should be able to deduce the biotic and environmental significance of a great many of the Proterozoic stromatolites by comparison with modern algal assemblages and modern environments of stromatolite formation.

I recently made an attempt (Gebelein, 1974) t o explain the changes in stromatolite form and microstructure through the Proterozoic (the Riphean of the U.S.S.R.) on the basis of analogy of Proterozoic microstructure with Recent microstructure and associated biotas, and on the basis of modem geographic distributions of blue-green algal biotas. In general, this idea works well, within the Proterozoic itself. Major changes in stromatolite forms could

510 C.D. GEBELEIN

be correlated with the periods of recolonization of continental platforms following periods of tectonic or eustatic regression. Monty (1973a) also considers this same factor, although from a different angle, to be important in the development of the Precambrian stromatolite sequence.

In my formulation, I certainly did not rule out continued evolution of the blue-green algae as a potential factor in stromatolite change through time. It seems assured that slow, continued evolution of the blue-green algae occurred during the Proterozoic. However, as during the Archean and Pro- terophytic, selection pressures toward development of innovative physiological and phenotypic traits in the blue-green algae were relatively low, at least until the very late Proterozoic. The fossil record apparently bears out this hypothesis. Comparison of the Gunflint biota with the biotas from the Beck Spring Dolomite (Cloud et al., 1969; Gutstadt and Schopf, 1969), the Belt Supergroup (Pflug, 1964, 1965, 1966), the Skillogalee Dolomite (Schopf and Barghoom, 1969), and particularly with the Bitter Springs flora (Schopf, 1968, 1972; Schopf and Blacic, 1971) indicates: (1) a high level of phenotypic conservatism within the blue-green algae of various ages; (2) an apparent, though not conclusive, increase in the diversity of form within the Oscillatoriaceae and the Nostocaceae (and the presence of many different taxa in the different deposits); and (3) the apparent first appearance of the complex blue-green algae, the Rivulariaceae, by 1.0 b.y. (although uncertain occurrences of this family may extend back to the Early or Middle Protero- zoic; Schopf, 1970). Within the limits of the above discussion, it can confidently be stated that some of the changes in stromatolites through the Proterozoic were related to algal evolution.

All workers would agree that development of stromatolite form and stromatolite-building biotas reached a peak in the Early to Middle Protero- zoic. The very latest stages of the Proterozoic (less than 800 m.y.) were marked by rapid decline in the abundance and diversity of stromatolites - a situation which can most clearly be related to the development of metaphytes and metazoans (Garrett, 1970b; Awramik, 1971a). Thus far, we have not considered the importance of the absence of metaphytes and metazoans on our comparison between Precambrian and modem stromatolites. The effect of the rise of higher plants and animals (discussed in more detail below) has been to limit the “realized ecologic hyperspace” of stromatolite- building blue-green algae. Prior to this event, blue-green algae and stromatolites occupied a much broader range of environments than they did during the Paleozoic and younger eras. In particular, the subtidal photic zone has often been cited as a prime site of formation for stromatolites (Serebryakov and Semikhatov, 1974), a situation very different from most of the Phanero- zoic. Data from many workers in the Proterozoic indicate that evidence of desiccation is absent or rare in Proterozoic and older stromatolites (cf. Walter, 1972a). Certainly, the size of some of the Proterozoic columnar stromatolites (combined with the rarity of microdisconformities in these


structures) necessitates a subtidal origin for these forms. Numerous workers cite the lateral continuity of individual stromatolite beds or microstructures as evidence for a uniform, therefore subtidal, environment. My own experience in the Recent would indicate that the shallow subtidal also may be exceedingly variable; on the whole, however, I think the weight of evidence certainly favors a subtidal origin for most of the Proterozoic and older forms.

If this is the case, how much Recent data can we apply to the Pre- cambrian? It seems that a great deal of the Recent data can still be applied to the interpretation of Precambrian stromatolite form and microstructure, as long as we keep in mind that formation occurred in the subtidal zone. Blue- green algal associations function in much the same way, whether they are intertidal or subtidal (Neumann e t al., 1970). Mechanisms for precipitation (discussed above) and especially sediment trapping are similar in these two settings. Environmental controls on stromatolite or bioherm morphology would be related to current and sediment transport patterns in the subtidal as well as in the intertidal. Evidence for such controls is apparent in Precam- brian stromatolites (see esp. Bertrand-Sarfati, 1969; Walter, 1972a; Serebryakov and Semikhatov, 1974). Thus, we lose very little interpretive potential from the Recent, as long as we are aware that the analogy is not one-to-one in terms of all environmental parameters. If we knew more about the controls on the distribution of Recent subtidal mat communities, we very probably could extend a much wider range of interpretive data from the Recent to the past. Unfortunately, there are virtually no data of this kind available.

The changes which took place in the biosphere and hydrosphere toward the end of the Precambrian bring the stromatolite picture into a nearly modern focus. They have their origin with the evolution of the eucaryotic cell, dated with some confidence from 1.3 b.y. (Cloud et al., 1969), and possibly older (Hofmann and Jackson, 1969). The effect of these very primitive mitotic algae apparently was small on the stromatolite-building biotas. There is no evidence that this development, per se, limited in any way the environmental range of stromatolite-building blue-green algae or decreased stromatolite diversity during most of the Proterozoic. More important in terms of stromatolite biotas was the development of meiosis and the ensuing development and expansion of metaphytes and metazoans. The earliest evidence suggestive of meiotic cells appears to be in the Bitter Springs rocks (Schopf, 1968; dated at about 800-900 m.y.). The absence of any advanced thallophytes in these deposits led Schopf to postulate that meiosis was quite new at 900 m.y. The development of meiosis would herald an era of tremendous genetic diversification, and the rapid rise of metaphytes. Again, the timing of certain events may help to understand the controls on stromatolite-building biotas. The acme of stromatolite diversity is reached at about 800 m.y., i.e. not until after the development of meiosis (Awramik, 1971a). The decline in stromatolite diversity indicated by Awramik is abrupt in time,

512 C.D. GEBELEIN

and very marked. This decline covers most of the Vendian (680-570 m.y). This interval was marked by widespread glaciation, and this factor may have triggered several events which led to the decline of stromatolite diversity and abundance (Schopf et al., 1973; Cloud, 1974): (1) glacial turnover of ocean and fresh waters would have brought abundant nutrients into the photic zone, where increased productivity and diversification of metaphytes could occur; (2) strong latitudinal gradients in climate would have existed, providing a wide variety of environments for genetic diversification of metaphytes (and metazoans); (3) peneplanation by glaciers would provide wide shallow seas for metaphyte development upon glacial waning; (4) increased productivity associated with the metaphyte bloom would have led to a rapid increase in atmospheric and hydrospheric oxygen - resulting in increased oxidation of organic matter in stromatolites, and of course providing a more hospitable environment for the concomitant development of metazoans.

How was this decline in stromatolites (and presumably in stromatolite- building biotas) related to the development of metaphytes and metazoans? Can we reasonably apply data from the Recent, where interaction between these groups is strong, to the Late Proterozoic setting? Garrett (1970b) suggests, and others have concurred, that the decline in stromatolite diversity may be correlated with the development of burrowing and grazing metazoans. Data from the Recent indicate that distribution of burrowing and grazing animals is the control on distribution of stromatolites in the Recent normal marine environment. This same relationship appears to be the controlling factor in the decline of stromatolites in the Late Proterozoic and in the positioning of stromatolites during the Phanerozoic.

Monty (1973) suggests alternative mechanisms for the Late Proterozoic decline: (1) decrease in atmospheric C 0 2 , which he says would reduce blue- green algal photosynthesis (I suggested above that changes in carbonate ion concentration, hence in cementation, may have occurred due to metazoan and metaphyte hard-part evolution); (2) competition with red and green algae for space and nutrients. Monty sees no evidence for burrowing in stromatolites which would indicate a metazoan control on stromatolite distribution and abundance. I suggest that these points are not valid for the following reasons: (1) Decreased photosynthesis related to decreased C02 is unlikely as a causal factor, as this parameter should affect the photosynthesis of metaphytes as well as blue-green algae. (2) There is no evidence for competition of blue-green algae with red and green algae (but it is not clear what type of evidence could be found). I t seems clear that metaphytes in modern environments are not limiting the development of blue-green algae, although they may limit the development of strornatolites in reef environments (see below): many subtidal mats contain a complex mixture of these various algal types. (3) The appearance of metaphytes may, in fact, have led to the development of new stromatolite microstructures in the Vendian. Walter (1972a) indicates that vermiform structures in stromatolites, whose size implies a


eucaryotic algal origin, do not occur in stromatolites older than Vendian. (4) Absence of abundant burrow structures in preserved stromatolites is exactly what one finds in modem normal marine settings. As discussed in Ch. 8.1, the boundary between modem marine algal stromatolites and burrowed or massive sediments is extremely abrupt. One rarely sees evidence for an intermediate texture. The conclusion remains, then, that the distribution, abundance and diversity of stromatolites from the Late Proterozoic to the present has been controlled, in large part, by the activities of metazoan burrowers and grazers.

Thus, Late Proterozoic and Phanerozoic stromatolites became confined more and more to environments where the activity of metazoans was restricted or eliminated. The inability of invertebrates to survive hypersaline conditions, as in Shark Bay, left this environment open for stromatolite building. The evidence for Phanerozoic high-relief stromatolites is largely confined to such hypersaline environments (see below for reef environments). The rationale for this occurrence is two-fold: (1) invertebrates were eliminated by chemical parameters; (2) cementation, and thus high-relief stromatolite forms, occurs within hypersaline environments because of an inorganic precipitation process. During the Phanerozoic stromatolites also flourished in freshwater settings (cf. Monty, 1972). The rationale for these occurrences is as follows: (1) the level of invertebrate activity is much lower in freshwater than in normal marine waters; (2) the lack of a buffering system in freshwater allows carbonate precipitation to occur as a result of blue-green algal photosynthesis; (3) the ability to precipitate a hard fabric greatly reduces the potential effect of burrowers, and sharply curtails the effect of grazers. A similar set of circumstances may account for the manganous stromatolites reported in the deep sea by Monty (1973a): (1) invertebrate burrowing and grazing are limited in this organic-carbon limited environment; (2) very slow rates of sedimentation allow precipitated fabrics to form; (3) this precipitation, as above, curtails the disruptive effects of burrowers and grazers.

Thus, blue-green algal and bacterial stromatolite-building biotas were restricted to the very margins of their potential ecological hyperspace during most of the Phanerozoic. That this restriction was rapid and (with one important exception) nearly complete by the Ordovician is recorded in the extremely abrupt decrease both in stromatolite diversity and abundance in the rock record. Therefore, the principles of stromatolite distribution derived from the Recent may be applied with some confidence throughout the Phanerozoic, and, in fact, provide valuable data to explain the timing and nature of the restriction of stromatolites at the conclusion of the Pre- cambrian.

An important exception to the direct application of Recent distributions to the Phanerozoic record concerns the role of blue-green algae in biohermal and reef-like deposits. There are many occurrences of stromatolitic bioherms and reef structures in the Precambrian. Stromatolites continued to

514 C.D. GEBELEIN

contribute to reef construction during the Phanerozoic. The occurrence of reefal forms, which are fairly common in the Paleozoic, rare in the Mesozoic, and lacking in the Cenozoic, may be related to the absence of competition and to the possibility of precipitation. Modem reefs, for example, are known to have the highest rates and amounts of inorganic precipitation of all normal marine environments. The hydrostatic pumping of water through the reef allows this cementation. Also, the abundance of burrowing and grazing animals is relatively low (fish probably being the most important surface grazers) on modem reefs. Thus, it would seem likely that reef environments would be the last normal marine environment in which blue-green algae could play a significant rock-building role.

This role has diminished gradually throughout the Phanerozoic. The rationale for this may be as follows. The origin of the invertebrate reef-building organisms would account for the dramatic and rapid decline in stromatolite reef building in the Early to mid-Paleozoic. These invertebrates have a much more efficient system of calcification, and, through their various growth and metabolic adaptations, were able to successfully subdivide the reef-building role into numerous ecologic niches. Blue-green algae were still able to play a part in reef building in environments: (1) for which there were no specialized invertebrates, and (2) in which grazers were limited. Such conditions apparently were met most often in the deeper portions of the reef (as in the Devonian Canning Basin forereefs: Playford and Cockbain, 1969). Gradually, however, these remaining reef niches were occupied by invertebrate reef builders.

The rarity of stromatolites in Mesozoic reefs and their absence in the Cenozoic may be related to the development of the calcareous red algae as a dominant member of the reef-building community. These latter organisms effectively outcompeted the blue-green algae in all reef portions, including the deep forereef (Monty, 1973a). A final factor which may have eliminated forever the possibility of precipitation in marine waters by blue-green algae was the rapid rise of the calcareous plankton during the Mesozoic. The removal of carbonate from solution may have adjusted the carbonate equilibria so as to make it virtually impossible for blue-green algae to precipitate carbonate in normal marine waters.

SUMMARY

The above discussions present, I hope, an optimistic outlook on the application of what we have leamed in the Recent to the problems of interpreting ancient stromatolltes. It is my contention that nearly all of the principles derived from the study of the Recent have important application to the geological record of stromatolites. I hope that I have also emphasized that the present is not the past, and that application of Recent data to ancient rocks


must be done in the full understanding of the state of evolution of the hydrosphere, atmosphere and biosphere at that particular time. The most important of these qualifying conditions are: (1) the importance of bacteria in the Early Precambrian stromatolites; (2) the importance of both environmental (unidirectional and repetitive) and evolutionary factors in explaining the change in stromatolite structures (keeping in mind that these changes should be slow during each of the eras - Archean, Proterophytic, Protero- zoic, and Phanerozoic - and rapid in the transitional periods); (3) the effects of the absence of metaphytes and metazoans on the distribution and diversity of stromatolites in the Early ,and Middle Precambrian; (4) conversely, the effects of the development of metaphytes and metazoans on the Phanerozoic distribution and diversity of stromatolites; ( 5 ) the importance of the development of calcareous red algae in modifying the role of blue- green algae in Mesozoic and Cenozoic reefs. Keeping these points in mind, it should be possible to very profitably apply the data on Recent stromatolites to the solution of many of the remaining problems in the study of stromatolites.


Chapter 10.1

USE OF STROMATOLITES FOR INTRABASINAL CORRELATION: EXAMPLE FROM THE LATE PROTEROZOIC OF THE NORTHWESTERN MARGIN OF THE TAOUDENNI BASIN

J. Bertmnd-Sarfati and R . Trompette

INTRODUCTION

Stromatolites have been used for intrabasinal correlation for a long time (Rezak, 1957). The purpose of this chapter is to give a good example of such a method of correlation in the western part of the Taoudenni Basin (Mauritania, Western Africa).

The Taoudenni sedimentary basin which is filled with Late Proterozoic to Carboniferous sediments overlies unconformably a metamorphic and granitic basement of Archean or Early Proterozoic age.

The stromatolites are almost entirely confined to Supergroup 1 of Late Proterozoic age and are particularly abundant along the northwestern margin of the basin. In the Adrar region of Mauritania (Trompette, 1973), four unconformable groups have been distinguished of which three can be followed eastwards as far as Algeria. They are from base to top (Fig. 1): (1) Char or Douik Group, essentially a detrital deposit with very scarce

stromatolites ; (2) Atar or el Mrei‘ti Group which includes most of the stromatolite-

bearing carbonate beds; (3) Tifounke Group, limited to the western part of the basin, lithologically

similar to the Atar or el Mreiti Group and also containing stromatolites; (4) Assabet el Hassiane or Oued Cheikhia Group which is mainly detrital

and is devoid of stromatolites.

STROMATOLITES AS ELEMENTS FOR CORRELATION

Following the first descriptions by Menchikoff (1946) of Conophyton from Chenachane, stromatolites were used for correlation purposes by geologists of the S.E.R.M.I. (Society for Mining and Industrial Studies and Realizations) in 1956. They recognized two horizons: a “giant Conophyton

sw

ADRAR OF MAURITANIA

111111. RICHAT - TOURINE

I FORT GOURAUD CHEGGA

.,*- NE

BIR AMRANE TlLEMSl WID EL MZEREB DAR CHENACHANE HAIDA

sw El Hmk NE _c_

0 0 0 0 tillite. base of Supergroup 2

U, Erosional Unamforrnltv Periglacial

U, Erosional Unconformny

U, Erosional Unconformlty Glaclal

U1 malor Unconformlty ----- ________

STRUCTURAL SKETCH

Fig. 1. Stratigraphic section of the Upper Precambrian “Supergroup 1” along the northwestern edge of the Taouden *i Basin.

STROMATOLITES IN BASIN ANALYSIS 519

- Tillite , base of “supergroup 2”

@l Conophyton and Jacurophyton

Fig. 2. Distribution of stromatolitic Unit 1 along the northwestern edge of the Taoudenni Basin.

horizon” equivalent to our Unit 3 containing Tungussia cumata Bertrand- Sarfati whose circular section was mistaken for Conophyton, and a “Collenia horizon” which is our Unit 5 characterised by the presence of Jurusania bioherms. Despite errors in identification, the correlations proposed by the S.E.R.M.I. remain valid.

Recent work (Bertrand-Sarfati, 1972c), based on a detailed study of stromatolites outcropping along the northwestern margin of the Taoudenni Basin from Adrar in Mauritania to Algeria, allows one to distinguish seven marker units within the Atar and Tifounke Groups:

Unit 1 with Conophyton and associated columnar forms. The Conophyton group is characterised by its unbranched columns and C. ressoti by the microstructure of its laminations which are regularly banded films. In the region of Aguelt el Mabha and el Mreiti, Unit 1 is missing (Fig. 2). On either side, in Tourine-Fort Gouraud, el Mzereb, Chegga and Dhar Cheir sections, small poorly shaped Conophyton occur only along a single horizon. At the two extremities of the basin, however, Conophyton accompanied by Jacutophyton and Baicalia mauritanica Bertrand-Sarfati in Adrar (Mauritania) and by Baicalia anastornosa Bertrand-Sarfati in Chenachane and Tilemsi ould Haida, occur in three distinct horizons.

Unit 2 with cylindrical columns forming biostromes and having parallel branches and a microstructure of micritic mats: Gymnosolen aff. ramsayi Steinmann and G. hankii Bertrand-Sarfati. These two forms never occur together in the same section and are absent in the Adrar, Tourine and Tilemsi ould Haida sections.

520 J. BERTRAND-SARFATI AND R. TROMPE'ITE

MAURITANIA

TlLEMSl OULD HAlDA

0 200 km 25'- 10' I

I /" Undulose 1 biostrornes Seriria radians

Unit 4 Tungussia globulosa

Tariouferra hernispherica

Unit 3

Fig. 3. Distribution of stromatolitic Units 3 and 4 along the northwestern edge of the Taoudenni Basin.

0 c 200 1 km M A U R I T E :w DHAR TlLEMSl OULD.

CHEGGA HAlDA

l?' 2 5 O

\ ALGERIA 1 A Til l i te, base of "Supergroup 2 '

/ Planar laminations

ORT-GOURA

WRAR OF MAURITANIA

MABH

Unit

A \ Jurusania aff. nisvensis

Jurusania lissa ' 1 Jurusania derbalensis

Fig. 4. Distribution of stromatolitic Unit 5 along the northwestern edge of the Taoudenni Basin.


Unit 3 has branched forms with a peculiar microstructure of hemispherical tussocks, which form either bioherms or biostromes. Between Tilemsi ould Haida and el Mzereb, it consists of a lower horizon with Tungussia globulosa Bertrand-Sarfati and an upper horizon with Tungussia cumutu. In Adrar the equivalents of these two horizons contain Turioufetia hemisphericu Bertrand-Sarfati (lower horizon) and Serizia radians Bertrand-Sarfati (upper horizon), characterised by the same type of microstructure as the Tungussiu horizons in the east of the basin (Fig. 3).

Unit 4 with undulose and elongated biostromes some tens to hundreds of metres long, occurring between Tilemsi ould Haida and Aguelt el Mabha. Correlation of this unit is based only on the external aspect of the biostromes easily identifiable in the field and on 1 :50,000 aerial photographs.

Unit 5 with bioherms a metre or so in diameter composed in the centre of columns belonging to the group Jurusuniu: J. derbalensis Bertrand-Sarfati J. lissu Bertrand-Sarfati, J. aff. nisuensis Raaben, all characterised by the same type of microstructure. These columns are vertical and separate in the center of the bioherm, but become oblique and linked in the marginal parts of the bioherm which is capped by planar laminations. This unit, often divis- ible into several sub-units, can be followed from Adrar to Chegga (Fig. 4).

Unit 6 with Gymnosolen directus Bertrand-Sarfati and Nouutila frutectosa Bertrand-Sarfati is composed of forms whose only common feature is their microstructure. The unit is well represented in Adrar whereas in Richat and Tourine-Fort Gouraud its identification remains hypothetical. I t disappears to the northeast.

Unit 7 has forms displaying a granular microstructure with peloids; the columns of Tifounkeiu rarnificata Bertrand-Sarfati and T. globulosa are subordinate to forms with planar or pseudocolumnar laminations. This unit is known in Adrar and Tourine-Fort Gouraud as well as in the Richat area, several hundred kilometres towards the centre of the basin.

CRITERIA USED IN SELECTING MARKER HORIZONS

In the selection of marker horizons along thenorthwestern margin of the Taoudenni Basin, three criteria have been used:

(1) Bioherms or biostromes (Units 4 and 5): these constructions ranging from a metre to several hundred metres in size are defined by their columns rather than their shape which is rarely specific.

(2) Columns (Units 1 and 2): these are easily identifiable by their morphology, longitudinal section and type of branching. (3) Microstructure of laminations (Units 3,6 and 7); this is a fundamental

feature especially when it is simple and therefore easy to define. The first two criteria have the advantage that they can be directly used in

the field. However, in all cases, we have tried to define each marker unit using the maximum number of criteria.

522

INTERPRETATIONS AND CONCLUSIONS

J. BERTRANDSARFATI AND R. TROMPETTE

(1) Stromatolites prove to be good elements for correlation on the scale of the margin of the Taoudenni Basin. The sequence of marker horizons is identical in all the sections with the exception of Chegga, where there is some doubt as to the relative position of Unit 1 and 2. This shows the need to use several horizons for correlation: an isolated stromatolite can only rarely be used.

The precision of the correlations based on stromatolites has been checked by classical methods of correlation (position of unconformities, marker horizons, etc.). For instance, the facies containing Vermiculus (worm-like structures 0.1-10.0 cm in size, probably of algal origin) always overlies Unit 3 or 4 (Fig. 3). (2) The reliability of the correlations depends on the good definition and

identification of the stromatolites, whose study must be pursued. The use of the microstructure of the laminations for correlation purposes is fundamental. Its value on the scale of the northwestern margin of the Taoudenni Basin is clear: the sequence of appearance of the microstructures is invariable.

(3) The sequence of stromatolite marker horizons leads to the recognition between el Mzereb and Aguelt el Mabha of a rise separating two sub-basins; across the rise several marker units are missing, probably as a result of non- deposition.

The western sub-basin is characterised by greater thicknesses and a larger number of stromatolite horizons. The eastern sub-basin is divided into two by a basement ridge which has influenced the distribution of stromatolites (e.g. Units 1 and 5 ) . (4) The basic problem of the meaning of the observed sequence of marker

units is not yet resolved. Two questions can still be raised: were there facies migrations during sedimentation or can ecological variations produce similar stromatolite sequences during different epochs? The successful correlations along the northwestern margin of the Taoudenni Basin, however, strongly suggest that the marker units are essentially synchronous.

( 5 ) Lastly, long-range correlations using stromatolites seem possible. The forms found along the northwestern margin of the Taoudenni Basin are very similar to those occurring in the Upper Riphean (950 k 50 to 680 k 20 m.y.) of the U.S.S.R. The comparison is supported by the geochronological results obtained by Clauer (1972) on rocks from Adrar in Mauritania (Fig. 1).

The excellent continuity and quality of outcrop together with the abundant variety of stromatolite forms found along the northwestern margin of the Taoudenni Basin, provide ideal conditions for testing the stratigraphical value of stromatolites. It is clear that stromatolites can be used for detailed intrabasinal correlations.


Chapter 10.2

PALEOECOLOGY OF CONOPHYTON AND ASSOCIATED STROMATOLITES IN THE PRECAMBRIAN DISMAL LAKES AND RAE GROUPS, CANADA

J.A. Donaldson

INTRODUCTION

Conophytons (columnar and pseudocolumnar stromatolites characterized by conical laminations) are abundant in a sequence of Helikian (Middle Proterozoic) dolostones that extends more than 300 km northeastward from Great Bear Lake to Coronation Gulf, Northwest Territories. When these conophytons were first reported (Donaldson, 1969), and in a subsequent interpretation of depositional environments (Donaldson and Jones, 1971), the stromatolitic dolostones were assigned to the Hornby Bay Group, following existing stratigraphic interpretations. Regional mapping has since revealed that the conophyton-bearing strata unconformably overlie the Hornby Bay rocks, and they therefore have been reassigned to the younger Dismal Lakes Group (Baragar and Donaldson, 1973). Conophytons also occur in the Hadrynian (Upper Proterozoic) Rae Group. The stromatolites of both groups are herein described, compared and interpreted, with emphasis on their utility for environmental interpretation.

Fig. 1 shows the regional distribution as well as the stratigraphic relationships of the Dismal Lakes and Rae Groups. The two groups form part of a sequence (in the Bear Province of the Canadian Shield) that unconformably overlies Aphebian (Lower Proterozoic) rocks of the Wopmay Orogen. (Stromatolites in strata of the Epworth Group, the main supracrustal component of the Wopmay Orogen, are discussed by P. Hoffman in Ch. 10.7). The Dismal Lakes and Rae Groups pass northwestward beneath Paleozoic cover, but possible correlatives occur to the north (Shaler Group of Victoria Island) and west (Brock Inlier). Eastward, the Rae Group appears to be discontinuously exposed for more than 200 km in a series of islands through Coronation Gulf. Dismal Lakes strata are unconformably truncated by the Rae Group; the Parry Bay Formation of Bathurst Inlet 300km east of Coppermine is probably correlative with dolostones of the Dismal Lakes Group.

5 24 J.A. DONALDSON

+ + + +

.... .... ........,. + + + + + + +

12

RAE G R O U P DISMAL, LAKES GROUP ml Dolostone,sondstone,siltstone Dolostone

COPPERMINE RIVER GROUP Sandstone;siltstone

HORNBY BAY GROUP Loo ',, Red beds

KI Bosalt Dolostone,shole

Fl ,.... Sandstone

Fig. 1. Geology of the region between Great Bear Lake and Coronation Gulf, Northwest Territories.

Radiometric studies have indicated that the Dismal Lakes Group has an age of about 1,200-1,250 m.y. and that the age of the Rae Group lies between 675 and 1,200 m.y. (Baragar and Donaldson, 1973). These estimates derive from K-Ar and Rb-Sr isochron measurements on igneous rocks that appear to provide reliable boundaries for the sedimentary strata (Muskox Complex beneath the Dismal Lakes Group, volcanic rocks of the Copper- mine River Group that overlies the Dismal Lakes Group and underlies the Rae Group, and Coronation Sills that intrude the Rae Group). These data, summarized in numerous reports of the Geological Survey of Canada, will not be reviewed here.

STRATIGRAPHY

The Dismal Lakes Group is unmetamorphosed, and primary structures are well preserved. Throughout most of the region the strata form a monoclinical


sequence that dips 5-10’ northward, although folds appear to be common in the area north of Great Bear Lake, not yet studied in detail. The lower part of the group, somewhat greater than 300m thick, consists of sandstones, pebbly sandstones and black shales. The upper part, about 800m thick, consists mainly of stromatolitic dolostones that have been subdivided into four units on the basis of lithology and bedding characteristics. In ascending order these units comprise: tan dolostone and red mudstone, a lower unit of laminated dolostone, a unit of massive-weathering dolostone and an upper unit of laminated dolostone. The uppermost strata are conformably overlain by basalts of the Coppermine River Group.

The Rae Group, at least 1300 m thick, forms a northward-tilted homocline in which dips less than 5’ prevail. Except for minor contact effects due to intrusion of the Coronation Sills, rocks of the Rae Group are unmetamorphosed. Sandstones, shales and mudstones compose about two-thirds of the group; dolostones and limestones, mainly forming a single prominent unit in the upper half of the group, constitute the stromatolite-bearing strata of interest in this report. Good exposures occur in sections beneath sills that cap islands in Coronation Gulf.

STROMATOLITE NOMENCLATURE

The procedure suggested some years ago by Cloud (1942) is adopted here as a sensible approach to stromatolite nomenclature. By this method, a distinctive stromatolite showing characteristics diagnostic of a type to which a Linnaean name has been assigned can be designated by the same name, but without capitalization or italics. Such an approach provides a basis for communication between “environmentalists” and “evolutionists” without com- promising their interpretations.

As used herein, the term “conophyton” refers to columnar and pseudocolumnar stromatolites characterized by conical laminations in which the apices point upward. Distinctively thickened axial zones are commonly but not invariably present, and although branching is extremely rare, it does occur in both Dismal Lakes and Rae conophytons.

Also present in the Dismal Lakes and Rae Groups are kussiellas, tungussias, gymnosolens, baicalias, jurusanias, inzerias and colonnellas and other varieties resembling named stromatolites of Raaben (1964b) and other workers (well summarized by Hofmann, 1969a). I t should be emphasized that the names have been applied mainly on the basis of field observations, because adequate reconstruction by serial sectioning is generally impractical for stromatolite columns more than 20cm in diameter, a size commonly exceeded in the Dismal Lakes and Rae occurrences. Nevertheless, for those stromatolites that have been systematically reconstructed, initial field identifications have been substantiated.

526 J.A. DONALDSON

Fig. 2. Typical conophytons of the Dismal Lakes Group, showing vertical attitude, lateral linkage, upward persistence in uniformity of column diameter and lack of intercolumn detritus. Length o’f hammer is 35 cm.


CONOPHYTONS OF THE DISMAL LAKES GROUP

Conophytons are the predominant stromatolites in the central massive unit of the Dismal Lakes Group. In contrast to the well-laminated units. above and below, this unit is characterized by massive-weathering beds, up to 3 m thick, of very pure dolomite. Diagenetic recrystallization has extensively obscured or destroyed primary textures and structures, but the conophytons are well preserved locally, especially where parting along laminations has exhumed accretionary surfaces (Fig. 2). The Dismal Lakes conophytons invariably are close packed, laterally, linked and have parallel axes that are perpendicular to bedding. Diameters from less than 10 cm to as much as 10 m have been recorded, but most fall in the range between 10 and 50cm, and diameters of individual conophytons tend to be relatively uniform. Zones of axial thickening (Fig. 3) are common but not ubiquitous, and many conophytons show only sporadic thickening along the central zone of stacked apices. Lamination textures show a great diversity, ranging from smooth and uniform to discontinuous and lumpy. These textures, resembling those regarded as characteristic of several named “forms”, appear to . represent various stages of diagenetic recrystallization of originally smooth and continuous laminae.

Synoptic morphology

For any conophyton-bearing stratigraphic unit, configuration of the depositional interface at any particular time (synoptic surface) can be established by considering the geometry of individual columns, size range of the columns and relationships between them. Fig. 4A illustrates two possible synoptic profiles for an occurrence of unlinked conophytons in which the columns are of different sizes. Because the columns are not laterally linked, determining whether apices of the columns maintained comparable elevations during growth, or whether only the conical surfaces rose above the intercolumn sediment, might here prove difficult or impossible (lamination stratigraphy, comparable to the matching of varves, could provide a basis for resolving this question). Fig. 4B shows the situation for conophytons of different sizes that are laterally linked. Here there can be no doubt that the apices of contemporaneous conophytons stood at various elevations during growth, and synoptic profiles can be readily defined. Such relationships are well displayed in numerous conophyton beds of the Dismal Lakes Group (Fig. 5).

Apical half-angles tend to be less than 20°, and thus the complementary inclinations of the conical accretionary surfaces are commonly greater than 70°, reflecting a relief somewhat greater than corresponding column diameters. For the largest conophytons, relief of depositional surfaces was clearly spectacular, locally exceeding 10 m.

J.A. DONALDSON

Fig. 3. Axial zone of a Dismal Lakes conophyton. Note that most laminae are smooth and of relatively uniform thickness. Porosity along axis (now infilled by diagenetic barite) appears to be partly of primary origin, perhaps a result of gas-bubble entrapment.

STROMATOLITES I N BASIN ANALYSIS 5 29

Fig. 4. A. Upper panel shows an oblique view of an outcrop surface, parallel to bedding, in which conophytons of different diameters occur as discrete columns. Dashed line represents the trace of a vertical section for which two equally valid interpretations of synoptic profiles are shown by thick lines in the lower panel (at a time when intercolumn detritus had accumulated to the level shown). B. Upper panel shows an oblique view of an outcrop surface, parallel to bedding, in which conophytons of different diameters are laterally linked. Dashed line indicates the trace of a vertical section for which only one synoptic profile can be inferred for any particular time of growth, as represented by the thick line in the lower panel.

Mode of growth

Whatever may be the control of accretionary-surface geometry for conophytons, the very nature of this geometry renders origin by sediment entrapment difficult to accept. Marked relief throughout a vast field of laterally linked cones, such as is indicated for the Dismal Lakes Group, would have a pronounced baffling effect on incoming sediment; widespread and uniform distribution of clastic sediment in such an environment would be extremely unlikely. Nor is there evidence (such as pronounced transverse elongation and significant inclination of the columns) to substantiate existence of the strong currents that would be required to sustain transportation of clastic detritus across extensive areas of sharply flexured substrate. Marked relief of the depositional interface should -also cause accumulation of abundant intercolumn detritus in an area of clastic influx, yet such detritus is extremely rare. Continuous successful trapping of all detritus on steep conical surfaces is difficult to accept, as is a sediment-trapping origin for the generally smooth and uncrenulated microlaminae of uniform thickness that are responsible for the conical surfaces. Origin by precipitation provides a better explanation for

530 J.A. DONALDSON

Fig. 5. Large conophytons of different diameters, central massive unit of the Dismal Lakes Group. Surface of outcrop is parallel to bedding; hammer is 35 cm long. Because the stromatolites are laterally linked, synoptic profiles analogous to that shown in Fig. 4B are inferred. Asymmetry apparent in foreground reflects growth on flank of the large conophyton in upper part of illustration.

the steeply inclined conical laminations. Purity of the carbonate and scarcity of terrigenous detritus support such origin for conophytons of the Dismal Lakes Group.

Associated sedimentary structures

On the basis of detailed measurement of five sections in the Dismal Lakes Group, distinct associations of primary sedimentary structures and stromatolite varieties have become apparent. Ripple marks, cross-bedding, intraformational conglomerates and oolites are particularly abundant in the lower and upper laminated units. Because these structures are generally indicative of moderate to strong current or wave action, they can be used as a reasonable basis for interpretation of environmental conditions during stromatolite growth. Similarly, evaporite casts and desiccation cracks are generally reliable indicators of at least some subaerial exposure, and thus their association with stromatolites has utility for environmental interpretations.

Fig. 6 summarizes the association of sedimentary structures with several


----

Fig. 6. Association of stromatolite morphologies (graphically depicted on the left) with common sedimentary structures. Structures depicted by symbols are, from left to right; ripple marks, cross-bedding, oolites, halite and gypsum casts, desiccation cracks and intraformational conglomerates. Inferred conditions of turbulence and depositional environments are indicated by lines (continuous in most likely environments, dashed to indicate probable range). Most of the oncolitic stromatolites range from 5 to 20 cm in diameter; for the other stromatolite varieties, there is a general increase in synoptic relief from a few centimetres (crenulate variety at top of diagram) to several metres (conophytons at bottom of diagram).

distinctive varieties of stromatolites in the Dismal Lakes Group, and presents qualitative inferences about depositional environments. Of particular interest here is the lack of sedimentary structures indicative of waves, strong currents or subaerial exposure in association with the Dismal Lakes conophytons, leading to the suggestion that these stromatolites formed in a subtidal environment.

Additional subtidal evidence

Inference of subtidal origin for the Dismal Lakes conophytons is supported by considering the effects of various environmental conditions on the stromatolites during growth. Recalling the previously discussed synoptic surfaces, it seems likely that even in the most hospitable intertidal environments, occasional storms would decapitate at least some cones, yet such events appear to be unrecorded in the Dismal Lakes conophytons. Consideration of the accretionary surfaces for laterally linked conophytons of different

532 J.A. DONALDSON

diameters necessitates synoptic surfaces in which correlative stromatolite apices stood at quite different elevations, a relationship which is more likely to obtain in a subtidal rather than an intertidal environment.

CONOPHYTONS OF THE RAE GROUP

Although conophytons are relatively scarce in the Rae Group, they are particularly interesting because of several unusual characteristics. Unlike the Dismal Lakes conophytons, those of the Rae Group typically occur as discrete columns 3-10 cm in diameter, and stromatolite-derived fragments locally occur between the columns. Furthermore, the conophytons are closely associated with other -stromatolites, and gradations between domal and conical laminations are common (Fig. 7). Nevertheless, most segments of columns showing conical laminations have well-developed axial zones, an attribute of all named conophyton forms. Another characteristic of the Rae conophytons is a distinct inclination of the columns (Fig. 7). Like the Dismal Lakes conophytons, they show some branching.

Environmental interpretation

In contrast to those of the Dismal Lakes Group, most conophytons of the Rae Group show characteristics suggestive of (or compatible with) strong currents and turbulent conditions. A shallow subtidal to intertidal environment is inferred to account for the column inclinations, lack of lateral linkage and the local occurrence of inter-column detritus.

CONCLUSIONS

Conophytons of the Dismal Lakes Group appear to have formed in an environment unlike the well-studied intertidal and supratidal environments in which analogues for other stromatolites in the rock record are presently forming. Nor do the closest modern analogues to conophytons, the small conical tufts forming around hot springs in Yellowstone National Park (Walter et al., Ch. 6.2) provide an adequate environmental analogue for beds of laterally linked conophytons that can be traced for hundreds of kilometres.

Paucity of associated sedimentary structures suggestive of strong currents, turbulence and subaerial exposure, together with an evaluation of the synoptic surfaces of the conophytons, has led to the conclusion that the Dismal Lakes conophytbns originated in a subtidal environment. Furthermore, the remarkably uniform and persistent laminations that record the conical geometry seem best explained by carbonate precipitation rather than by sedimen t entrapment .


Fig. 7. Conophytons of the Rae Group. Note that the most distinctive conical laminations occur in discrete and locally branched columns that are distinctly inclined. Gradation from hemispherical to conical lamination is well displayed in the vertical columns above the hammer, and some stromatolite-derived detritus occurs in the undulatory layers below the hammer, where some of the inclined columns are truncated.

534 J.A. DONALDSON

Conophytons in the Rae Group show features suggestive of strong currents and at least intermittent turbulence. Accordingly, a shallow subtidal to intertidal environment of deposition is inferred.

Most previous studies of conophytons have involved detailed investigations of morphology and microstructure, with little or no attention to other parameters. In this study the utility of a comprehensive sedimentological and stratigraphic approach to stromatolite investigations has been emphasized. Additional data on the association of stromatolite varieties and sedimentary structures for Precambrian strata elsewhere in the world could provide a sound basis for evaluating the question of whether certain stromatolites record an evolutionary sequence, or merely record responses to different environmental conditions.

ACKNOWLEDGEMENTS

Field work for this study was supported by the Geological Survey of Canada. Subsequent laboratory investigations have been supported by grants from the National Research Council of Canada. Brian G. Jones and A.H. Taylor collaborated in systematic studies of the stromatolites as well as in preparation of earlier reports. I am also indebted to J. Bertrand-Sarfati, R. Horodyski, H. Hofmann, M.E. Raaben and M.R. Walter for extremely helpful discussions.


Chapter 10.3

LACUSTRINE STROMATOLITES, EOCENE GREEN RIVER FORMATION, WYOMING

Ronald C . Surdam and John L . Wray

A widespread lake complex existed in western United States during Middle Eocene time. In southwestern Wyoming sediments now known as the Green River Formation were deposited in an interior basin occupied by ancient Lake Gosiute. The Eocene Green River Formation is perhaps the most famous accumulation of lacustrine sediments in the world, because it contains enormous deposits of oil shale and trona. Ever since the classic work of Bradley (1929a, 1931) it has been under almost continuous scrutiny by geologists. The most modern work is documented in Bradley (1963, 1964, 1973), Deardorff (1963), Culbertson (1966, 1971), Bradley and Eugster (1969), Deardorff and Mannion (1971), Surdam and Parker (1972), Eugster and Surdam (1973), Eugster and Hardie (1975), Surdam and Wolfbauer (1975), and others.

According to Bradley and Eugster (1969), Lake Gosiute at its highest stand covered an area of approximately 39,000 km2, and at its lowest stand about 3,900 km2 (Fig. 1). The size of the lake changed numerous times throughout the 4 m. y. it existed. During this time the lake was characterized by three major stages, each of which corresponds to a member of the Green River Formation. The major stratigraphic units from bottom to top are the Tipton Shale Member, Wilkins Peak Member, and Laney Shale Member (Fig. 2).

Recent studies of the Green River Formation suggest that Lake Gosiute was in fact a playa-lake complex, or in other words, the oil shale and associated sediments were the result 01 lacustrine deposition in a closed basin (Eugster and Surdam, 1973). By its nature a closed-basin, lacustrine complex is a dynamic feature; the area of the lake, depth of water, and salinity vary greatly according to seasonal inflow and evaporation (Langbein, 1961). Thus, the Green River Formation should contain abundant evidence of repeated fluctuations in lake level and shoreline position. I t is in this regard that algal stromatolites in the Green River Formation are significant in interpreting the geologic history of ancient Lake Gosiute.

Bradley (1929a) recognized and first described in detail “algal reefs”

536 R.C. SURDAM AND J.L. WRAY

0 2 5 50 0 40 80 - u miles kilometers

Fig. 1. Location of ancient Lake Gosiute in southwestern Wyoming and adjacent areas of Utah and Colorado and the outline of its hydrographic basin. Also shown are the minimum and maximum extent of the lake. (After Bradley and Eugster, 1969.)

WEST EAST

L A N E Y S H A L E

WASATCH FY.

W l L K l N S P E A K

Fig. 2. Schematic stratigraphic diagram for the Green River Formation of southwestern Wyoming. The Wilkins Peak Member contains the trona deposits.

STROMATOLITES IN BASIN ANALYSIS 5 37

I a 0 IA

a

1 a

W

z W W

W a

= I TIPTON SHALE

I= MEMBER ZONE #2

EXPLANATION

OIL SHALE

MARLSTONE

LIMESTONE

SHALE

MUOSTONE ALGAL LIMESTONE

TRONA

OR DOLOSTONE

b OSTRACODES

Q GASTROPODS - PELECYPOOS

0

El:, 8.. METERS FEET

Fig. 3. Schematic stratigraphy of the Tipton Shale Member of the Green River Formation. Zone No. 1 is 1.5-15 m thick; Zone 2 is 10-180 m thick; and Zone No. 3 is 0.9-6 m thick. Invertebrate remains are common constituents throughout much of Zone No. 1; however, with the exception of algae there was a dramatic decrease in biological activity by the end of this interval. Thus, the correlation between the first appearance of cryptalgal structures in the Tipton Shale Member at the top of Zone No. 1 and the disappearance of other fossil remains can be related to hypersaline conditions. (Surdam and Wolfbauer, 1975.)

Fig. 4. Stromatolites in Zone No. 3 of the Tipton Shale Member at Boars Tusk locality (Section 5 in Fig. 7).

(stromatolitic limestone units) in the Green River Formation. Two regionally significant stromatolite units occur in the Tipton Shale Member (Fig. 3) (Surdam and Wolfbauer, 1975) and nine to twelve such units are noted in the Laney Shale Member (Trudell et al., 1973), whereas stromatolitic


limestones are rarely present in the Wilkins Peak Member. Bradley (1929a) noted that these algal beds are important because they denote very shallow water and the proximity of shore. The stromatolite units are intimately associated with rocks containing mudcracks, flat-pebble conglomerates, saline crystal casts, ooliths and pisoliths, and coquinas containing pulmonate gastropods, all indicative of subaerial exposure or shallow-water deposition. Both the external morphology (Figs. 4, 5) and the internal fabrics (Fig. 6) of the algal stromatolites in the Green River Formation are similar to near- intertidal Recent columnar stromatolites and associated cryptalgal fabrics at Shark Bay, Western Australia, described by Logan et al. (1974).

Fig. 5. Artificial arrangement of weathered columnar stromatolites (Tipton Shale Member, Green River Formation).

The two stromatolitic algal horizons in the Tipton Shale Member, one at the top of Zone No. 1 and the other in Zone No. 3 (Fig. 3), are useful in a basin analysis of the Green River Formation sediments. These stromatolites are exposed continuously along the White Mountain scarp from Boars Tusk (T. 23 N., R. 105 W.) south to Rock Springs, Wyoming, a distance of approximately 55 km (Fig. 7). In the vicinity of Boars Tusk (Fig. 7, Section 5) the stromatolites of Zone No. 3 are up to l m thick and individual algal heads measure 30-40cm in diameter; however, some heads are so closely packed that the upper surface of the unit has a hummocky appearance (Bradley, 1929a; Bradley, 1926, fig. LXII). The amount of growth relief is difficult to ascertain, but appears to have been only a few centimeters. Southward along the White Mountain scarp, the algal heads of Zone No. 3 become smaller (15-30cm in diameter) and less domal in character. Approximately 23km south of Boars Tusk the stromatolites are more planar, consisting of stromatolitic algal structures 7-15 cm across (Fig. 7,

STROMATOLITES IN BASIN ANALYSIS 5 39

Fig. 6. Internal laminar morphology of columnar stromatolite (Tipton Shale Member, Green River Formation).

Section 3). These low-relief stromatolites resemble Recent forms occurring in lakes near the Coorong Lagoon, South Australia (Walter et al., 1973). Ten km farther south the unit has lost most of its distinctly stromatolitic algal character. Slightly further south again the unit is composed of very fine-grained dolostones characterized by mudcracks, flat-pebble conglomerates, salt crystal casts, and ripple marks (Fig. 7, Section 1).

Preserved tuffaceous ash beds in this part of the Green River Formation provide a means of interpreting depositional history within a chronologic framework. The relationship between the stromatolite-dolostone lithologic unit and the tuff beds is illustrated in Fig. 7. Because correlative tuff beds are time.stratigraphic units, it is readily apparent that the stromatolite unit is diachronous. The stromatolite unit is interpreted as representing a transition from a wave-swept shoreline along the northern margin of the basin at a high stand of the lake (maximum areal extent of lacustrine sediments) to a mudflat environment in the Rock Springs area at a relatively low stand of the lake (minimum areal extent of lacustrine deposits). During this contraction of the lake the shoreline migrated 55 km to the south and the time-stratigraphic position of the stromatolite unit climbed at the rate of about two stratigraphic feet per mile (ca. 33cm per km). Bradley (1964), using other geologic evidence, estimated that the topographic gradient in the Lake Gosiute basin during the deposition of the Green River Formation was one to two feet per mile (ca. 16-33cm per km).

A reasonable estimate of the amount of shoreline fluctuation can be determined by mapping the minimum and maximum areal extent of algal


SOUTH

LEGEND

I TUFF E3 OIL SHALE

Bfi DOLOSTONE

U SHALE LINE El MARLSTONE 0 SANDSTONE AND

LOCATfON E57 LIMESTONE Qc,, LIZ[\STRATIGRAPHIC

' TIME-STRATIGRAPHIC

SCALE

SILTSTONE vertical 0 10 0 3 H m H E E = I n

feet meters

horizontal 0 2 4 0 7 --

miles kilometers

T .22 N.. R. 105 W.

Fig. 7. Stratigraphy of uppermost Tipton Shale Member and lowermost Wilkins Peak Member along the White Mountain scarp north of Rock Springs, Wyoming. Note the exaggerated vertical scale. Circled inserts enclose sketches of stromatolite forms found in Zone No. 3 in Sections 2-5. Reticulate pattern in Section 1 indicates mudcracks and associated sedimentary structures. (Section 1 modified after Bradley, 1964; Sections 2-5 modified after Stuart, 1965; see Surdam and Wolfbauer, 1975.)

stromatolites in an individual unit (Fig. 8). The regional distribution of algal stromatolites in Zone No. 3 of the Tipton Shale Member shows that the lake shrank thousands of square kilometers during the transition from Tipton Shale to Wilkins Peak deposition. The areal distribution of the lowermost trona bed (No. 1) in the Wilkins Peak Member (Fig. 8), which lies just above the Tipton-Wilkins Peak contact, most probably marks a very low stand of the lake and thus helps to corroborate the extent of shoreline variation during the Tipton-Wilkins Peak transition.

Recognition and interpretation of the stromatolitic carbonates formed in ancient Lake Gosiute contribute to a better understanding of the sedimentology , stratigraphy, paleolimnology, and mineral deposits of the Green River Formation. By mapping the regional distribution and types of lateral changes characterizing individual stromatolite units, the fluctuations in shoreline position of Eocene Lake Gosiute can be quantified.


Fig. 8. Minimum and maximum areal extent of stromatolitic algal unit in Zone No. 3 of Tipton Shale Member. Stippled area represents carbonate mudflats exposed as the lake receded from a high to low stand. Note the areal extent of the lowermost trona bed in the Wilkins Peak Member of the Green' River Formation, which represents a maximum recession of lake level. The Rock Springs Uplift is shown as a dashed line and the margin of the Green River Basin appears as a solid line. (After Surdam and Wolfbauer, 1975.)

ACKNOWLEDGEMENTS

The authors gratefully acknowledge support by the National Science Foundation (GA-33789) and Marathon Oil Company.


Chapter 10.4

DEVONIAN STROMATOLITES FROM THE CANNING BASIN, WESTERN AUSTRALIA

P.E. Playford, A.E. Cockbain, E.C. Druce and J.L. Wray

INTRODUCTION

A diverse assemblage of stromatolites occurs in the Devonian reef complexes of the Canning Basin in the Kimberley district of Western Australia. These complexes are well exposed in a belt of rugged limestone ranges some 300 km long and up to 50 km wide along the northern margin of the basin. They occur in a structural subdivision named the Lennard Shelf (Fig. 1).

The complexes developed as reef-fringed limestone platforms around islands of Precambrian rocks and along the mainland shore of the Kimberley Block. They range in age from Middle to Late Devonian (Givetian to late Famennian).

Playford and Lowry (1966) recognized four main facies in the complexes: reef, back-reef, fore-reef, and inter-reef (Fig. 2). The reef facies forms a narrow, discontinuous, massive rim, which is commonly only 10-30m wide, around the margin of most platforms. I t was built largely by algae (both skeletal and non-skeletal) and stromatoporoids. The major part of each platform is made up of well-bedded back-reef deposits which were laid down in shallow subtidal to supratidal environments of the shelf lagoon. These deposits are mainly stromatoporoid and cryptalgal limestone, with coral limestone and oolite in some areas. The fore-reef deposits consist of talus derived from the platforms, together with variable contributions from terrigenous sources and indigenous organisms. They show steep depositional dips, commonly up to 35' in talus, and ranging up to near vertical where detritus has been bound by algae. The fore-reef deposits grade near the foot of each fore-reef slope into flat-lying inter-reef deposits, which are largely terrigenous, with some interbedded carbonates. The depth of water in the inter-reef basins ranged from a few tens of metres to perhaps 300 m or more.

Details of the stratigraphic nomenclature applied to the Devonian sequence are given by Playford and Lowry (1966). The limestone platforms consist of two units: the Pillara Limestone, which mainly embraces the extensive back-reef deposits, and the Windjana Limestone, which includes

544 P.E. PLAYFORD ET AL.

I + + + + + + + + 1 2 4 O + + + + + + + + A

&+ + + + + + 7 + + + + + + + _I I + + + + + + + -

I 5 AUSTRALIA

+ + + (PRECAMBRIAN) +, *

\ CANNING \b

BUGLE GA McWHAE F

0 100 k P h C l h l - T - -

+ + +

+ - +

I r n V I I .

I 124O

Fig. 1. Locality map.

STROMATOLITES SEA LEVEL 4

Fig. 2. Diagram illustrating facies relationships and distribution of stromatolites in Devon- ian reef complexes of the Canning Basin.

the narrow reef rim. Several formations are recognized among the fore-reef and inter-reef deposits; those of most importance to the present study are the Frasnian to Famennian Virgin Hills and Napier Formations (which in part are laterally equivalent), and the Frasnian Sadler Limestone.

Stromatolites occur in the reef, back-reef, fore-reef, and inter-reef facies of the complexes (Fig. 2), and they also cap drowned stromatoporoid-algal reefs (Fig. 3). Brief accounts of these occurrences have been given by Playford and Cockbain (1969) and Playford (1973).


SEA LEVEL _ .

Possibly > 100 rn I I

STROMATOPOROID-ALGAL PINNACLE REEF

Fig. 3. Diagram illustrating development of stromatolite bioherms over drowned stromatoporoid algal reefs.

SHALLOW-WATER STROMATOLITES

Shallow-water cryptalgal limestones make up much of the limestone platforms in the Canning Basin. They are believed to have been deposited in supratidal to very shallow subtidal environments.

The reef framework in many areas consists largely of cryptalgal limestone, generally having a fenestral fabric, associated with skeletal framebuilders (especially stromatoporoids, algae, and corals). Spaced columnar stromatolites make up part of this limestone, but much of it is massive and lacks recognizable growth forms. Stromatolite columns are indistinct where the inter-areas have been progressively filled with fenestral limestone similar to that forming the columns. However, columnar shapes are well preserved in some places, especially where the columns are covered by terrigenous detritus (Fig. 4A).

Fenestral columnar stromatolites similar to those in the reef facies also occur in the back-reef deposits. They are associated with well-bedded fenestral (birdseye) and oncolitic limestones. In some areas, especially in the Famennian sequence, fenestral limestone constitutes the major part of the bac k-ree f facies.


Fig. 4. A. Shallow-water columnar stromatolites from the Pillara Limestone (reef-flat environment) at Windjana Gorge. The stromatolite heads are covered by calcareous sandstone. B. Thin section of modem intertidal columnar stromatolite from Hamelin Pool, Western Australia, showing well-developed fenestral fabric.

STROMATOLITES IN BASIN ANALYSIS

DEEP-WATER STROMATOLITES

547

General description

Field evidence indicates that stromatolites in the fore-reef and inter-reef facies, and those capping drowned reefs, grew to considerable depths, probably exceeding 100 m below sea level. We refer to these forms as deep-water stromatolites.

A diverse stromatolite assemblage occurs in the fore-reef deposits, but well-developed occurrences are uncommon and are confined to condensed sequences in the Virgin Hills and'Napier Formations. In some areas stromatolites extend below the foot of the fore-reef slope into the flat-lying inter-reef deposits (Fig. 2).

Stromatolites similar to those in the fore-reef deposits cap ridges and pinnacles of drowned stromatoporoid-algal reefs (Fig. 3). Some of these form giant stromatolite bioherms, which are best developed on the north side of the Oscar Range (Figs. 6 and 7). Comparable, but smaller, stromatolite developments also occur on top of allochthonous blocks in the fore-reef deposits.

The fore-reef stromatolites grew on depositional slopes with original inclinations ranging from a few degrees to near vertical. Algae (and possibly bacteria) are believed to have been responsible for maintaining slopes at angles steeper than the angle of rest for loose debris. The algae include skeletal forms (especially Renalcis and Sphaerocodiurn), but in many cases non- skeletal forms seem to have been more important. In areas where deposition was very slow, the non-skeletal algae (and/or bacteria) formed recognizable stromatolite heads, but where deposition was more rapid their action in stabilizing steep layers of sediment has largely been inferred rather than proved.

Some of the deep-water stromatolites lack recognizable algal (or bacterial) filaments. These types are believed to have formed through the trapping and binding of biogenic and terrigenous sedimentary particles by non-skeletal algae and/or bacteria, with or without concomitant precipitation of calcium carbonate. Other stromatolites contain recognizable algae, belonging to the genera Sphaerocodium, Girvanella, Frutexites, and Renalcis.

It is not known whether the deep-water algae were photosynthetic. Some of the stromatolites could well have been formed by heterotrophic algae or bacteria. Many, but not all, of the columnar stromatolites grew vertically on depositional slopes, but it is not clear whether this was controlled by light (phototropism) or gravity (geotropism).

C. Thin section of Devonian shallow-water columnar stromatolites from the Windjana Limestone (reef facies) at Geikie Gorge. The stromatolite is growing on a stromatoporoid. The resemblance between the fenestral fabric of this Devonian stromatolite and that of the modern form from Hamelin Pool is striking.


TYPE PLAN SECTION TYPE PLAN SECTION

.. A R

I LONGITUDINAL B ~ R I DOMAL 1 Fig. 5. Diagram illustrating principal types of deep-water stromatolites.

Fig. 6. Map of part of the northern Oscar Range near Elimberrie Spring showing Elimberrie stromatolite bioherms nos. 1, 2, and 3.


B CROSS-SECTION A - B (natural scale)

, , ..- -

/ /

A

__- / -- __ - - -

W STROMATOLITE STRATIFICATION

FORE-REEF STRATIFICATION - JOINT (OF DEVONIAN AGE)

+ AXIS OF STROMATOLITE BIOHERM

_ _ _ _ 0 STROMATOLITE BIOHERM

0 FORE-REEF LIMESTONE

DIP & STRIKE OF STRATIFICATION I ~~ ~ ~~ ~ ~~

Fig. 7 . Map and cross-section of Elimberrie stromatolite bioherm no. 2.


Holdfasts of corals and crinoids encrust the surface of some stromatolites (Fig. l lA) , and ammonoids, nautiloids, and conodonts are conspicuous elements of the associated open-marine pelagic fauna.

The main types of deep-water stromatolites are illustrated on Figs. 5, 8, 9, and 10. Most of these forms grade from one to another, the named types being essentially end-members of series. The following descriptions briefly summarize some of the characteristic features of these stromatolites.

Spaced-columnar stromatolites are the most abundant forms. They commonly grew vertically on the fore-reef slopes. The maximum known height of such columns is l m , but their relief during growth was normally not more than about 5 cm, and they often show lateral linkages (Fig. 8A). Such linked forms can also be referred to as “pseudocolumnar”.

Contiguous columnar stromatolites (Fig. 9A) are less common. Most of these also grew nearly vertically on the depositional slopes.

Branching columnar stromatolites consist largely of the alga Sphaeroco- dium, and this genus also occurs in other columnar forms. A continuous range is found between columns composed entirely of Sphaerocodium filaments and those in which the filaments are rare.

Longitudinal stromatolites occur in some areas. In cross-section they resemble columnar types, but they actually form elongate ridges, directed down the fore-reef depositional slopes (Fig. 9B). This elongation was probably caused by down-slope currents.

Scalloped stromatolites are similar in vertical section to contiguous columns, but in plan they have a scalloped form, extending parallel to the strike direction of the original depositional slopes. The points of the scallops are directed up-slope. The down-slope face of each stromatolite is thickened and is composed largely of white micrite, whereas the rest is laminated and is commonly red. The micritic margin often contains abundant encrusting foraminifers.

Reticulate stromatolites are unusual forms, which developed low on the fore-reef slopes and extended into the flat-lying inter-reef deposits. They are characterized by reticulate patterns of meandering grooves between very low flat-topped mounds (Fig. 8B). The best lamination in these stromatolites is that marking successive grooves.

Undulous stromatolites show simple to intricate undulating to bulbous growth forms. Many have the appearance of being secondarily contorted, but as successive layers are commonly encrusted by crinoid and coral holdfasts

Fig. 8 . A. Laterally linked spaced-columnar deep-water stromatolites in the Virgin Hills Formation (fore-reef facies) at Ngumban Cliff, Bugle Gap. These stromatolites grew vertically (apparently phototropic or geotropic growth) on a depositional slope of about 12 . They grew extremely slowly and their relief during growth did not exceed 5cm. The inter-areas contain fragmental crinoid material. Note the horizontal geopetal filling in an ammonoid near the bottom right-hand corner of the photo. B. Reticulate deep-water stromatolites from the Virgin Hills Formation (inter-reef facies) at Bugle Gap.


(which require a hard surface for attachment) it is clear that they must have grown in their undulous form as hard, rigid bodies (Fig. 9C).

Domal stromatolites occur in the fore-reef deposits of some areas. They show some resemblance to the undulous forms, but have a more regular and well-defined mound shape. They tend to be elongated in the strike direction of the depositional slopes on which they grew. Bioclastic debris commonly accumulated on the up-slope side of each mound, and crinoid holdfasts are also concentrated there. Many of the domal stromatolites have grown over, or in association with, coral thickets composed mainly of the genus Aulopora. Most domal forms are about 0.5-2 m long, and these had a relief during growth of 0.2-1.5 m.

Depth relationships

The fore-reef deposits were commonly laid down with steep initial dips on the flanks of the limestone platforms. The depositional nature of these dips was first recognized by Guppy et al. (1958) on the basis of observed stratigraphic and structural relationships. This interpretation was confirmed by other workers, including Rattigan and Veevers (1961) and Playford and Lowry (1966), but they did not describe any method of separating the initial component of an observed dip from later components due to compaction or tectonism. More recent studies have shown that the amount of post-depositional tilt can be deduced, within limits of a few degrees, using geopetal fabrics that show the orientation of the rock in space at the time of deposition (Playford and Cockbain, 1972).

Many types of geopetal fabrics have been recognized in these rocks, and they can be broadly grouped into two types-cavity fillings and organic growth forms. The cavity fillings are “fossil spirit levels”, generally consisting of laminated sediment in the lower part and sparry calcite above, with the lamination and the spar-sediment contact approximating the original horizontal. They are best developed in fossils, especially closed brachiopods. The geopetal organic growth forms consist mainly of calcareous algae and stromatolites that grew nearly vertically, presumably under the influence of light or gravity. Fig. 9. A. Polished slab (parallel to the depositional dip) of contiguous columnar deep- water stromatolites from the Virgin Hills Formation at McWhae Ridge. They grew vertically on a depositional slope of about 35’. Note two crinoid holdfasts attached to the dark (iron-rich) lamina on the left-hand side of the slab. B. Polished section (parallel to the depositional strike) through a slab of longitudinal deep-water stromatolites from the Virgin Hills Formatioz at Bugle Gap. They grew on a depositional slope (away from the viewer) of about 20 and are elongated in the slope direction. C. Polished section through a slab of undulous deep-water stromatolites from the Virgin Hills Formation at Ngumban Cliff. Note the cluster of crinoid holdfasts overgrown by the stromatolite on the left. Other fossil material consists of crinoids, nautiloids, and ammonoids.


The sediment which has filled or partly filled cavities in fossils consists mainly of calcarenite (including bioclastic material) and calcilutite which is identical to sediment in the surrounding matrix. Lithification of the fore-reef limestones occurred very soon after deposition, as is shown by such features as penecontemporaneous neptunian dykes and slide breccias. Sediment must have entered the shells while the fore-reef sediments were being deposited, and lithification of both the geopetal sediment and the surrounding fore-reef strata occurred soon afterwards.

The spar-sediment contact and the layering in the sediment are normally parallel from one cavity to another, and it is clear that the sediment must usually have settled close to horizontal in each cavity. However, in some cases where sediment poured in on one side it maintained a significant depositional dip inside a cavity. It is therefore necessary to average numerous geopetal measurements when using cavity fillings to determine dip components.

The evidence of steep depositional dips in the fore-reef facies has important palaeoecological implications, as it points to a means of estimating relative water depths at the time of deposition. Where an observed dip is entirely depositional, the difference in original water depth between any two points on a bedding plane is the same as the present elevation difference between those points. Appropriate corrections need to be applied in order to estimate depths where there has been post-depositional tilting.

There are a number of localities where beds containing stromatolites in their original growth positions are exposed over considerable elevation ranges, allowing relative water-depth estimates to be made. Notable examples are at McWhae Ridge, Geikie Gorge, and Windjana Gorge.

McWhae Ridge is a faulted reef spine with associated off-reef deposits at the southern end of the Lawford Range platform. At this locality the Sadler Limestone and overlying Virgin Hills Formation were laid down with steep depositional dips against and over the reef spine. Geopetal fabrics are well developed in the Sadler Limestone, and detailed measurements of these have been made. The average of measurements on the geopetal fillings of 106 brachiopods, 39 Receptaculites, and 8 other fossils indicates that the observed dip of 41’ is made up of two components, a depositional dip of about 34’, and a post-depositional dip (due to tectonism and/or compaction) of about 7’. At the same locality the Sadler Limestone contains beds of Giruanella oncolites which developed on top of the reef spine after reef growth had ceased (apparently due to drowning) and were intermittently swept off, coming to rest on the steep depositional slopes on each side of the spine.

Fig. 10. A. Thin section of a deep-water stromatolite showing dark “microcolumns” of iron-rich Giruanella filaments. Quartz silt and fossil fragments are incorporated in the stromatolite. B. Thin section of a deep-water stromatolite showing dark iron-rich “shrubs” of Frutex- ites associated with quartz silt, fossil debris, and some Girvanella “microcolumns” in the lower part.


Fig. 11. A. Thin sectipn of a deep-water stromatolite showing abundant crinoid and coral holdfasts which encrusted successive stromatolite surfaces. The stromatolite contains abundant Frutexites. B. Girvanella oncolites and capped oncolites at McWhae Ridge. .The oncolites developed on a drowned reef spine and were periodically swept from the top (probably by wave and current action), coming to rest on a depositional slope of about 34O, in water at least


Conical Sphaerocodium caps grew vertically on some of the oncolites, at angles of about 35' to the depositional slope (Fig. 11B). The average direction of elongation of 95 caps is almost at right angles (within about 1%') to the original horizontal as indicated by the geopetal cavity fillings. The oncolite-bearing bed is exposed up the sides of the ridge to the top, and elevation data indicate that the lowest exposed Sphaerocodium caps must have grown in water at least 35 m deep.

The Virgin Hills Formation at McWhae Ridge was laid down over the Sadler Limestone with similarly high depositional dips. It is bright red in colour and contains a rich pelagic fauna of goniatites, nautiloids, and conodonts. Conspicuous stromatolite beds occur in the formation, interfingering with other limestones, and pinching out down the flanks of the ridge. This pinching out may have been the result of decreasing light penetration with increasing water depths, or it could have been a response to increasing rates of sedimentation on the flanks of the ridge. The difference in elevation between the lowest stromatolites and the eroded crest of the reef spine is nearly 55 m, and making allowance for post-depositional structure it is con- servatively estimated that the minimum depth of water in which those stromatolites grew was about 45 m. However, palaeotopographic cross- sections through the ridge suggest that the actual water depth was considerably more, probably exceeding 100 m.

The evidence at Windjana Gorge and elsewhere similarly indicates that stromatolites grew on fore-reef slopes and over drowned pinnacle reefs and allochthonous blocks in water up to 100 m or more in depth.

Age relationships

Druce (1976), using conodonts, has shown that well-developed deep-water stromatolites in the Virgin Hills Formation of the Bugle Gap area are of three different ages. The oldest horizon is late Frasnian, the middle horizon marks the base of the Famennian, and the youngest horizon is early Famen- nian. The youngest is of the same age as the stromatolite-bearing horizon in the Napier Formation at Narlarla (Napier Range).

The well-developed deep-water stromatolites are associated with strongly condensed sequences (i.e. very slow deposition). This is shown by the fact that in stromatolite-bearing beds the conodont zones are very closely spaced and conodont concentrations are high. A Devonian conodont zone represents an average of about 0.5 m.y., and on this basis the growth rate of some deep- water stromatolites may have been as low as 2pm per year. Growth could have taken several hundred thousand years for the taller individual columns, and a few million years for the large stromatolite bioherms.

35 m deep. Vertical (phototropic or geotropic) conical caps of Sphaerocodium grew on some of them, especially those in the final layer. Most oncolites have cores of crinoid columnals or brachiopods.


At McWhae Ridge stromatolite beds in the Virgin Hills Formation contain abundant palmatolepid conodonts (probably deep-water forms, see Druce, 1973) associated with relatively large numbers of icriodid (probably intermediate depth) conodonts having high degrees of morphological variability. Icriodids are lacking and palmatolepids are abundant in sediments above and below the stromatolite horizons. Druce (1976) suggests that this could mean that these stromatolites represent periods of relative shallowing in the prevailing deep-water environment.

It is important to note that in addition to the Canning Basin occurrences other Phanerozoic stromatolites are known to be associated with strongly condensed sequences. Examples have been reported from the Jurassic and Cretaceous of Europe by Szulczewski (1968), Radwansky and Szulczewski (1966), Marcinowski and Szulczewski (1972) and Jenkyns (1971). Modern oceanic ferromanganese stromatolites (Monty, 1973a) are also features of areas of very slow sedimentation. It therefore seems likely that there is a characteristic association of Phanerozoic deep-water stromatolites with condensed sequences.

Iron and manganese deposition

Much of the Virgin Hills and Napier Formations is coloured bright red by finely divided hematite. Minor amounts of black manganese oxide also occur. The reddest sediments are those containing the smallest amounts of platform- derived debris, laid down in relatively deep water, and exhibiting the slowest rates of sedimentation.

Red is also the dominant colour of deep-water stromatolites in the Canning Basin. In many cases the iron-oxide colouring is finely and uniformly dispersed, but in others there is a marked concentration of iron oxide, and smaller amounts of manganese oxide, in and around algal filaments. This is especially so with the genus Frutexites, which is preserved as minute “shrubs” of branching filaments containing high iron concentrations (Fig. 10B). Electron microprobe analysis (carried out by courtesy of C.S.I.R.O., Perth) indicates that the metallic iron content of these filaments is as high as 30%.

Concentrations of iron oxide also occur in and around Giruanella filaments (Fig. 10A) and (less commonly) in Sphaerocodium filaments. It is possible that some algae precipitated iron, but it seems more likely that iron bacteria living in close association with the algae were responsible. Such associations are common today (Choldony, 1922), and they no doubt also occurred in the Rast. The finely dispersed hematite in fore-reef and inter-reef sediments associated with the stromatolites is also likely to have a bacterial origin. It was probably precipitated originally in the form of iron hydroxide.

The genus Frutexites also occurs in Ordoviciari stromatolites of the U.S.S.R. (Maslov, 1960), and in Jurassic stromatolites of Poland


(Szulczewski, 1963). In each case it is characterized by the concentration of iron oxide in its filaments.

Red basinal deposits associated with Phanerozoic carbonate complexes also occur in other parts of the world, for example, in the Jurassic of Sicily (Jenkyns, 1971), the Triassic of the Alps (Fischer, 1964), the Carboniferous of Ireland (Schultz, 1966), the Devonian of the Camic Alps (Bandel, 1972), and the Silurian of the Michigan Basin (Mesolella et al., 1974). Among these occurrences, the red sediments of Sicily, Ireland, and the Michigan Basin are also associated with stromatolites. Maslov (1960) states that many stromatolites (of different ages) in the Soviet Union are red, and Walter (1972a) notes that some Precamljrian stromatolites in Australia occur in red sediments.

Iron and manganese are the only base metals known to have been concentrated in the Canning Basin stromatolites. However, the sole economic lead-zinc orebody found to date in the area, at Narlarla in the Napier Range, occurs above beds containing deep-water stromatolites. A genetic relationship seems possible, although there are no known lead-zinc orebodies near other deep-water stromatolite occurrences in the Devonian outcrop area.

Elim berrie stromatolite bioherms

On the north side of the Oscar Range in the vicinity of Elimberrie Spring there is a group of three giant stromatolite bioherms, referred to as Elimberrie bioherms 1 ,2 , and 3 (Figs. 6 and 7).

The largest bioherm, no. 1, is approximately 1 km long. It has developed over two stromatoporoid-algal pinnacle reefs. These reefs are thought to have been drowned as a result of abwpt subsidence to depths (possibly in excess of 100m) too great for continuing growth of the reef-building organisms. Successive stromatolite layers then developed over the drowned pinnacles (Fig. 3), the separate caps merging with continuing growth to form a single bioherm. Elimberrie bioherms 2 and 3 are also thought to have formed in this way, but the present level of erosion is not low enough to expose pinnacle reefs below these bioherms.

The best exposed bioherm is no. 2 (Figs. 7, 12). It is ca. 500 m long and has a similar shape, but on a greatly increased scale, to the domal stromatolites previously described. Its relief during growth is thought to have been at least 25 m, and possibly more than 100 m. Large convex outgrowths on the flanks of the bioherm meet at typical v-junctions of the same type as those found on the small domal mounds. Dips on the flanks are up to near vertical, and welldeveloped geopetal fabrics show that the dips are depositional.

Some stromatolite layers in the bioherms contain spaced-columnar or scalloped stromatolites, but others are massive. Skeletal algae, especially Sphaerocodium, occur in some beds. Receptaculites (possibly an alga) is abundant in parts of bioherms 1 and 3.

5 60 P.E. PLAYFORD ET AL.

Fig. 12. A. Aerial view of the southern crestal part of Elimberrie stromatolite bioherm no. 2, looking north. The width of the area covered by the photo in the centre is about 150 m, and the total length of the bioherm is about 500 m. Note the characteristic v-junctions of major convex outgrowths.


Conodonts indicate that the Elimberrie bioherms are of Famennian age, and that they represent condensed sequences. The core of bioherm no. 2 yields abundant icriodid (probably intermediate-depth) conodonts, whereas the flanks yield palmatolepids (probably deep-water forms) only.

Jenkyns (1971) described red stromatolites associated with ferromanganese nodules and crusts in red pelagic Jurassic limestones capping drowned reefal platforms in western Sicily. This occurrence appears to be analogous to that of the reef-capping stromatolites in the Canning Basin. Machielse (1972) has reported a thin discontinuous layer of stromatolites capping Upper Devonian pinnacle reefs in the subsurface at Rainbow Lake, Alberta. The stromatolites extend nearly 170m down the flanks of the reefs, and Machielse favoured the hypothesis that they developed during intermittent cessations of reef growth associated with lowered sea levels. However, in view of the Occurrence of deep-water stromatolites capping drowned pinnacle reefs in the Canning Basin it is suggested that the Rainbow reefs may have been drowned by rapid subsidence and then capped by deep-water stromatolites. This idea is also important in assessing deep-water versus sabkha hypotheses for the origin of the immediately overlykg Muskeg evapori tes .

COMPARISONS WITH MODERN STROMATOLITES

The best examples of columnar stromatolites known from modem seas are at Hamelin Pool, a barred hypersaline embayment forming part of Shark Bay in Western Australia. They were discovered by geologists of West Australian Petroleum Pty Ltd in 1954 and were first described by Logan (1961), who stated that the stromatolites are confined to the intertidal zone. By analogy he suggested that ancient stromatolites are also intertidal phenomena.

The Hamelin Pool intertidal model was accepted by many authorities. However, there were some early dissenters, notably Fischer (1965) and Monty (196513). Playford and Cockbain (1969) stated that Devonian stromatolites in the Canning Basin must have grown to depths of at least 45 m, and accordingly concluded that the Hamelin Pool model could not be applied to all ancient stromatolites. Achauer ixid Johnson (1969), Walter (1970a, 1972a), and Hoffman (1974) reached similar conclusions.

Monty (1971) suggested on theoretical grounds that, far from being necessarily restricted to the intertidd zone, there was no reason why stromatolites should not grow to great depths of the ocean, in complete darkness. Monty (1973a) has now shown that ferromanganese nodules and crusts

B. Ground view of the northern crestal part of Elimberrie stromatolite bioherm no. 2, looking north. The man standing in the centre left gives the scale. The structure is depositional, formed by successive accretion of stromatolite layers.


forming today in abyssal water depths and capping oceanic seamounts are bacterial stromatolites. Hofmann (1969a) had been the first to suggest that such ferromanganese crusts could be regarded as modem deep-water stromatolites, but he had not shown that they were of organic origin.

Recent work by the Geological Survey of Westem Australia at Hamelin Pool has shown that stromatolites there are not confined to the intertidal zone as previously supposed. Living subtidal stromatolites are widespread in Hamelin Pool, and they extend to water depths of at least 3.5m (Playford and Cockbain, Ch. 8.2).

The Hamelin Pool stromatolites (both intertidal and subtidal) show close similarities to shallow-water stromatolites in the Devonian of the Canning Basin (Fig. 7). A characteristic feature of most Hamelin Pool forms is that they have well-marked fenestral fabrics. Lamination is generally, but not always, crudely developed or absent, and the columns commonly have irregular margins. These features are similarly characteristic of the shallow- water Devonian stromatolites.

The Devonian deep-water stromatolites do not closely resemble the Hamelin Pool forms, but show some similarities to the ferromanganese oceanic stromatolites described by Monty. In common with oceanic stromatolites, the Devonian deep-water forms are generally finely laminated, lack fenestral fabrics, show evidence of iron and manganese precipitation, mark condensed sequences, and are associated with pelagic faunas. The reef- capping stromatolites in the Canning Basin are also analogous to stromatolites capping modem seamounts. However, the Canning Basin deep-water forms generally contain only small amounts of iron and minor manganese, whereas modem oceanic stromatolites are composed largely of iron and manganese oxides. Moreover, there is no evidence to show that any of the Devonian forms grew in abyssal water depths.


Stromatolites are important constituents of Devonian reef complexes in the Canning Basin. They grew through a wide range of environments, from shallow reef-fringed limestone platforms to deep inter-reef basins.

Cryptalgal limestone, ranging from massive in the reef facies to well bedded in the back-reef facies, makes up much of the platforms. Discrete columnar stromatolites occur in these deposits, which are thought to have formed in shallow subtidal to intertidal environments.

Deep-water stromatolites occur in the fore-reef and inter-reef facies and as cappings on drowned reefs. They grew to depths of at least 45 m, and probably more than l o o m , below sea level. They are best developed in strongly condensed sequences, and grew very slowly; their average growth rates probably amounted to no more than a few microns per year. The deep-water


TABLE I

Differences between deep-water and shallow-water stromatolites from Canning Basin

Feature Deep-water stromatolites Shallow-water stromatolites

Morphology fenestral fabrics absent usually finely laminated

diverse assemblage of forms

Occurrence

Associated fauna

Chemistry

Colour

characteristically in condensed sequences grew on depositional slopes, drowned reefs, and allochthonous blocks pelagic faunas common

some forms encrusted with crinoids and corals iron and some manganese precipitation important in certain forms

commonly red or reddish brown

fenestral fabrics common usually weakly laminated or unlamin ated columnar forms only, associated with oncolitic and fenestral limestones not in condensed sequences

grew on near-horizontal limestone platforms

associated with reefal and b io stromal organisms, especially stromatoporoids and Renalcis not encrusted by corals or crinoids

no significant iron or manganese precipitation commonly white or pale yellow

stromatolites characteristically grew on hard surfaces which stood above the surrounding sea floor, received little or no sediment, and were at depths too great for extensive colonization by skeletal reef builders. The largest examples are the Elimberrie stromatolite bioherms, which grew on drowned stromatoporoid-algal pinnacle reefs.

Distinguishing characteristics of deep-water and shallow-water stromatolites in the Canning Basin are summarized in Table I.

Shallow-water stromatolites from the Devonian limestone platforms show close similarities to modem intertidal and shallow subtidal stromatolites known from Hamelin Pool. The Devonian deep-water forms have certain characteristics in common with modem oceanic stromatolites; they do not closely resemble the Hamelin Pool forms.


Chapter 10.5

LOWER CAMBRIAN STROMATOLITES FROM OPEN AND SHELTERED INTERTIDAL ENVIRONMENTS, WIRREALPA, SOUTH AUSTRALIA

P. G. Hasle t t

INTRODUCTION

During his early investigations of the Lower Cambrian sequences in the northern Adelaide geosyncline, Mawson (1925) interpreted many of the structures found within the limestones as being of algal origin. Recent work of Preiss (197213, 1973b) and Haslett (1976) has confirmed the presence of algal stromatolites in most of the Lower Cambrian sequences. These sequences are largely composed of carbonate rocks which are considered to have formed in broad shallow seas (Daily, 1956; Dalgarno, 1964). Local zones of faulting and diapiric intrusion have, however, resulted in areas of considerable complexity within this broad and uniform depositional environment (Dalgamo, 1964; Dalgarno and Johnson, 1968; Daily and Forbes, 1969).

Detailed study of the Lower Cambrian stratigraphy and carbonate sedimentology at Wirrealpa (Fig. 1) (Haslett, 1976) has revealed numerous stromatolite beds associated with marked lithofacies changes in a sedimentary sequence adjacent to the northern edge of the Wirrealpa Diapir (Fig. 5). This sequence is laterally equivalent to the Wilkawillina Limestone, which forms the basal carbonate unit of the Cambrian over large areas of the central Flinders Ranges (Daily, 1956). The environmental model proposed herein for the stromatolite-bearing sequence at Old Wirrealpa is consistent with information collected over a far wider area of the central Flinders Ranges (Haslett, 1976).

At Old Wirrealpa Spring, 11 km NW of Wirrealpa, excellent outcrop and a lack of structural complexity allow variations in stromatolite morphology and distribution to be readily mapped. Selected samples were cut into three mutually perpendicular slabs, acidetched and stained with alizarin red S for laboratory study. Large thin sections of selected specimens were also examined. Preservation of features is generally good, but certain microscopic features have been obscured by diagenesis.

No attempt has been made to apply a rigorous classification to the stromatolites studied. Instead, the structures have been described in general

566

1 I 1 k ’ S O U T H A U S T R A L I A I I 1 I I

P.G. HASLE’IT

I 0 Scale 200

Precambrian . Fig. 1. Wirrealpa locality map.

terms, chiefly following the orderly scheme of Hofmann (1969a). Associated carbonate lithologies are referred to the classification of Dunham (1962). The term “gravelly” (see legend, Fig. 5) is used in the sense of Folk (1968) to indicate that lithoclasts within the grainstones are in the pebble to boulder size range.

LOCAL GEOLOGY

Wirrealpa Diapir

The Wirrealpa Diapir forms an area of complex, poorly exposed rocks, to the south and west of the stromatolite-bearing sequence at Old Wirrealpa. The mode of emplacement of this diapir is not well understood. It is mainly composed of sedimentary rocks, which may correlate lithologically with the lowermost part of the late Precambrian sequences elsewhere in northern


South Australia (B. Murrell, pers. comm., 1972). Evidence from detailed mapping in the vicinity of the diapir suggests that periodic intrusion of the Wirrealpa Diapir took place over a considerable span of Lower Cambrian time.

These episodic movements had marked effects on sedimentation. Uplifted areas acted as sources of carbonate and non-carbonate detritus which was incorporated into nearby Cambrian sediments. Large boulders up to 10 m in diameter, composed of characteristic diapiric lithologies, occur within the Cambrian sequence adjacent to the diapir, particularly on the western extremity of the stromatolite-bearing sequence. Although the southern contact of the diapir with the stromatolite-bearing sequence is poorly exposed, better exposed areas at Wirrealpa suggest that this contact represents an onlap of Cambrian sediments over an older eroded surface of diapiric material.

The large size of exotic boulders within the Cambrian, and the development of a Cambrian karst topography associated with areas of diapiric intrusion, attest to there having been considerable, and at times, subaerial, relief of these areas. Furthermore, the palaeogeographic distribution of local, Lower Cambrian lithofacies indicates that the areas of rising diapir acted as barriers to wave and tidal action during sedimentation (Haslett, 1976).

Local Cambrian sedimen tology

The geology of the Lower Cambrian stromatolite bearing sequence at Old Wirrealpa is summarized on the lithofacies map and diagrammatic cross- section (Fig. 5 ) . The sequence basically consists of massive cross-bedded ooid and lithoclast grainstones to the west, passing laterally into finely laminated calcareous mudstones to the east. A marked cyclicity in sedimentation, which is particularly evident in the western part, has probably occurred in response to local diapiric movements. Stromatolite beds occur over the whole lateral extent of the sequence. Evidence of other organic activity is limited to rare worm trails and bioturbation in the mudstone facies, and rare phosphatic skeletal remains as nuclei to ooids in the grainstone facies.

The eastern part of the sequence consists of dark-coloured calcareous mudstones, most showing fine lamination. Much of the lamination is believed to be of algal origin, because most beds display sporadic lateral development of low domes and nodular stromatolites. Desiccation mudcracks, although rare, are present throughout the sequence. Thin interbedded flat-pebble conglomerates are common, and are composed of locally derived cryptalgalaminite pebbles. These flat-pebble conglomerate beds are broadly lenticular, with sharp basal contacts and pebble-rich lower units. Uppermost boundaries are usually gradational into calcareous mudstones or stromatolites. Columnar stromatolites associated with the conglomerate beds are best developed where the conglomerate lenses are thickest. However,

568 P.G. HASLETT

columnar stromatolites are also found on the margins of narrow, steep-sided channels which cut through both older stromatoltes and underlying flat- pebble conglomerates (see cross-section, Fig. 5).

The desiccation features and flat-pebble conglomerates indicate shallow intertidal to lower supratidal environments of deposition (Lucia, 1972). Deposition of the flat-pebble conglomerates is believed to have taken place preferentially on broad depressions in a shallow tidal flat, the pebbles apparently having been derived from adjacent shallower areas, some of which were periodically exposed. With the exception of the conglomerate beds, the whole of the eastern sequence is believed to represent carbonate sedimentation under lowenergy conditions, on a sheltered, shallow intertidal mudflat.

Massive grainstone wedges, which thicken markedly toward the diapir edge, are the dominant lithological units in the western part of the sequence. Nearest the diapir the grainstones consist almost entirely of diapirderived debris, a large proportion of which is non-carbonate material. These grainstones contain poorly sorted, angular lithoclasts as well as blocks up to l O m in diameter. The upper parts of the large blocks are commonly encrusted by columnar stromatolites. Eastward there is a gradual lateral change in the nature of each grainstone unit, the size and amount of non- carbonate clasts decreasing, and the grainstones showing much better sorting. Light-coloured ooids, well-rounded quartz sand and well-rounded pebbles are the major clastic components of the grainstones. Pronounced medium-scale cross-stratification, commonly of the herringbone or chevron type occurs, indicating a strong tidal influence on sedimentation (Conybeare and Crook,

Fig. 2. Channel cutting pikt of a thick stromatolite bed within the grainstone sequence. Doming of the stromatolites may be seen at the right centre of 'the photograph. Note the rounded pebbles near the hammer in the channel bottom.

Fig. 3. Domed biostromal stromatolite bed within the grainstone sequence.


1968). Most upper bedding surfaces on the ooid grainstones show well- developed symmetrical ripple marks. Farther east, as the grainstone beds thin into the dominantly mudstone sequence, non-carbonate lithoclasts disappear and ooids become very dark in colour. Here the grainstones are poorly sorted and commonly merge at their easternmost extremities into packstones and wackestones.

Grainstone beds are usually overlain by thick units of columnar stromatolites. Clasts within these stromatolite beds show a similar decrease in grain size and non-carbonate component to the east, as described above for the underlying grainstones. Channels partly filled with well-rounded lithoclasts intersect the stromatolite beds to the west (see Fig. 2). This sequence of grainstones and stromatolites is considered to represent shallow, highenergy marine deposition, probably adjacent to an exposed headland of diapiric material.

Stromatolite beds in the west are typically overlain by a sequence of pale- yellow dolomitic mudstones, minor flat-pebble conglomerates, calcareous mudstones and thin irregularly bedded but well-sorted quartz sandstones, which locally fill welldeveloped desiccation cracks in the dolomitic mudstones. The shallow and restricted conditions suggested by this sequence may have been the result of shoaling due to local stromatolite growth under stable environmental conditions. Continued stability and lowenergy conditions are indicated by calcareous mudstone sedimentation for brief periods, even in the westernmost parts of the sequence, prior to renewed grainstone deposition.

STROMATOLITES

Thickness and macrostruct ure

The thickest units of columnar stromatolites occur within the western grainstone sequence, where stromatolitic beds are up to 10 m thick. Most, however, have a maximum thickness in the order of 1-1.5m and show gradual thinning to the east, where the underlying grainstone lenses thin into calcareous mudstones.

Stromatolite beds to the west generally have planar basal contacts with underlying grainstones. Stromatolite biostromes with relief of up to 25 cm (Fig. 3) occur adjacent to the diapir, and these pass gradually into tabular stromatolite beds to the east. There is a corresponding change within the bed from radiating to erect columns. In those cases where stromatolites develop on rippled ooid grainstone surfaces, stromatolite elongation is perpendicular to the ripple crest, indicating common current directions for both structures (see Fig. 4).

Beds of columnar stromatolites within the mudstone sequence are more lenticular and thinner than their counterparts to the west. Maximum

570 P.G. HASLETT

a.

b.

C .

CURRENT DIRECTIONS

Stromatolite bioherrn elongation in .open tidal flat. e n = 9

Measured from

Current ripples open tidal flat.

n = 17

Stromatolite bioherm elongation in restricted tidal flat.

Fig. 4. Current directions from open and sheltered intertidal areas.

thickness of the beds does not exceed 2m, and the relief of constituent domed stromatolites is less than 10cm. Columnar stromatolites are invariably associated with flat-pebble conglomerate beds, and the maximum thickness of stromatolites appears to occur where conglomerate lenses thicken. In many places both stromatolites and flat-pebble conglomerates are cut by channels. The columnar stromatolites may be seen to pass laterally into domeshaped stromatolites and algal laminites.

Current directions

A scarcity of bedding-plane exposures limits the number of measurements of stromatolite elongation that can be taken. A plot of the readings that have been obtained, however, shows good agreement within particular beds (Fig. 4). Current directions measured from stromatolites in the western part of the sequence, also show good agreement with those measured from ripple marks in associated grainstones. Stromatolites in the eastern part of the Fig. 5. Lower Cambrian lithofacies map and diagrammatic cross-section north of the Wirrealpa Diapir.


sequence do not show such marked elongation, and give a wider spread of current directions.

Some biohermal stromatolite beds in the eastern sequence show weak asymmetric growth, similar to that reported by Hoffman (1967). The asymmetry indicates a supply of carbonate mud from the northwest.

Lateral changes in stromatolites

In mapping the nature and distribution of stromatolites in the sequence, all beds were traced laterally and their characteristics noted. More detailed work was concentrated on two beds in the lower, more markedly cyclic part of the sequence. The two beds were selected for their good exposure and the apparent wide range of sedimentary environments which are laterally represented. Details of differences in stromatolite morphology from west to east can be most easily described with reference to representative samples A-K as marked on the diagrammatic cross-section (Fig. 5). Some very large allochthonous boulders derived from the diapir form sites for growth of columnar stromatolites (A). These boulders, which presumably remained immobile after initial deposition, have their topmost parts encrusted by stromatolites in layers up to 20cm thick. The stromatolites are slender, erect, cylindrical, close-spaced and about 5 mm in diameter. Rare branching is dendroid to anastomosed (Hofmann, 1969a). Laminae tend to be diffuse, slightly wavy and convex, with low relief. Wall structure is poorly developed. Intercolumnar sediment is up to 2mm in grain size, but most sediment within the columns is less than 0.3 mm. Stromatolites are not generally well developed within the coarse lithoclast grainstones between points A and B; however, minor crusts cap some boulders, and detached thumb-sized columns also occur.

At point B, stromatolite beds become well developed. The stromatolite columns tend to be radiate and constringed, with ragged margins. Anastomosing of columns is common, although furcate branching also occurs. The average diameter of individual columns is about 5 mm and most columns tend to be short. Thin-section examination reveals diffuse to lumpy laminations, with a moderate to low degree of inheritance. Some laminae are wavy and convex with a very slight tendency to lap over column margins. Microstylolitization has removed many column edges. Intercolumnar lithoclasts are up to 2cm in size, whereas material within the columns is usually less than 2 mm in grain size.

Stromatolites of the type which occur at B have a very notable feature about their microstructure. Within columns, generally at the same level from column to column, irregular patches of micrite and pseudospar occur. Rare cellular structures similar to Renalcis (Johnson, 1966, p. 25) may be found (Fig. 6). Other globular, branched areas of pseudospar are also apparent within the columns (Fig. 7). Branching of these areas is invariably divergent

576 P.G. HASLETT

Fig. 6. Preserved Renalcis-like algal remains from stromatolite in the grainstone sequence. (Thin section JW 10.) Fig. 7 . Globular patches of pseudospar within a stromatolite column in the grainstone sequence. Note the tendency of the patches to branch upward. (Thin section JW 10.)

upwards. Although preservation is fa r from perfect, it appears that algal species capable of carbonate precipitation have at times played a part in the development of these stromatolites.

Some distance farther east (C),lithoclasts seldom exceed 1 cm in diameter. Stromatolites are similar to those described above, being short, highly constringed columns with common anastomosing and dendroid branching. Laminations are extremely irregular and bulbous, with low inheritance.

Stromatolites at point D are associated with well-sorted grainstones which only rarely contain large lithoclasts. The erect cylindrical columns, some strongly constringed, may be up to 1.5cm in diameter. The columns are close-spaced, with anastornosed to umbellate branching. Wall structures are strongly developed and commonly show intense black pigmentation (Fig. 8). The diffuse to lumpy laminations show high relief. A basic difference appears to exist between the irregular but relatively continuous laminae in the central parts of the columns, and the clotted irregular fabric of the pigmented areas. The pigmented areas sometimes form laminae but are mainly concentrated along column walls (see Fig. 8).

As the stromatolite beds thin to the east, and associated grainstones lens out into mudstones and flat-pebble conglomerates, the stromatolite columns generally become broader in cross-section ( E ) . Branching of the erect to decumbent columns is usually umbellate. Laminations show diffuse to clotted textures with low to moderate inheritance. As seen in profile, the


Fig. 8 . Stromatolite with highly pigmented column margins, grainstone sequence. (Rock slab NW 30.) Fig. 9. Broad columnar stromatolite with associated flat-pebble conglomerates from the mudstone sequence. Arrows indicate irregular laminations which may be traced from one column to the next. (Thin section JW4.)

laminae are convex to penecinct with moderate to high relief. Large flat mudstone pebbles which appear within the stromatolite beds often cap columns and appear to have prevented continued growth. They subsequently became colonized and acted as a base for new columns. The columnar stromatolite beds are in general overlain by irregular stratiform stromatolites.

At point F, squat columnar stromatolites are closely associated with flat- pebble conglomerates within the mudstone sequence (Fig. 9). Columns are 2-5 cm in breadth, and are cylindrical to turbinate, with furcate branching. The erect columns show distinct, even laminations which are convex and have low relief. Partial linking of the close-spaced columns occurs locally. Several thick irregular layers (Fig. 9) contrast with the more normal thin, regular and strongly inherited laminations. The irregular layers are readily traceable from one column to the next. Stromatolites at G are dominantly stratiform, but show a strong similarity to the broad columnar forms with respect to their laminations. The porous, irregular layers are thick, and loaf-shaped. The high degree of inheritan& in overlying regular laminations means that many irregularities are reproduced with subsequent stromatolite growth. Some irregularities soon disappear, but some are enhanced by further growth, and furcate branching may result.

To the east, cryptalgalaminites and nodular stromatolites occur. Columnar stromatolites are associated with flat-pebble conglomerate beds, particularly

578 P.G. HASLETT

Fig. 10. Walled columnar stromatolite from the mudstone sequence, showing spar-filled cavities within columns. (Rock slab JW 30.)

in areas where the conglomerate beds are thickened. Some columnar stromatolites which have grown on flat-pebble conglomerates have fine carbonate mudstone between and above the columns (H). These stromatolites are upright and slender, with diffuse laminations but well- developed, highly pigmented wall structures (Fig. 10). Large irregular spar- and sediment-filled cavities are generally found within the columns. The cavities show radial shrinkage cracks and probably formed during periods of subaerial exposure. Another type of columnar stromatolite forms low, discrete bioherms (K ). Digitate branching of the close-spaced cylindrical columns is common. Laminae are convex, regular to wavy, and have low relief. Walls are not well developed. Intercolumnar sediment consists of carbonate mud, and of carbonate intraclasts up to 5 mm in length. Columns usually coalesce at the tops of the bioherms and pass into algal laminites.

Many columnar stromatolites associated with the thickest parts of flat- pebble conglomerate beds (I) have grown from oncolites which have finally come to rest (Fig. 11). Broad columnar stromatolites, upon umbellate or furcate branching, form numerous smaller constringed columns. The branching typically occurs at particular levels throughout the stromatolite. Laminae may be planar or low-relief convex, and are markedly down-curved on column margins. Most laminations are regular and somewhat diffuse, but thick irregular clotted layers with considerably higher relief and lower inheritance than the more normal laminations occur at certain intervals. These irregular layers are similar to those described above (F and G) and are traceable from one column to the next. Textures of the irregular layers are fairly well preserved, and are somewhat similar to those found in Madigunites mawsoni, and described by Walter (1972a) as a vermiform microstructure.

Stromatolites which grow on the edges of steep-sided channels which


Fig. 11. Branched columnar stromatolite from the mudstone sequence, showing irregular porous laminations within parts of the columns. (Thin section JW 19.) Fig. 12. Stromatolite from the margin of a tidal channel within the mudstone sequence. The effect of rare irregular laminae on the shape of later laminations with high inheritance may be seen. (Thin section JW 25.)

intersect flat-pebble conglomerate beds (J), have formed adjacent to, but at a topographically lower level than, those just described (I). They consist of cylindrical to turbinate close-spaced columns which show digitate branching and distinct convex to geniculate laminations (Fig. 12). The regular laminae have a moderate to high relief and generally show a moderate degree of inheritance. Lenticular bands with an irregular clotted structure also occur within these stromatolites. They tend to have higher relief than normal regular laminae and cause irregularities in the column structure. The irregular layers correspond in adjacent columns, and commonly occur at levels of column branching.

DISCUSSION

Workers on Recent stromatolites (Logan, 1961; Monty, 1965b, 1967; Gebelein, 1969) have recognised that the characteristics of a particular stromatolite are moulded by complex interrelationships between physical, chemical and biological aspects of the environment in which that stromatolite forms. Recent stromatolites have been shown to form by the binding, trapping or precipitation of sediment, mainly by communities of blue-green algae. It is only recently that the biological complexity and

580 P.G. HASLETT

environmental sensitivity of such algal communities has been really appreciated (Carr and Whitton, 1973). There still exists, however, a polazation af views on whether the major factors that determine stromatolite morphology are biological (Cloud and Semikhatov, 1969b; Walter, 1972a) or physico-chemical (Logan et al., 1964; Hofmann, 1969a). The contentious points, which bear on the validity of stromatolite biostratigraphy , have been summarized by several different workers (Hofmann, 1969a; Cloud and Semikhatov, 196913; Walter, 1972a), and will not be repeated here.

At Old Wirrealpa, stromatolite morphology appears to be related to lithofacies distribution. Most variation in form occurs laterally, with little or no vertical change within equivalent lithofacies. Precise correlation between stromatolite morphology and lateral changes in physical, chemical or biological factors is difficult to make in ancient sequences. At Old Wirrealpa no preserved organic matter has been found, chemical factors such as pH and salinity are indeterminable, and even physical factors are difficult to assess, due to outcrop limitations and diagenesis. Despite these difficulties, it is possible to make some tentative environmental analogies with modern stromatolites.

Mechanisms of growth

Algal stromatolites at Old Wirrealpa appear to have grown predominantly by the trapping and binding of sediment. This is most obvious in the western sequence where large amounts of non-carbonate sand and silt occur as layers within stromatolite columns. The ability of algae to trap and bind sediment in modem environments has been found to depend on factors such as current velocity, sediment size and supply rate, and the species of algae involved (Gebelein, 1969). Stromatolites do not occur in the bottom of channels at Old Wirrealpa, where energy levels and abrasive action must have been high. Columnar stromatolites which encrust large boulders in the western sequence appear to have formed under high-energy conditions but would, because of their elevation, have been less subject to abrasion.

Carbonate precipitation appears to have contributed minor amounts to stromatolite growth, and although evidence is scarce, much of the carbonate precipitation appears to have taken place under algal influence. Clotted, highly pigmented and cellular structures of possible organic origin have been observed within columnar stromatolites associated with clastic grainstones (see p. 575 and Figs. 6 and 7). Such bands of apparent precipitation may be traced from column to column at approximately equivalent levels. Although best devebped on column margins, some bands may be traced from within columns, down margins and as bridges across to adjacent columns. The thickening of these pigmented bands on the vertical column margins in contrast to the more normal laminations, which are thickest on column tops,


suggests that dominant growth was by precipitation. An almost total absence of noncarbonate detritus from within these bands, even in the western sequence, indicates that their development was associated with low current velocities and low sediment-supply rates. It may be that such layers are ancient analogues of calcified layers found in Recent algal mats, correlated by Monty (1967) with periods of emergence.

Similarly, the irregular sparry layers described from well-laminated columnar stromatolites within the mudstone sequence may also be the result of carbonate precipitation. Most of these once very porous bands occur on column tops (see Fig. ll), but some form continuous dome-shaped layers in stratiform and nodular algal limestones. The greatest thickness and frequency of these layers seems t o occur in columnar stromatolites from slightly more elevated parts of the tidal flat (compare samples I and J, Figs. 11 and 12). These layers may represent growth of different algal species, or a reaction of certain species to a period of subaerial exposure.

Effects o f sediment grain size

Despite the ill-sorted nature of sediment supplied to stromatolites in the western part of the sequence a t Old Wirrealpa, trapped and bound material within the columns is restricted to sand- and mud-sized grains. Detrital material caught between columns commonly contains larger well-rounded pebbles. Persistent agitation in this environment probably caused such pebbles to be dumped on growing algal mats at various times, terminating growth in these positions. The size of such pebbles, however, would have reduced their chances of being bound, and further agitation could have removed them and allowed renewed growth upon columns. Such a mechanism probably played an important role in the development of irregular laminations and to some extent may have influenced column branching in this part of the sequence (see p. 583).

Columnar stromatolites to the east are composed of very fine calcareous muds, with minor fine quartz sand and silt. Sediment between columns is usually of equivalent grain size, but may consist of intraclast chips, coarser quartz sand and occasional larger flat conglomerate pebbles which were trapped edgewise between columns. The periodic incursion of high-energy conditions into these sheltered environments, probably during storms, has resulted in local accumulation of flat-pebble conglomerates upon the stromatolite beds. Unlike their grain-size equivalents to the west, these pebbles appear to have been seldom shifted after initial deposition. Their flattened shape and the relatively rare incursion of high-energy conditions into this environment probably account for this. Depending on the thickness of the flat-pebble conglomerate beds, columnar stromatolite growth may have been effectively stopped.

The selective trapping and binding of finer sediment fractions by algae, as

582 P.G. HASLETT

illustrated by these Cambrian stromatolites, has been reported in studies of Recent environments (e.g., Black, 1933; Gebelein, 1969). The greater range of grain sizes within columns in the western part of the sequence at Old Wirrealpa, compared with that to the east, probably reflects to some degree the greater range available in the west, due to a poorly sorted supply. The possibility that this wider range of grain sizes might be contributed to by differences in the nature of algal communities living in the two environments cannot be discounted.

Stromatolite laminae

As a general rule, stromatolite laminae in the carbonate mudstone lithofacies show far greater regularity than those associated with grainstones in the western sequence. Columnar stromatolites within the mudstones have finer laminae of more uniform thickness and with a greater degree of inheritance than those to the west. Cryptalgalaminites and domed stromatolites have even finer and more uniform laminations than columnar varieties. As has been reported by Gebelein (1969) for some stromatolites of Bermuda, high degrees of lamina regularity may be related to areas of lower rates of sediment supply. As discussed above, other factors such as the degree of sorting of supplied sediment, and the nature of the actual algae, may also influence the regularity of the laminations. Stromatolites within the mudstone sequence commonly show down-folding of laminae at column margins to form walls. Such walls are usually best developed where intercolumnar sediment is fine grained. Broad, squat columnar types have better- developed walls than the slender columnar varieties.

The irregular nature of laminations within stromatolites in the western part of the sequence may be partly due to large pebbles temporarily preventing growth (see above, p. 581) but erosion may also play a part in this irregularity. The thickness of individual laminae varies markedly from one lamina to another, and within a single lamina there may be pronounced thickness variation. Wall structures such as described above are rare, the column margins usually being ragged. An exception occurs, however, in that some column peripheries have clotted and highly pigmented margins of what is interpreted to have been organically precipitated carbonate. These margins differ from the dominantly laminated central portion of the column, and are probably best described as mantles (Raaben, 1964: see Hofmann, 1969a, p. 18).

Distribution of columnar stromatolites

All stromatolites within the grainstone sequence at Old Wirrealpa are columnar. Columnar stromatolites in the mudstone sequence only occur in close association with flat-pebble conglomerate beds, which appear to have


accumulated in areas which were slight topographic depressions on the sheltered tidal flats.

The growth of columns and domes in present-day stromatolites is due, fundamentally, to enhanced stromatolite growth in certain parts of a mat and/or limited growth in adjacent parts. Substrate irregularities are believed to initiate column growth in certain instances (Logan, 1961; Walter, 1972a). Continued column development may take place due to the restriction of active mat growth to column tops and sides, and the retardation of growth between columns by such factors as excessive wetting, mechanical erosion, and very heavy sedimentation (Logan et al., 1964). Monty (1967) suggests that initiation and growth of columns may be related to the growth of certain distinct algal species. In areas with high rates of sediment supply Gebelein (1969) has reported nodular, rather than flat laminar, stromatolite growth.

Cambrian columnar stromatolites at Old Wirrealpa do not necessarily form on very irregular substrates, and some conglomerate beds are overlain by only slightly domed forms. The columnar stromatolites appear to have grown most prolifically in areas where sediment of suitable size for dgal trapping and binding was supplied at a high rate, mainly as a consequence of topographic position on the very shallow tidal flats. Judging by the greater frequency of desiccation mud cracks in those parts of the sequence in which flat-laminated and domed stromatolites are found, these stromatolites appear to have formed in even shallower areas, where sediment-supply rates might be expected to have been very low. Thus, sedimentation rate may well have been an important contributing factor in the distribution of columnar stromatolites within the Cambrian sequence.

Branching

Much of the above discussion of columnar-stromatolite distribution is also relevant to stromatolite branching. All of the columnar stromatolites show branching of one type or another. Stromatolites from within the western or grainstone sequence usually exhibit markedly divergent types of branching. This branching tends to occur a t irregular intervals during column development. Growth constraint (Hofmann, 1969a) caused by pebbles resting on active algal mats was probably the most common cause of initial branching of stromatolites in the western grainstone sequence. However, the low inheritance of laminae in these stromatolites means that there is a fairly small chance that any single growth constraint will give rise to branching, without the largely fortuitous intervention of other environmental factors. In most instances constraints in lamina growth are “evened out” by succeeding laminae.

In contrast, the branching of columnar stromatolites in the mudstone sequence is mostly parallel to slightly divergent. Columns within a particular

584 P.G. HASLETT

bed commonly branch a t about the same level, and their structure, to use the term of Walter (1972a), appears to “anticipate” branching. The regularity and high inheritance shown by column laminae means that any irregularities in the structure are reproduced for a considerable time by successive laminar accretions. These variations are frequently enhanced by other environmental factors, such as erosion or preferential deposition in hollows, and distinct branches result. Irregular bands of probable precipitative origin may provide the necessary irregularity within stromatolite columns to enable this sort of branching to begin (see Fig. 11).


Lower Cambrian algal stromatolites at Old Wirrealpa formed in open to protected intertidal environments. The nature, sorting and rate of supply of sediment to the growing stromatolites varied markedly from one environment to the next. Stromatolite growth took place predominantly by trapping and binding of the finer sediment fractions, although minor distinct intervals of carbonate precipitation also occurred. These intervals commonly show an irregular black pigmentation and contain poorly preserved cellular structures.

Columnar stromatolites formed where rates of supply of finer sediment were relatively high. Cryptalgalaminites and domed stromatolites developed in upper intertidal areas where sediment supply was lower.

Columnar stromatolites which grew on open tidal flats show irregular laminae and poorly developed wall structures. The slender, ,commonly constringed, columns branch at irregular intervals, and the branches tend to be divergent.

Columnar stromatolites from sheltered tidal flat environments are usually regularly laminated, with individual laminae showing high inheritance. Wall structures are commonly developed by lamina downfolding on column margins. The columns, most of which are closely spaced, have parallel branching which often occurs at distinct levels within the stromatolite.

ACKNOWLEDGEMENTS

The author wishes to acknowledge the expert assistance of the School of Surveying, South Australian Institute of Technology, in providing map control over the study area. Financial support for this contribution by the Centre for Precambrian Research, University of Adelaide, is also appreciated. Finally, the author thanks B. Daily, J.A. Donaldson and W.V. Preiss for their helpful criticisms of this paper.


Chapter 10.6

STROMATOLITES FROM THE MIDDLE PROTEROZOIC ALTYN LIMESTONE, BELT SUPERGROUP, GLACIER NATIONAL PARK, MONTANA

Robert J. Horodyski

INTRODUCTION

The Belt Supergroup, a thick Middle Proterozoic sedimentary sequence consisting of argillaceous, arenaceous, and calcareous rocks, extends from western Montana and northern Idaho into adjacent parts of Alberta, British Columbia, and Washington state (Ross, 1970; Harrison, 1972). The section exposed at Glacier National Park is near the eastern margin of the Beltian depositional basin; it is very well exposed, structurally simple, essentially unmetamorphosed, and contains an abundant and diverse assemblage of stromatolites.

The Altyn Limestone (actually a dolostone) is of restricted distribution, occurring only in a narrow outcrop band which immediately overlies the Lewis Overthrust on the east side of Glacier National Park and an adjacent part of Canada (Ross, 1959). The interval exposed in Glacier National Park attains a thickness of several hundred meters and is composed predominantly of silty and sandy dolostone. Although rocks of the Altyn Limestone have not been dated radiometrically, dates are available for other units of the Belt Supergroup and for the underlying metamorphic basement. These data indicate that the Altyn Limestone is approximately 1.3 ? 0.2 billion years in age (see Harrison, 1972).

Stromatolites have long been known from the Altyn Limestone in Glacier National Park, having been described previously by Fenton and Fenton (1931, 1937), Rezak (1957), and White (1970, 197313). The stromatolites are particularly well exposed on the southeast side of Appekunny Mountain immediately east of Appekunny Falls. The stromatolite beds range from several decimeters to about 2 m in thickness; locally, several of these beds are superimposed to form composite stromatolitic units several meters thick. Individual stromatolites are columnar, are commonly branched, and are composed of silty dolomite.

586

DEPOSITIONAL ENVIRONMENT

R.J. HORODYSKI

A 26 m thick interval, representative of the stromatolite-bearing portion of the Altyn Limestone at Appekunny Falls, was selected for detailed study. Two lithologic types predominate in this interval: (1) medium-bedded sandy dolarenites, forming meter-thick units which are overlain by (2) thinly interbedded silty dololutite and rippled sandy dolarenite, occurring as 1-2m thick units. In the second lithologic type, interbedding of rippled coarser- grained sediments with non-rippled finer-grained sediments resembles the “wavy bedding” of Reineck and Wunderlich (1968), suggesting a tidally influenced depositional environment, probably a tidal flat. The underlying medium-bedded dolarenite units have curved to slightly undulatory lower erosional surfaces, contain sparce centimeter-sized dolostone intraclasts near their base, and commonly display rather poorly developed, low-angle, planar cross-stratification. Although the precise depositional setting for these lower sands is not known, they evidently were deposited under somewhat higher- energy conditions and in deeper water than the overlying thinly interbedded dololutites and dolarenites. Deposition as intertidal sand bars seems unlikely because they lack the reactivation surfaces and type of cross-stratification characteristic of such deposits (see Klein, 1970). The curved erosional surfaces and rapid lateral thickness changes exhibited by some of these dolarenites suggest deposition within laterally migrating tidal channels cut to depths of about 0.5-2m. However, the geometry of other dolarenite beds in the unit suggests that deposition also might have occurred on an undulatory, rather than channeled, surface in the shallow subtidal zone.

Centimeter-sized dolostone intraclasts and medium sand- to gknule-sized quartz grains between stromatolite columns indicate that relatively high- energy conditions prevailed during stromatolite growth. As is shown in Fig. 1, the stromatolite beds occupy the same general position as the medium- bedded dolarenites: they typically overlie a sandy dolarenite bed and underlie thinly interbedded dololutite and dolarenite. This relationship indicates that they formed at water depths greater than that of the thinly interbedded dololutite and dolarenite; and, because shrinkage features in the overlying dolarenites are rare (and are not necessarily the result of subaerial desiccation), it seems likely that the stromatolites were generally, or perhaps continuously, submerged during growth. Irregular erosional surfaces typically are present at the base of the stromatolite beds. These surfaces have somewhat greater relief than those associated with the dolarenite beds (e.g., one erosional surface beneath a stromatolite bed exhibits a relief of 80cm throughout a lateral distance of several meters); some of the stromatolites may therefore have formed in tidal channels. The markedly streamlined shape of some Altyn stromatolites, evidently the result of current rather than wave activity, further suggests a tidal channel environment for at least some of the stromatolites.


SANDY DO L AR E N IT E - erosional surface

INTERBEDDED RIPPLED SANDY DOLARENITE AND NONRIPPLED S I LT Y DOLOLUT IT E

- gradational contact p?( #@, STROMATOLITE BED WITH

- erosional surface SANDY DOLARENITE

a DOLOSTONE INTRACLASTS AT BASE

- erosional surface INTERBEDDED RIPPLED SANDY DOLARENITE AND 'NONRIPPLED SILTY DOLOLUT ITE

Fig. 1. Diagrammatic representation of a stromatolite bed and adjacent strata in the Altyn Limestone.

LAMINAR STRUCTURE

Stromatolites in the Altyn Limestone are composed almost entirely of two types of laminae (Fig. 2A); compositionally, they are thus much simpler than stromatolites which occur in other formations of the Belt Supergroup (e.g., eight types of laminae have been recognized in stromatolites from the lower Missoula Group in Glacier National Park: Horodyski, 1975). The two main types of laminae in the Altyn stromatolites are compositionally similar, both being composed primarily of dolomite pseudospar. They differ principally in texture, one being finely (15-60 pm) crystalline, and the other very finely (5 - 15 pm) crystalline.

Laminae of the coarser variety are typically smooth and continuous, range from 0.04 to 5mm in thickness. and, as shown in Fig. 2B, they generally exhibit fining-upward size grading. Medium silt- to very fine sand-sized detrital quartz and feldspar compose 4-15% of these laminae, indicating that the accumulation of detrital particles played a role in their formation. In fact, the importance of detrital sedimentation was probably greater than suggested by the abundance of terrigenous detritus, because submillimeter- sized dolostone intraclasts are abundant constituents of the Altyn sediments. Due to their small size, such intraclasts are difficult to detect in the stromatolitic laminae; however, an estimate of their abundance may be made from the somewhat coarser-grained dolarenites which form a major part of

588 R.J. HORODYSKI

Fig. 2. Lamination of Altyn stromatolites. A. Transmitted light photograph of a thin section of a stromatolite showing very finely crystalline precipitated laminae (p = darker laminae) and finely crystalline, organically stabilized detrital laminae (s = lighter laminae). B. Enlargement of a portion of A showing precipitated laminae (p), organically stabilized detrital laminae (s), and mixed precipitated and organically stabilized detrital laminae

C. Photomicrograph of a thin section of a sandy dolarenile illustrating the predominance of sand-sized dolostone intraclasts (i) over terrigenous detritus (t). D. Transmitted-light photograph of a stromatolite thin section showing interlayering of organic films (darker layers) and calcareous laminae. The rosette-like crystals are pseudomorphs after barite.

( m ).


the Altyn Limestone. In the very fine and fine sand-sized dolarenites the ratio of detrital quartz and feldspar to dolostone intraclasts of approximate hydraulic size equivalence is about 1 : l O (see Fig. 2C). Thus, it appears that the coarser-grained laminae in the Altyn stromatolites formed chiefly by the accumulation of detrital particles (both autochthonous and allochthonous). In situ carbonate precipitation in laminae of this type was probably of secondary volumetric importance ; however, carbonate cementation might have been important in strengthening the growing stromatolite. That this lamina variety accumulated through organic stabilization (viz., by either trapping or agglutination) of sedimentary particles, rather than solely through the physical accumulation of sedimentary particles, is demonstrated by the laminae maintaining a fairly uniform thickness as they pass from the horizontal portions to the steeply inclined sides of stromatolitic growth surfaces.

Laminae of the finer-grained variety (Fig. 2B) range from 0.01 to 1 mm in thickness, are somewhat wavy and discontinuous, and commonly occur in couplets with the coarser lamina type. These finer laminae generally have sharp rather than gradational contacts with underlying comer laminae (see Fig. 2B), suggesting that the two lamina types might have been deposited in somewhat differing manners. This possibility is supported by the observation that in such couplets the thickness of the upper finer-grained laminae is not directly related to the thickness of the lower coarser-grained laminae (see Fig. 2A). The finer-grained laminae could have been produced either by trapping or agglutination of a very fine-grained carbonate detritus in an organic mat, or by in situ carbonate precipitation within such a mat. It should be possible to distinguish between these alternatives by comparing the abundance of terrigenous material in these laminae with that of texturally equivalent, physically deposited sediment occurring in associated strata. Specifi- cally, if these stromatolitic laminae resulted from the accumulation of very fine-grained particles, they should contain about the same proportion of terrigenous material as texturally equivalent, physically deposited sediment; on the other hand, if these laminae originated largely as a result of in situ precipitation, they should be relatively deficient in terrigenous material. Scanning electron microscopy of etched slabs and optical microscopy of petrographic thin sections revealed the finer stromatolitic laminae to be deficient in terrigenous material relative to texturally equivalent, physically deposited sediment. Thus, it seems evident that in situ precipitation was the major process resulting in formation of the finer-grained lamina variety.

Stromatolitic laminae texturally intermediate between these two distinct lamina types are also of common occurrence. Such intermediate types typically contain less terrigenous material than the coarser-grained laminae, and more than that occurring in the finer-grained laminae. This suggests that

E. Quartz and hyalophane pseudomorphs after barite, obtained by dissolving part of a calcareous stromatolite in hydrochloric acid.

590 R.J. HORODYSKI

they probably formed through a combination of both sediment stabilization and in situ carbonate precipitation.

Wavy 1-5pm thick organic films were detected in the upper portion of one stromatolite bed (Fig. 2D). These films probably represent sediment- poor organic mats, subsequently compressed during diagenesis. Associated with these organic films are pseudomorphs after barite (Figs. 2D, E), presently composed largely of quartz and a lesser amount of the barium-rich feldspar hyalophane. (Mineralogic composition was determined by X-ray diffrac- tion and an approximate chemical composition, Ko.sB%.z(Al, Si),OS, was determined by electron microprobe analysis.) The association of barite with stromatolites of both the Altyn Limestone and lower Missoula Group (Horodyski, 1975) of the Belt Supergroup, and its apparent absence in associated non-stromatolitic sediments, suggests that barium might have been .concentrated by organisms of the stromatolitic microbiotas and precipitated as barium sulfate during diagenesis.

ORIGIN OF THE CARBONATE

The origin of the carbonate occurring in both stromatolites and sediments of the Altyn Limestone is a subject of major importance in understanding the mode of formation of the Altyn stromatolites. As previously discussed, the finer-grained stromatolitic laminae appear to have been produced by in situ carbonate precipitation within an organic mat; however, these laminae account for only about 20% of the total volume of the stromatolites. The carbonate occurring in these stromatolites seems to have been derived mainly from the accumulation of submillimeter-sized carbonate intraclasts. Such intraclasts are also a major component of the associated dolarenites. The composition and texture of these intraclasts closely resembles that of the finer-grained (precipitated) stromatolitic laminae and, being relatively deficient in terrigenous material, differs from that of the finer-grained, physically deposited associated sediments; thus, it seems likely that the intraclasts in both the dolarenites and stromatolites of the Altyn Limestone were derived mainly by mechanical fragmentation of carbonate laminae, originally precipitated within algal mats.

STROMATOLITIC MACROSTRUCTURE

The Altyn stromatolites occur in discrete beds ranging from several decimeters to about 2 m in thickness. Some beds are distinctly pod-shaped and pinch out laterally; others extend continuously for several tens of meters (and possibly much further). Three varieties .of macrostructure occur; all stromatolites within a single bed exhibit similar macrostructural attributes.


Fig. 3. Two sets of stereophotographs of a single model of a somewhat divergently branched columnar stromatolite from the Altyn Limestone. This model was made by cutting serial sections through a stromatolite specimen, cutting cardboard templates of the column shape for each section, assembling these templates in their true positions, and filling open spaces with modelling compound. Representative laminae are outlined and part of a column was removed from the front of the model in the lower photographs. Bar for scale is 5 cm.

592 R.J. HORODYSKI


Two of these macrostructural varieties exhibit columnar branching; both types are composed of 4-20cm diameter, 0.1-2m high columns that are subcircular, oblong, or lobate in transverse section (Fig. 3). The two forms differ in branching pattern, one exhibiting markedly divergent branches (Figs. 4A, 5A), and the other being much less divergently branched (Fig. 5B). The columns of the former type developed under conditions of open- space packing on broad surface irregularities standing several decimeters above the surrounding depositional surface. The columns of the less divergently branched variety developed under more closely spaced conditions giving rise to stromatolite beds that extend laterally for tens of meters or more (Fig. 4B).

A third macrostructural variety also occurs. In contrast to the forms described above, the columns of this variety typically are unbranched, are elongate and smooth in transverse section, and commonly are. inclined. Column dimensions are similar to those of the two branched forms except that they are elongate in transverse sections, attaining lengths of several decimeters (or a few meters in some extremely elongate examples; Fig. 4C). Incli- nations from the vertical of 30-45" are common (Fig. 4D), and inclinations locally exceed 60' (Fig. 5C). Stromatolites of this third macrostructural type are very closely spaced (Fig. 5C), occurring in beds that extend laterally over distances of at least a few tens of meters (Fig. 4E). Although the columns within a 10 m long segment of a single bed typically are inclined in the same direction, stromatolitic columns in overlying beds commonly are inclined in a different direction (Fig. 5C).

In situ precipitated laminae and organically stabilized detrital laminae are present in these inclined elongate stromatolites; both types of laminae are asymmetric in thickness and shape. These laminae are similar in composition, grain-size of detrital constituents, and lamina thicknesses to those of the branched forms; they differ, however, in degree of uniformity and lateral continuity. The laminae of the inclined elongate stromatolites are smooth and continuous (Figs. 6A, B), whereas those of the branched columnar forms tend to be irregular and somewhat discontinuous (Fig. 6C).

Fig. 4. Altyn stromatolites in outcrop, near Appekunny Falls. A. Branched columnar stromatolites showing marked divergence of branches. These stromatolites occur as an isolated moundshaped structure in an intraformational conglomerate bed separated from underlying strata by an undulatory erosional surface (e). Columns and representative laminae are outlined. Scale in 5-cm divisions. B. A bed (5') composed of slightly divergently branched columnar stromatolites. C. Plan view of extremely elongate stromatolites. Scale in 5-cm divisions. D. Erect, branched, columnar stromatolites overlain by inclined, unbranched, highly elongated stromatolites (shown in plan view in C). Columns and representative laminae are outlined in parts of the outcrop. Scale in 5-cm divisions. E. A massive unit composed of eight superimposed stromatolite beds. Sketch of the stromatolites in part of this unit is shown in Fig. 5C.

594 R.J. HORODYSKI

I c L. ~ ~

A I l m I

J B r l m 1 C

Fig. 5. Shape of representative Altyn stromatolites. Markedly divergently branched, columnar stromatolites (A); less divergently branched, columnar stromatolites (B); and inclined, generally unbranched, elongate stromatolites (C). Sketches were made from photographs of outcrops in which columns and representative laminae were outlined with a felt-tipped pen. Stippled areas indicate portions of the outcrop where detailed structure is not clear.

FACTORS INFLUENCING STROMATOLITIC MACROSTRUCTURE

One of the most significant features of the Altyn stromatolites is that they clearly demonstrate the influence of physical conditions on the development of stromatolitic macrostructure. That the two branched columnar varieties of stromatolite were produced by the same (or very similar) microbiologic components is indicated by the close similarity of their microstructure and is consistent with their stratigraphic proximity, being separated by only 2.5 m of strata. Differences in the branching style in these two stromatolite types seems to reflect differing degrees of spacing during growth. The divergently branched forms grew as isolated mounds (Figs. 4A, 5A) elevated somewhat above the syrrounding substrate; a lack of lateral constraints on the developing stromatolitic columns permitted lateral expansion which resulted in their highly divergent branching pattern. Stromatolites with less divergently branched columns occur in beds that are laterally more extensive; this form developed under relatively closely spaced conditions (Fig. 4B, 5B). Lateral


Fig. 6. Laminar structure of Altyn stromatolites. A. Slabbed section of an unbranched elongate stromatolite. B. Thin section (transmitted-light photograph) of an unbranched elongate stromatolite. C. Thin section (transmitted-light photograph) oE a branched columnar stromatolite. D. Thin section (transmitted-light photograph) of part of an unbranched elongate stromatolite showing two erosional surfaces (e). Note the smooth, uniform lamination in the unbranched elongate forms (A, B, D) and the irregular, discontinuous lamination in the branched columnar form (C).

expansion of these stromatolites was inhibited by adjacent columns; the resultant stromatolites therefore developed only slightly divergent branching patterns. In addition, the shape of successive biohermal growth surfaces influenced the branching pattern of the constituent stromatolites. Develop-

596 R.J. HORODYSKI

ment of inclined branches seems to have been favored on the inclined sides of mound-shaped bioherms (Figs. 4A, 5A), whereas more erect branches developed on the generally horizontal growth surfaces of laterally extensive beds densely packed with stromatolites (Figs. 4B, 5B).

The.marked degree of elongation and streamlining exhibited by the third variety of Altyn stromatolite (Fig. 4C) was probably the result of unidirectional (or bipolar) currents flowing parallel to the direction of elongation (see Hoffman, 1967). However, other evidence of paleocurrent direction could not be obtained from any of these stromatolite beds; the stromatolites are so closely spaced that the beds are essentially devoid of physically deposited sediments and thus do not contain primary sedimentary structures of the type needed for paleocurrent analysis. Continuity of the stromatolitic columns and asymmetry of shape and thickness of the laminae indicate that the inclinations exhibited by these stromatolites are primary, rather than being the result of tilting or sliding of vertically growing stromatolites. Either increased biologic activity on one side of the stromatolites or an asymmetry in direction of sediment supply could account for these inclinations; however, the precise nature of the casual factors have not been identified.

Effects of currents are also evident in individual stromatolitic laminae. The smooth and uniform laminae of the elongate forms (Figs. 6A, B) are probably the result of rather vigorous currents; the irregular and discontinuous laminae exhibited by the branched columnar stromatolites (Figs. 6C) apparently reflect relatively less intense current conditions. Stronger currents during growth of the elongate stromatolites are further suggested by the common occurrence of erosional features in these forms (Fig. 6D).

SUMMARY

Stromatolites occurring in the Middle Proterozoic Altyn Limestone of Glacier National Park, Montana, are dolomitic, frequently branched, columnar structures that attain heights of several decimeters to about 2m. Strata of the stromatolite-bearing portion of the Altyn Limestone appear to have been deposited in a tidally influenced setting, probably the lower portions of a tidal flat and adjacent subtidal areas, with the stromatolites having formed in a shallow subtidal environment.

The Altyn stromatolites are composed largely of two types of laminae. One of these formed through the organic stabilization of submillimeter-sized dolostone intraclasts and detrital particles; the other appears to be the result of in situ carbonate precipitation, probably occurring within an algal mat.

Millimeter- and submillimeter-sized dolostone intraclasts are the major carbonate component of the Altyn sediments and stromatolites. The texture and composition of these intraclasts is similar to that occurring in the finer- grained, apparently precipitated stromatolitic laminae, and differs from that


of the finer-grained, physically deposited Altyn sediments. It seems likely, therefore, that most of the carbonate in both the Altyn stromatolites and sediments was derived ultimately from carbonate precipitated within algal mats.

Three macrostructural varieties of stromatolites have been identified; two of these are branched columnar forms and the third is an inclined form which is elongate in transverse section. Similarities in microstructure of these stromatolites, in conjunction with their stratigraphic proximity, suggest that the two branched forms, and probably also the elongate inclined forms, were produced by identical or very similar microbiotas.

The macrostructure of the Altyn stromatolites was strongly influenced by the physical environment. The closeness of spacing of columns and the shape of successive growth surfaces of the entire stromatolitic bioherm were of major importance in determining the branching pattern of the columnar forms. Stromatolites growing on isolated mounds elevated somewhat above the surrounding substrate developed in to divergently branched forms. The lack of lateral constraint on the developing columns permitted lateral expansion and the development of a highly divergent branching pattern; in addition, the inclined sides of growth surfaces of the entire stromatoliti'c bioherm favored the development of inclined branches. Stromatolites growing under more closely spaced conditions and forming relatively more horizontal biohermal growth surfaces developed into less divergently branched forms. Lateral expansion of these stromatolites was inhibited by adjacent columns and the growth of relatively more erect branches may have been favored by the more nearly horizontal nature of the bioherm growth surface. The inclined elongate forms apparently developed under relatively strong current conditions, their elongate nature probably being a streamlining effect caused by unidirectional (or bipolar) currents;

ACKNOWLEDGEMENTS

This study is part of a graduate research project undertaken at the University of California, Los Angeles. It was supervised by Prof. J.W. Schopf, to whom I am deeply indebted. I thank the National Park Service and the many rangers who provided valuable assistance. Thanks are also due to Prof. J.W. Schopf, Prof. J.A. Donaldson, and Dr. M.R. Walter for critically reading this manuscript. Financial support for this study was received from NASA Grant NGR 05-007-407 and NSF Grant GB-37257 (Systematics Biology Pro- gram), both awarded to Prof. Schopf,-and from Geological Society of America Penrose Grants.


Chapter 10.7

ENVIRONMENTAL DIVERSITY OF MIDDLE PRECAMBRIAN STROM ATOLITES

Paul Hoffman

INTRODUCTION

Stromatolites occur in a wide variety of facies -carbonate and non- carbonate, shallow and deep-water, marine and non-marine. Four formations are selected to document this variability. All are from the Middle Precam- brian of the northwesternmost part of the Canadian Shield.

(1) Stromatolite biostromes in the upper part of regressive cycles in the interior of a broad carbonate shelf (Rocknest Formation).

(2) Stromatolite bioherms between tidal channels along the outer margin of a carbonate shelf (Taltheilei Formation).

(3) Deep-water columnar stromatolites in terrigenous sediments of a foreshelf basin (McLean Formation).

(4) Non-marine stromatolites at the distal ends of alluvial fanglomerates (Murky Formation).

I t is concluded that the mere presence of stromatolites is not in itself sufficient to interpret environment of deposition, but that specific types of stromatolites may be environmentally diagnostic. Above all, stromatolites must be studied within their local lithologic and regional stratigraphic context, not in isolation.

The following publications may be consulted for more detail: regional geologic setting (Hoffman, 1973b); Taltheilei and McLean Formations (Hoffman, 1974a); Rocknest Formation (Hoffman, 1976).

REGIONAL GEOLOGIC SETTING AND GEOCHRONOLOGY

The major Middle Precambrian tectonic elements of the northwestern part of the Canadian Shield (Fig. 1) are: (1) the Slave craton; (2) the intracratonic Athapuscow and Bathurst aulacogens; (3) the pericratonic Wopmay orogen. The facies transitions from the craton into the foreland of the orogen are shown in Fig. 2, and from the craton into the Athapuscow aulacogen in Fig. 3.

600 P. HOFFMAN

Fig. 1. Major Early Proterozoic tectonic elements of the northwesternmost part of the Canadian Shield.

Ages of the four stromatolitic formations, based mainly on Rb-Sr isochrons of relevant intrusive rocks, are: (1) Rocknest Formation - between 1,865 and 2,200m.y.; (2) Taltheilei and McLean Formations -between 1,795 and 1,865 m.y.; (3) Murky Formation .-between 1,300 and 1,865 m.y. Thus, all but the Murky Formation are of pre-Riphean age.

The geochronology of the pre-Riphean formations is important because they contain certain columnar stromatolites with modes of branching typical of Middle and Upper Riphean stromatolites elsewhere (Raaben, 1969a). These pre-Riphean stromatolites have not been subjected to the rigorous serial sectioning required to elucidate their intricate detail, but the generalijies of their modes of branching are evident in the field (Fig. 4).

STROMATOLITES OF A CARBONATESHELF INTERIOR

The Rocknest formation constitutes the upper part of the miogeoclinal shelf in the foreland of the Wopmay orogen (Fig. 2). Sedimentary cycles are


Wesi

I

I

2

3 km

4

5

East TECTONliE FRONT

(BASEMENT UNKNOWN)

1 1 ; 1 slromatolilic shaly and cherty dolomite

-1 silty mudstone with turbidites - crossbedded orthoquartzite and quartz-pebblestone

pillow basalt and basalt breccia

arkose and granite-pebbleslone

Fig. 2. Facies transitions from the Slave craton into the foreland fold belt of the Wopmay orogen (see Fig. 1 for location).

well developed, except along the outer margin of the shelf and in the foreshelf basin to the west. Stromatolite biostromes within individual cycles can be traced for tens of kilometers across, and hundreds of kilometers along the shelf.

A typical cycle is shown in Fig. 6. There are nearly 200 such cycles, each

ATHAPUSCOW AULACOGEN SLAVE PLATFORM

GROUP

CHRISTIE BA'

RTHEl

K*HO'btLLL*

MSAN

UNlON 1SL

NOR13

u FlNE CROSSBEDDED OUARTZITE

0 RED CROSSBEDDED LITHWENITE

GREYWACKE TURBlDlTES ~

B R E D PLATY SILTSTONE

RED FISSIIE MUDSTONE

REDCONCRETONARI MUDSTONE

GREEN TO BLACK MUDSTONE

lTHS

SUBAERIAL BASALT

ASH- AND LAPILLI.FA*LL TUFF

m ARKOSIC DOLOMITE

STROMATOCITIC DOLOMITE

STROMATOLITK. LIMESTONE - DIGITATE FENESTRAL LIMESTWE

@ PLATY RHYTHMIC MIRLSTONE

DIGITATE MNUSTONE

RED FISSILE MARLSTONE

RED MUDSTONE WLTH SlROYATOLlTlC LIMESTONE OLISTOLITES

Q, 0 N

cd

Fig. 3. Facies transitions from the Slave craton into the Athapuscow aulacogen (see Fig. 1 for location).


Tungussiform

Kussieiiiform

Baicaiiform

Fig. 4. Modes of branching of columnar stromatolites (after Raaben, 1969), all of which occur in the Early Proterozoic.

2-20 m thick, in the formation as a whole. At the base of each cycle is a thin intraclast packstone bed, interpreted to be the condensed record of transgression. The remainder of the cycle is regressive, and results from progradation of intertidal and supratidal dolomite over sublittoral marlstone.

The upper part of each cycle is stromatolitic. In simple cycles, there is a vertical succession of stromatolites from laterally linked domes and columns of Omachtenia (Fig. 5a), 0.5 m in diameter, to flat-laminated and fenestrate sheets of Stratifera. The domes and columns tend to be elongate in plan and preferentially oriented, and are associated with edgewise flat pebbles and oncolites, and lenses of intraclast or ooid grainstone. The overlying stromatolite sheets have prone tabular oncolites and are more silicified. The vertical change in stromatolite morphology is interpreted to reflect decreasing wave action from lower to higher tidal flats.

The upper parts of more complex cycles have, in addition to the stromatolites described above,.units of black cherty dolomite up to several meters thick. These units contain a different stromatolite assemblage, in which the following types are prominent:

(1) Discrete smooth-walled columns of 2-5 cm diameter with a Gymno- soleniform (Raaben, 1969a) mode of branching (Fig. 5b).

(2) Tiny digitate forms resembling Alenia (Fig. 5c), in which vertical filament moulds are preserved where silicified. (3) Closely spaced vertical columns, less than 1 cm across, of Conophyton

(Fig. 5d), forming low domal bioherms 1-3 m across. (4) Flat-laminated sheets, some with incipient Alenia-like growths, others

with fenestrate or clotted (thrombolitic) fabrics. The black dolomite units have little evidence of turbulence and are interpreted as forming in stagnant ponds within or behind the main tidal flats.

Although lacking cycles, the uppermost part of the formation in the shelf interior also has extensive stromatolite biostromes, including a thick columnar unit with a Tungussiform (Raaben, 1969a) mode of branching (Fig. 5e).

604 P. HOFFMAN

Fig. 5. Stromatolites in the Rocknest Formation. a. Laterally linked domes. b . Columns with Gymnosoleniform mode of branching. c. Alenia-like form. d. Conophyton. e. Columns with Tungussiform mode of branching. The scale is divided in 3-cm intervals.


reg1 jive

dark cherty stromatolitic dolomite (Alenia -I ike , Gymnosolen)

light cherty stromatolitic dolomite IColonnella. Omechtanie, Stratifera)

~~

light intraclastic oolitic dolomite

dark argillaceous dolomite

t . transgressive I m a s t ic dolomite

Fig. 6. Typical cycle in the Rocknest Formation.

supratidal ponds I

intertidal sandflats

sublittoral nearshore I

beach lag I

In conclusion, the shelf interior facies is characterized by extensive stromatolite biostromes occurring mainly in the upper parts of regressive cycles .

STROMATOLITES OF A CARBONATE SHELF MARGIN

The Pethei Group (Fig. 7) undergoes a facies change from carbonate shelf over the Slave craton to foreshelf basin in the Athapuscow aulacogen. The outer margin of the shelf is sharply defined and well exposed in the Taltheilei Formation.

In the shelf interior, the Taltheilei Formation consists of extensive biostromes of alternating, columnar and flat-laminated stromatolites. In the thickest columnar biostromes, there is a basal zone of Conophyton in which the columns are tilted toward the shelf interior. Above, there is a thin zone of Baicaliform (Raaben, 1969a) branched columns, which passes into a thick zone of Ornachtenia (Fig. 8a), in which the columns are strongly elongate parallel to paleocurrents and normal to the shelf edge (Fig. 8b).

Toward the shelf margin, the columnar biostromes become thicker at the expense of the flat-laminated ones. Within 1-5 km of the outer edge of the shelf, the columnar biostromes reach 10-20 m in thickness and develop

606 P. HOFFMAN

S ATHAPUSCOW AULACOGEN SLAVE CRATON N

Fig. 7. Facies transitions in the Pethei Group from the Slave craton into the Athapuscow aulacogen.

depressions within them filled with cross-bedded intraclast grainstone. Reconstruction of these depressions in plan suggests that they were anastornosing channels oriented normal to the shelf edge. Cross-bedding indicates bimodal transport, suggesting tidal current action. As the channels become deeper, the intervening stromatolites appear in cross-section as large bioherms (Fig. 8c), the tops of which stand several meters above the channel floors. Both the bioherms and their constituent columns tend to be elongate normal to the shelf edge.

At the shelf edge itself, the channels widen at the expense of the bioherms. Less than a kilometer into the basin, the stromatolites disappear and the cross-bedded grainstones change to thinly flat-bedded, finer-grained, poorly sorted packstone. The lower slopes are made up of monotonous very thin- bedded lime mudstone with shale partings, and sheets of slump breccia composed of blocks of slope and shelf edge facies.

In conclusion, the outer margin of the shelf is marked by a belt of transverse tidal channels, between which are large columnar stromatolite bioherms. This contrasts with the extensive biostromes of the shelf interior.

STROMATOLITES OF A FORESHELF BASIN

The McLean Formation is in part the basinal stratigraphic equivalent of the Taltheilei Formation described above (Fig. 7). The lower foreshelf slope


Fig. 8. Stromatolites in the Pethei Group. a. Laterally linked columns of Omachtenia. The scaleis1.5 m long. b. Bedding surface showing elongate columns of Omachtenia. c. Margin of bioherm flanked by intraclastic channel fill in foreground; 3 m of section shown. d. Poorly laminated Conophyton-like columns in dark marlstone of the foreshelf basin floor. e. Calcareous columns in dark marlstone with thin inorganically deposited carbonate beds. Scale is divided into 3-cm intervals.

608 P. HOFFMAN

facies interfingers with stromatolitic marlstones of the basin floor. The marlstones also interfinger with thick tongues of greywacke turbidites, transported along the axis of the aulacogen parallel to the carbonate shelf edge. Absolute water depths in the basin cannot be determined, but the basin must have been a sufficient bathymetric trap to accommodate hundreds of meters of turbidites that nowhere lap out onto the adjacent shelf.

The basinal stromatolites are important in view of the general assumption that stromatolites are strictly shallow-water features. The stromatolites consist of poorly laminated, vertical to inclined columns (Fig. 8d), 3 cm in diameter. The columns must result from direct precipitation of carbonate, for the surrounding sediment is terrigenous and growth of the columns is snuffed out by thin beds of mechanically deposited carbonate (Fig. 8e). The columns are circular in plan and their inclination is parallel to the axis of the aulacogen. The laminations, although indistinct, are conical and there is, in places, an indistinct apical zone, suggesting that the columns are closely related to Conophyton.

The basinal stromatolites lap out of the aulacogen onto the adjacent shelf at the top of the Taltheilei Formation (Fig. 7). The passage from basin (McLean Formation) to shelf (Utsingi Formation) is gradational, the columns becoming more closely spaced and of slightly greater diameter, and the intercolumn sediment changing to dolomite relatively free of terrigenous admixture. That the shelf was submerged below wave base is suggested by the absence of a sharply defined shelf-edge facies change, and by the extreme monotony of the Utsingi Formation, which consists of hundreds of meters of poorly laminated Conophyton-like columns, unbroken by any other lithology. This reinforces the impression that Conophyton is a sublittoral form.

The vertical transition from the shallow Taltheilei shelf to the submerged Utsingi shelf is interesting. Here, there are cycles involving a reversible but ordered succession of the forms Conophyton, Jacutophyton, Baicalia and Omachtenia (in some cycles, the first or last named are absent). This succession is interpreted to reflect decreasing water depth (Fig. 9).

In conclusion, Conophyton is believed to be a sublittoral form in the Pethei Group. Poorly laminated Conophyton-like columns grew by direct carbonate precipitation in foreshelf basins of water, many tens, perhaps hundreds of meters deep.

STROMATOLITES IN A CONTINENTAL FAULT BASIN

The Murky Formation consists of alluvial fan deposits shed into localized continental fault basins during the final phase of tectonic activity in the Athapuscow aulacogen. It is made up of little deformed sheets of friable red


,,sea level

2 :onophyton? (poorly laminated)

.. . Omachtenia

lm

10 m

100m- water depth

Fig. 9. Postulated relation of stromatolite type to water depth.

Fig. 10. Morphology of stromatolites in the Murky Formation.

610 P. HOFFMAN

Fig. 11. Murky Formation. a. Small isolated colonies (scale in cm). b. Geologist with foot on the basal stromatolite sheet and hand on the bulbous crown. Crown in the foreground has fallen out of place.

conglomerate and breccia. Near the distal ends of the fanglomerate sheets axe units up to 10 m thick of red terrigenous silstone, interpreted as ephemeral lake deposits. It is in one of these silstone units that the calcareous stromatolites occur.

The unusual morphology of the stromatolites is shown Fig. 10. In


development, small isolated colonies (Fig. l l a ) coalesced to form an extensive basal sheet, from which grew discrete narrow pedestals with broad bulbous crowns (Fig. l l b ) standing nearly 2 m above the basal sheet. Internally, the stromatolite masses consist of closely spaced columns with a Kussielliform (Raaben, 1969a) mode of branching. The carbonate of the stromatolites must have been precipitated in place as the surrounding sediment is entirely terrigenous.

In conclusion, the occurrence of stromatolites in a non-marine non- calcareous basin adds to the diversity of environmental settings in which even the most ancient stromatolites have been found.


The mere presence of stromatolites is not sufficient to induce environment of deposition. They may be marine or non-marine, shallow or deep-water. However, certain types of stromatolites may be diagnostic of specific facies.

(1) Extensive stromatolite biostromes occurring in the upper parts of stacked regressive cycles are typical of the interior of a carbonate shelf.

(2) Large stromatolite bioherms separated by tidal channels are typical of the outer margin of a carbonate shelf.

(3) The distinctive columnar stromatolite Conophyton may be an exclusively subaqueous form, and poorly laminated Conophy ton-like columns may grow in foreshelf basins of water tens or even hundreds of meters deep.

(4) The elongation of stromatolites in plan is generally parallel to paleocurrent direction and commonly normal to the depositional strike.

(5) Stromatolite growth by direct precipitation, rather than sediment trapping, occurs in non-marine and deep-water marine basins where the particulate sediment is terrigenous.


Chapter 10.8

DISTRIBUTION OF STROMATOLITES IN RIPHEAN DEPOSITS OF THE UCHUR-MAYA REGION OF SIBERIA

S.N. Serebryakov

INTRODUCTION

The use of Riphean stromatolites in basin analysis is rather limited due to inadequate knowledge of their paleoecology. The widely employed extra- polation of the ecological features of the best known (intertidal) Recent stromatolites to Precambrian stromatolites meets well-grounded objections from many scientists (e.g., Maslov, 1959, 1960; Korolyuk and Sidorov, 1965; Trompette, 1969; Walter, 1970a, 1972a; Serebryakov, 1971, 1975; Bertrand-Sarfati, 1972b, c; Hofmann, 1973 ; Serebryakov and Semikhatov, 1973, 1974; Monty, 197313, c; and others). The most popular method of studying the paleoecology of the Riphean stromatolites, by analyzing microfacies of individual bioherms (e.g. Bertrand-Sarfati, 1970, 1972b), is not faultless either. In particular, the effect of biotic (associated with the systematic composition of algae) factors on the morphogenesis of stromatolites can hardly be detected in local material. This makes, in turn, the determination of the relative role of the environment more difficult. Furthermore, the interpretation of microfacies itself is based mainly on evidence from Recent stromatolites.

Some of these faults can be avoided by considering broadly the location of stromatolites in the Riphean deposits of an extensive region. This article is an attempt to determine a relative effect of biotic and environmental factors on the lateral and vertical distribution of stromatolites in the Riphean of the Uchur-Maya region.

REGIONAL GEOLOGY

The Uchur-Maya region of southeastern Siberia covers an extensive area (650 x 750 km) in the basins of the right tributaries of the Aldan river and includes three large structural elements (Figs. 1, 2): the eastern slope of the Aldan shield, the Uchur-Maya platform of the Siberian craton and the


Fig. 1. Schematic lithologic-facies profiles of the pre-Yudomian deposits of the Uchur- Maya region. Legend: 1 = sandstones (predominantly quartz) and conglomerates, grey-coloured; 2 = the same, red-coloured; 3 = siltstones; 4 = argillites; 5 = sandyklayey red-coloured rocks; 6 = variegated argillites of the Neruen suite; 7 = limestones;8 = dolomitic limestone and limey dolomites; 9 = dolomites; 10 = variegated clayey limestones, mark and calcareous argillites of the Totta suite; 11 = rhythmic interstratification of sandstones, sandy oncolitic and stromatolitic dolomites; 12 = the same and chemogenic dolomites; 13 = rhythmic interstratification of sandstones, siltstones, argillites and dolomites; 14 = variegated and dark-grey pelitomorphic platy limestones of the Malgina suite; 15 = light-grey and grey, massive and platy dolomites; 16 = grey dolomites with reddish-brown crust of weathering; 17 = dark-grey bituminous limestones of the Neruen suite; 18 = grey-coloured clastic and microphytolitic limestones and dolomites; 19 = red-coloured chemogenic, clastic and stromatolitic limestones and dolomites of the Ignikan suite; 20 = grey and dark-grey chemogenic, clayey-bituminous, clastic and microphytolitic limestones and dolomites of the Ignikan suite; 21 = pre-Riphean formations; 22 = surfaces of regional unconformity; 23 = surfaces of regional weathering; 24 = suite boundary lines; 25 = surface of pre- Yudomian erosion; 26 = K-Ar age of glauconite in m.y. (according to Kazakov and Knorre, 1970); 27-37 = forms of stromatolites (in brackets - morphologies of stromatolites found in the studied material analogous to the given form in microstructure but with a different morphology) : 27 = Omachtenia omachtensis Nuzhnov (Stratifera, nodular, rare Kussiella); 28 = Nucleella figurata Komar, Gongylina differenciata Komar, Kussiella kussiensis Krylov (Stratifera, Omachtenia); 29 = Baicalia baicalica (Maslov) (Stratifera), Colonnella kylachii Shapovalova, Suetliella tottuica Komar and Semikhatov (nodular, Stratifera), S . suetlica Shapovalova (Stratifera), Conophyton garganicum Korolyuk; 30 = Malginella malgica Komar and Semikhatov (Stratifera); 31 = Malginella zipandica Komar (Stratifera, Irregularia); 32 = Parmites aimicus (Nuzhnov) (Irregularia); 33 = Colonnella ulakia Komar (Baicalia, Tungussia); 34 = Minjaria sakharica Komar (Baicalia, Jacutophy- ton) , Conophyton metula Kirichenko (Jacutophyton), Conophyton lituum Maslov (Jacutophyton, Baicalia, Colonnella, nodular), Baicalia lacera Semikhatov (Jacutophyton, Conophyton), Baicalia ingilensis Nuzhnov (Jacutophyton); 35 = Baicalia maica Nuzhnov (Jacutophyton); 36 = Inzeria tjomusi Krylov (Jacutophyton); 37 = Inzeria confmgosa (Semikhatov) (Stratifera, Omachtenia?); 38 = area of deep burial of Riphean deposits; 39 = boundary of largest positive structures (Al. Sh. = Aldan shield, Om. Up. = Omnia uplift); 40 = zone of Nelkan fault; 41 = depressions in the Uchur-Maya platform (Uch. D. = Uchur, M.D. = Maya); 42 = Yudoma-Maya trough (Yud. M.T.); 43 = geographic position of lithologic-facies profiles (a-b and c-d, Fig. 1 ; e-f-g-h-i, Fig. 2 ; j - k , Fig. 6); 44 = position of sections shown in Fig. 5.

h i

N. SLOPE OF

1 h I

Fig. 2. Schematic lithologic-facies profiles of lower and upper subsuites of the Yudoma suite (Terminal Riphean). For position of profile line see Fig. 1 ; e , f , g , h , i - points of inflection of profile line. Legend: 1 = quartz and feldspar-quartz sandstones and gritstone with subordinate limestones and oncolitic dolomites; 2 = quartz and feldspar-quartz sandstones with subordinate gritstones, siltstones and argillites; 3 = complex, occasionally rhythmic alternation of predominating microphytolitic dolomites with stromatolitic, granular and sandy dolomites and sandstones; 4 = variegated cherty and clayey dolomites, argillites and cherty argillites with subordinate phytogenic and sandy dolomites; 5 = argillites, clayey and indistinctly granular dolomites with rare lenses of sandstones, stromatolitic and microphytolitic dolomites; 6 = indistinctly granular dolomites with intercalation of argillites, sandy and microphytolitic dolomites; 7 = bituminous and clayey dolomites with rare stromatolites, intercalations of argillites, siltstones and sandstones; 8 = bituminous limestones; 9 = distinctly granular dolomites with particular stromatolites and microphytolites; 10 = indistinctly granular limestones containing glauconite; I 1-1 8 = stromatolites with microstructures described from the following stromatolite forms: 11 = Boxonia grumulosa Komar; 12 = B. ingilica Komar and Semikhatov; 13 = Pankcollenia emergens Komar; 14 = Colleniella singularis Komar; 15 = Jurusania judomica Komar and cL

Semikhatov; 16 = Gongylina nodulosa Komar and Semikhatov; I7 = J. sibirica (Jakovlev); 18 = Linella simica Krylov; 19 = undefined stromatolites.


Yudoma-Maya trough. The latter is distinguished from the typical platform area by greater and more variable thicknesses, and, in some cases, by the facies of the Riphean deposits. It has been treated lately as an aulacogen. A number of lesser structures can be distinguished within the Uchur-Maya platform; the largest among them are the Uchur and Maya depressions separated by the Omnia uplift.

The composition of the deposits of the four subdivisions (phythems) of the Riphean is shown in Figs. 1 and 2. There are terrigenous, terrigenous- carbonate and carbonate units forming extensive sedimentary cycles (e.g., Nuzhnov and Yarmolyuk, 1959, 1963; Komar et al., 1970). Five cycles of this kind (which are regarded as groups) have been distinguished in the pre-Yudomian deposits (Fig. 1). At the bases of all the groups (except the Uii) there are unconformities; these are most distinct on the Uchur-Maya platform. The Yudoma suite (Fig. 2) corresponds to the Terminal Riphean (Yudomian) and overlies the older rocks with regional unconformity. It is intimately associated with the deposits of the Early Cambrian epoch, so that they form a single sedimentary cycle.

In the Uchur-Maya region the completeness of the sections and the thickness of the Riphean deposits gradually increase from west to east; the terrigenous suites are subject to the most significant lateral changes (Fig. 1). The variations of the carbonate suites are marked by changes in the proportions of limestones and sedimentary-diagenetic dolomites, and the amount and character of allochthonous material. Unfortunately, the marginal (nearshore) facies of the deposits are preserved only in the Omakhta suite and in the lower subsuites of the Yudoma suite.

The variations in the composition of the carbonate rocks can be partially accounted for by fluctuations in the salinity of the basinal waters (e.g. Akulshina et al., 1969). Supply of continental waters containing sodium and magnesium carbonate was evidently also of great importance; this facilitated accumulation of dolomite and limestone-dolomite facies (Serebryakov, 1968, Grigorev et al., 1969). Further from the coast (or as the transgressions developed), these facies were replaced by limestones, much like the phenomenon observed in the Neruen, Ignikan and Yudoma suites.

In the carbonate suites of all the zones of the Uchur-Maya region, one can commonly observe traces of local submarine erosion, cross-bedded and wavy-laminated structures, and ripple marks; clastic carbonate rocks (mainly endoclastites) are also widespread in occurrence. Desiccation cracks, typical of terrigencus suites, are found occasionally. This suggests deposition under generally shallow water and relatively high hydrodynamic energy for all of the Riphean sediments of the Uchur-Maya region.

This conclusion is supported by the practically ubiquitous occurrence of stromatolites and microphytolites. Stromatolites have been found in all the suites of the region, except the Ustkirba suite. The only stromatolites considered below are those of the carbonate and terrigenous-carbonate


IDESICCATIO)(I -- DEEPER WATER

623

/ m2 E5ii 1 7 0 3 7 n4 . . . ...

.... ..

Fig. 3. Sedimentary rhythms of the lower part of the Omakhta suite, Uchur River. Legend: 1 = stromatolites; 2 = oncolitic dolomites; 3 = sandy dolomites and dolomitic sandstones; 4 = finegrained silty sandstones; 5 = silty argillites; 6 = desiccation cracks; 7 = pseudomorphs after halite; 8 = ripple marks.

suites. (The stromatolites were studied jointly by the author, Komar and Semikhatov. Komar and Semikhatov made all the taxonomic determinations.)

The study of the Uchur-Maya stromatolites indicated that they formed subtidally. This conclusion is based on the uniformity of the stromatolite- bearing members of the suites over vast areas, persistence of microstructures within the same bioherm irrespective of its growth relief, and absence of any evidence that the bioherms were subjected to intense erosion. Neither is there any evidence of subaerial exposure of the stromatolites; in particular,

6 24 S.N. SEREBRYAKOV

no desiccation cracks have been found. According to Gebelein (1969), these are the only reliable indication of intertidal or supratidal formation of st ro ma toli tes .

Among the stromatolite-containing suites of the Uchur-Maya region, the evidence of shallowness is most pronounced in the Omakhta suite of the Uchur depression. There the suite is represented by rhythmically alternating terrigenous and carbonate rocks (Fig. 1). In the direction to the Aldan shield, the rhythms become gradually thinner, the chemogenic dolomites disappear, the rocks get more and more red-coloured, and numerous desiccation cracks and halite pseudomorphs are observed in the clastic members of the rhythms.

Rhythmicity in the Omakhta suite in the western, near-shore part of the Uchur depression is manifested by the repeated occurrence of the following rock types (Fig. 3): fine-grained silty sandstones with quartz and, then, dolomitic cement, then sandy dolomites, then oncolitic and stromatolitic dolomites with variable amounts of sand. Some of these rocks recur in the reverse order forming closed asymmetric rhythms (Serebryakov, 1971). In the sandstone beds which begin and conclude each rhythm, thin clay layers are common. Desiccation cracks, ripple marks and salt crystal moulds are confined to these layers, providing evidence of short periods of exposure. At the same levels there are some erosional surfaces. The carbonate members of the rhythms show no evidence of subaerial exposure. I t is apparent that the rhythms are of a transgressive-regressive character, stromatolites being confined to the periods of greatest submergence. It is significant that in the eastern zone of the Uchur depression, where there is less evidence of shallow deposition, the transgressive part of the rhythms is terminated by chemogenic dolomites overlying stromatolites. Thus, there is every reason to suppose that even the Omakhta stromatolites, which formed in the shallowest water, formed below the ebb tide level.

Thus, the observed distribution of stromatolites in the Uchur-Maya region should reflect their original distribution in the various subtidal environments. These environments differed at least in their position relative to the shore and in some cases in the conditions of sedimentation. Therefore, the first question to be answered is whether these differences affected the lateral distribution of the stromatolites. The second question is whether there exists any correlation between the vertical changes in the stromatolites and the variations in the composition and character of the enclosing rocks. These questions are considered below independently for the various morphological groups of stromatolites and for the stromatolites of a definite microstructure irrespective of their morphology.

The stratigraphic basis for analysing the distribution of the stromatolites is provided by the suites (formations) and subsuites which are litho-stratigraphic subdivisions; we are obliged to regard their boundaries as isochronous.


DISTRIBUTION OF ASSEMBLAGES OF MORPHOLOGICAL GROUPS OF STROMATOLITES

625

Study of stromatolites of the Uchur-Maya region (Nuzhnov, 1960, 1967; Nuzhnov and Shapovalova, 1965, 1968; Semikhatov and Komar, 1965; Voronov et al., 1966; Komar et al., 1970, 1973; Krylov and Shapovalova, 1970b; Komar, 1973) indicated that their groups, defined on a morphological basis, form laterally persistent stable assemblages confined to particular stratigraphic intervals. The spatial distribution of the assemblages was found to have three general features: (1) Stromatolite assemblages occur over practically the whole area of

outcrop of the carbonate and terrigenous-carbonate suites; they are not subject to any significant lateral changes. The variations in thickness and facies of the sediments are accompanied by changes in the relative abundance of some groups, the total abundance of the stromatolite and the thickness of their bioherms.

The Neruen suite provides a good example. Its thickness (Fig. 1) in the Yudoma-Maya trough is twice that in Uchur-Maya platform. At the same time there appear in the trough, among the carbonate rocks, beds of chemogenic, bituminous limestones and dolomites which result in a lesser abundance of the phytogenic and phytoclastic rocks. Extensive carbonate bodies appear in the trough in the lower and upper part of the suite represented on the platform mainly by argillites. In spite of these changes in the suite, Baicalia, Conophyton and Jacutophyton, normally forming “Jacutophyton cycles”, dominate everywhere (Shapovalova, 1965, 1968; Serebryakov et al., 1972). In the Yudoma-Maya trough, in comparison with the Maya depression, the individual stromatolite horizons get thicker and bioherms composed of any one group of stromatolites (predominantly Baicalia) are of greater size and more numerous. As a result, the relative abundance of the above groups is somewhat different. The horizons containing Baicalia account, in four sections of the Maya depression, for 58-60% of the total thickness of the stromatolite horizons and, in two sections of the Yudoma-Maya trough, for 67-7076. (2) The composition of the assemblages may remain the same, irrespective

of vertical changes in the composition and character of the enclosing deposits. Most spectacular in this respect are stromatolites of the Neruen suite. Conophyton, Jacutophyton, Baicalia and Colonnella are enclosed in its lower part in clastic and chemogenic dolomite varying in their fabrics and colour, and in its upper part, in less variable limestones and dolomitic limestones. These stromatolite constructions frequently cross the boundary of the lithologically different carbonate rocks without changing their morphology. Moreover, all these groups (except Jacutophyton) have been found enclosed in argillites (Fig. 4).

Absence of any direct interdependence between the morphology of


8

not exposed 740m

I

Fig. 4. Stromatolite horizon of the upper subsuite of the Neruen suite, Lyaki River. Legend: 1 = limestones; 2 = argillites;3-5 = stromatolite morphology;3 = Jacutophyton; 4 = Buicaliu; 5 = Colonnella; 6-9 = stromatolite microstructures typical for the following forms: 6 = Baicalia lacera Semikhatov; 7 = B. ingilensis Nuzhnov; 8 = Conophyton lituum Maslov; 9 = C. cylindricum Maslov (for jacutophytons the denominator indicates the microstructure of central column, the numerator that of the branches); 10-14 = colour of rocks: 10 = red; 11 = mottled (red and green); 1 2 = variegated; 13 = grey of various shades; 14 = dark grey to black.

stromatolites and the character of the enclosing rocks in the Neruen suite has been already considered in the description of the “Jacutophyton cycles” (Serebryakov et al., 1972). Yet the superposition in bioherms of different stromatolite morphologies can evidently be attributed to ecological factors which may have operated over extensive areas (see Ch. 6.4, fig. 7). This contradiction can obviously be accounted for by the fact that cyclic variations in the morphology of stromatolites and local variations in the character of rocks were controlled by different causes. For instance, the rate of sedimentation and intensity of subsidence are poorly manifested in the character of carbonate sediments; the role of these factors in the morphogenesis of stromatolites seems to be significant (Bertrand-Sarfati, 1972c; Serebryakov et al., 1972). The possibility that the environment may affect the shape of stromatolites but not the character of the sediment has been suggested by Maslov (1959) and Preiss (1973a).


(3) Different assemblages of stromatolites are found in lithologically similar beds of different ages; these assemblages may vary through a homogeneous section of carbonate rocks. Such a change in assemblages occurs in the uniform dolomites of the Tsipanda suite (see Fig. 1). In its lower part there are stratiform stromatolites (Mulginellu, Strutiferu, Irregulariu); in the middle part, stratiform and stratiform-columnar types (Irreguluria, Purrnites) are encountered, and in the upper part, these are replaced by columnar structures (Colonnellu, Buiculiu, Tungussiu, occasional Minjuriu, and forms resembling Tilernsinu). There is no lithological evidence that these changes in stromatolite assemblages are dependent on any changes in ecological conditions. Though such a dependence cannot be absolutely denied, the fact that the changes in morphological assemblages coincide with variations in the microstructures of the stromatolites indicates that these changes are probably due to changes in the communities of stromatolite-forming algae.

Dolomites similar in composition, colour and fabric to the Tsipanda ones occur in the Yudoma-Maya trough a t the base of the overlying Neruen suite. They are separated from the rocks of the Tsipanda suite by a regional weathering surface and a thin (5-20m) bed of argillites. Minjuriu and Tilemsinu (?) have not been found in these dolomites but Jucutophyton and Conophyton are abundant. This change in the morphological assemblage is accompanied by a partial change in the microstructures; but some of the stromatolites of the boundary horizons of the Tsipanda and Neruen suites have similar microstructures.

The Neruen stromatolites assemblage differs from the Tsipanda one not only in composition but also in the appearance of the cyclic structure of the bioherms (Jucutophyton cycles). Consequently, the change in the associations at the boundary of the suites under consideration was caused not only by the change in algal communities but also on the changes in the environment. It is significant that these changes are not manifested in the macroscopic and petrographic appearance of the rocks.

This is not an exceptional example of the repetition of uniform carbonate beds a t various stratigraphic levels. The dolomites of the Svetly and the Tsipanda suites are similar in their composition and position in the sedimentary cycles. Dolomites of the Svetly type have been encountered in the Neruen and Ignikan suites. Yet the stromatolite assemblages in these suites are different. The lithologically similar deposits are unlikely to have formed, in all cases, under different conditions. It is much more probable that we have to deal here with the age variability of stromatolite morphology; this view is reinforced by the fact that there are corresponding variations in their microstructures (see below).

The specific features of the distribution of the stromatolite assemblages allow two conclusions. Firstly, the vertical change in assemblages often has no direct relationship to changes in the enclosing deposits. Though the effect of environment on the stromatolites in the sections can be established in


some cases, the general sequence of assemblages can hardly be explained without admitting that it was dependent on the evolution of the stromatolite- forming algae. Secondly, there is very little relationship between the distribution of assemblages, the conditions of sedimentation and the distance from the coast. This does not exclude the effect of environmental factors on stromatolite morphology within the bioherms or even in certain areas of basins (see Ch. 6.4). The effect of these factors is not manifested in the characteristics of the rock, which makes their interpretation more difficult.

Those ecological factors which affected the morphology of the stromatolites hardly influenced their microstructure. In fact, the intrabioherm variability of stromatolites is normally unaccompanied by microstructural changes (Komar and Semikhatov, 1965, 1968b; Komar, 1966; Krylov, 1967a, 1972; Serebryakov, 1971; Walter, 1972a). The abundant evidence of the Recent and ancient stromatolites (see Ch. 7.1) suggests that particular microstructures owe their formation to particular communities or species of algae. If this is correct, the distribution of stromatolites with particular microstructures must to some degree reflect the original distribution of the stromatolite-forming communities.

DISTRIBUTION OF STROMATOLITES WITH PARTICULAR MICROSTRUCTURES

The specific features of microstructure usually provide the basis for distinguishing forms (“species”) of stromatolites; the concept of “group” is sometimes identified with the concept of “stromatolites of a certain microstructure”. Krylov (1967a, 1972) has repeatedly discredited such identifications, stating that whether a form does or does not belong to some group, according to the current classification, must be determined by the morphology of the structures. Meanwhile, the same microstructure frequently occurs in stromatolites with diverse morphologies characteristic of different groups. Because the classification of microstructures remains an unsolved problem, I have referred in the following analysis to the microstructures described at the time of the establishment of each particular form.

The distribution of stromatolites with particular microstructures in Riphean deposits of the Uchur-Maya region has recently been considered by Komar et al. (1973). Four types of distribution have been recognized:

(1) The same microstructure is characteristic of all the stromatolites of a suite or some portion of a suite over the whole area of its distribution. For example, a considerable part of the Malgina suite is composed of Malginella malgica Komar et Semikhatov and Stratifera of a similar microstructure. In the lower part of the overlying Tsipanda suite, Malginella, Stratifera and Irregularia have a common but different microstructure (that described from Malginella zipandica Komar). The boundary between the two microstructures coincides (not always exactly) with the sharp change in the composition of


the rocks at the boundary of the suites. The next level of microstructural change is within the uniform Tsipanda dolomites: in their middle part, there occurs Pamites aimicus (Nuzhnov) and Irreguluriu with a similar microstructure.

Y A

2

. CHELASIN

3

L +@

1)

- PLLYW-VU

4

Fig. 5. Vertical range of stromatolites with particular microstructures, in some sections of the Neruen and Ignikan suites (ranges in proportion to thicknesses). For position of sections see Fig. 1. Letters given in circles are stromatolites with microstructures described from the following forms: c = Conophyton cylindricum; It = C. lituum; mt = C. metula; s = Minjaria sakharica; u = Colonnella ulakia; i = Baicalia ingilensis; 1 = B. lacem; m = B. maica; t = Inzeria tjomusi; cn = I . confragosa.

The distribution of microstructures in the Ignikan suite is particularly interesting (Fig. 1). This suite is composed of clastic, microphytolitic, and chemogenic carbonate rocks of diverse composition in which dolomites typical for the Uchur-Maya platform are gradually replaced by limestones in the Yudoma-Maya trough. Yet the vertical sequence of stromatolites persists over the whole area of occurrence of the suite (Fig. 5). This allows three regional biostratigraphic subdivisions to be distinguished in the suite. The facies variations in the deposits consist only of some irregularities in the distribution of stromatolite bioherms and thickness variations in the beds containing stromatolites of a particular microstructure.

(2) Mass concentrations of stromatolites of some definite microstructure are confined to definite associations of deposits. The Yudoma suite (Fig. 2) can be used to illustrate this. The distribution of its stromatolites exhibits two opposite tendencies (Semikhatov et al., 1970). On the one hand,


Fig. 6. Schematic lithologic-and-facies profile of the lower Yudoma subsuite. For position of profile line see Fig. 1. Legend: I = sandstones; 2 = argillites; 3 = siltstones; 4 = dolomites; 5 = limestones; 6 = darkcoloured bituminous and clayey-bituminous limestones; 7 = sandy dolomites; 8 = variegated cherty-dolomite and clayey-herty-dolomite rocks; 9 = catagraphs; 10-1 1 = stromatolites with microstructures typical for the following forms: 10 = Boxonia grumulosa; I 1 = Jurusania judomica; 12 = K-Ar age of glauconites in m.y. (according to Kazakov and Knorre, 1970).

stromatolites with the same microstructure are present in rocks of diverse composition and genesis, for example in cross-bedded quartz sandstones, horizontally laminated clayey dolomites and massive pelitomorphic limestones. On the other hand, mass concentrations of stromatolites with a definite microstructure are confined to definite associations of rock types, which causes the stromatolites to be irregularly distributed. Thus, most


common among sandstones and microphytolitic dolomites are stromatolites which have the microstructures described for Boxonia grumulosa Komar, B. ingilica Komar and Semikhatov, Paniscolleniu emergens Komar and Colleniella singularis Komar; stromatolites with the microstructures described for Jurusaniu judomica Komar and Semikhatov occur in cherty and clayey dolomites which probably formed in the deeper sea; those with the characteristic Gongylina nodulosa Komar and Semikhatov microstructure are found in various types of pure carbonates. The boundaries of the rock associations with their contained stromatolite assemblages may be diachronous (Fig. 6).

Some of the rock associations occur repeatedly in the section due to the cyclic structure of the Yudoma suite. However, among its stromatolites there are forms which are confined to only one of the subsuites. Thus, Boxonia grumu Zosa and some other morphologically different stromatolites of a similar microstructure do not pass into the deposits of the upper subsuites despite the passage upwards of its characteristic rock association.

(3) Stromatolites of various microstructures are distributed in a disorderly fashion within a suite. Though some sequence in the change or alternation of microstructure can be observed in some sections, it is not maintained throughout the depositional area. The Svetly and Neruen suites can provide an example of this. Among the stromatolites of the Neruen suite, seven varieties of microstructures were distinguished (Fig. 5 ) . Five of them are encountered practically in the whole time-range of the suite, overlapping and replacing one another in vertical and lateral directions without any visible regularity. Moreover, in some cases one stromatolite structure may have different microstructures in its different parts (e.g. some Jacutophyton, see Serebryakov et al., 1972).

(4) Stromatolites occurring in different sections of the suites have different microstructures. This type of distribution has been found in the Omakhta suite in two isolated sections: in the Uchur depression and in the northern part of the Yudoma-Maya trough (see Fig. 1).

These considerations do not contradict the previously made suggestion that the spatial distribution of stromatolites of a certain microstructure reflects the distribution of communities of the stromatolite-forming algae. If this is so, it is unlikely that there was any direct dependence of the change of communities in time on environmental variations. This conclusion is supported also by vertical distribution of assemblages of morphological groups of stromatolites (see above). The condition in the Riphean basins of the Uchur-Maya region allowed, in some cases, the development of one or several coenoses over the whole of its ar6a; in other cases they were predominantly confined to limited areas.

In some comparatively rare cases, changes in the algal communities might be caused by the same environmental factors which affected the morphology of the stromatolites. This view is supported by the occasional coincidence in


a change of stromatolite morphology and change in their microstructure (some of the “Jucutophyton cycles”). In general, the following hypothesis may be proposed to explain the distribution of stromatolites in the Neruen suite (type 3): in the Neruen basin there existed at the same time five similar and closely related communities (species?) of algae which gave rise to stromatolites with five different microstructures. Each of the communities could develop within the entire range of the conditions present. For this reason the spatial distribution of the communities was irregular and their appearance in the given point of the basin was fortuitous. The difference in the microstructures of the Omakhta stromatolites (type 4 distribution), can be, in all probability, accounted for by the existence of two isolated basins populated by different communities (species?) of algae.

The same type of distribution of stromatolites with definite microstructures is observed in suites differing in composition (for instance, in the Svetly and Neruen) while different types are encountered in the lithologically similar suites: in the Svetly and Tsipanda, Neruen and Ignikan. This suggests that the realization of one type or another was determined not by the conditions of sedimentation but by the particular character of the algal communities.

CONCLUSION

(1) In the Uchur-Maya region, there were, during the Riphean, extensive shallow-water basins where the conditions of sedimentation were rather stable in space and time. As already pointed out in the literature (e.g., Keller et al., 1968; Trompette, 1969; Bertrand-Sarfati, 1972c) such basins were, evidently, characteristic of the Riphean.

(2) Stromatolites of the Uchur-Maya region formed subtidally as did, probably, the majority of other Riphean stromatolites (see Monty, 1973b, c; Serebryakov and Semikhatov, 1973,1974). They formed practically over the entire area of the basin. Therefore, u priori use of Riphean stromatolites as indicators of an intertidal or a nearshore environment is unfounded.

(3) Variations in the distribution of different communities (species?) of algae seem to have been the result of their different residence against environmental changes. The range of conditions allowing formation of stromatolites by one kind of alga or another was rather wide, which reinforces the interpretation of the subtidal formation of the stromatolites.

(4) The dependence of the distribution of the Uchur-Maya subtidal stromatolites on the conditions of sedimentation is poorly manifested. A moderate supply of terrigenous material did not limit the development of the stromatolites. It was not infrequent that short periods of terrigenous sedimentation failed to change the stromatolite assemblages (for example, in the Svetly and Neruen suites). Such changes, however, did occur within


homogeneous carbonates. Stromatolites were, normally, unaffected by changes in the carbonate composition of the rocks. Stromatolites of similar morphology and microstructure occur in limestone, sedimentary-diagenetic dolomites and mixed limestone-dolomite rocks. Thus, there is no reason to regard the Riphean stromatolites as indicators of definite salinity conditions in the basin.

(5) The morphology of the stromatolites, and sometimes also the distribution of the algal communities was greatly affected by environmental factors which are poorly, or not at all, manifested in the lithological characteristics of the rocks. The effect of these factors is more apparent in the analysis of the intrabioherm variability of stromatolites (see Ch. 6.4).

(6) Vertical variations in the algal communities and, as a result, in the morphological and microstructural assemblages of stromatolites, seem to have been associated mainly with the evolution of the stromatolite-forming algae. I t is not mere chance that a similar sequence of stromatolites has been observed in other Riphean sections within the U.S.S.R. (see Ch. 7.1).

(7) The conditions under which the Riphean stromatolites formed were, probably, essentially different from the conditions of growth of Recent algal structures.


Chapter 10.9

PALAEOENVIRONMENTAL AND GEOCHEMICAL MODELS FROM AN EARLY PROTEROZOIC CARBONATE SUCCESSION IN SOUTH AFRICA

K.A. Eriksson, J.F. Truswell and A . Button

INTRODUCTION

The lower part of the Transvaal Supergroup is largely nondetrital, made up in the main of carbonates and chert and lesser amounts of banded iron formations. It is developed in two structural basins, firstly around the Bushveld Complex in the Transvaal and secondly in the northern Cape (Fig. 1). Within these rocks several palaeoenvironmental models have been developed. In this paper these are related in space, defined wherever possible in terms of their geochemistry, and used in an analysis of the basin.

The rocks considered are about 2,250 m.y. old. Environmental aspects of the model may have relevance through the Proterozoic. But as will become apparent, aspects of the carbonate chemistry, and the presence of both widespread carbonates and banded iron formation in the basin, suggest that the geochemistry could only be fully developed in the time period 2,300-2,000 m.y.

STRATIGRAPHIC RELATIONSHPS

A full succession is developed in the eastern Transvaal. Here within the Olifants River Group carbonates and cherts of the Malman'i Dolomite form the bulk of the succession, and are overlain by the banded iron formations of the Penge Formation, above which are further carbonates and terrigenous clastics in the Duitschland Formation (Button, 1 9 7 3 ~ ) . The Malmani Dolo- mite has been subdivided into five lithostratigraphic zones in the eastern Transvaal (Fig. 2). Zones A , C and E are composed largely of darkcoloured carbonates and are essentially chert-free. Zones A and C are dolomitic; in zone E the dolomite is more ferruginous and there are remnants of limestone and minor development of proto-iron-formation and carbonaceous shale. The thicker zones B and D contain light-coloured recrystallized dolomite with abundant chert (Button, 1 9 7 3 ~ ) .

636 K.A. ERIKSSON ET AL.

G H A A P ond OLIFANTS RIVER GROUPS , .. . .. . . . .- ,

--- DEPOSITIONAL AXIS

z

[27 Iron Formation

. .. . SUBOUTCROP - - UPPER CONTACT

Fig. 1. Outcrop and locality map of the Ghaap and Olifants River groups.

-- Legend

Banded Iron Formation Limestone, Fe-rich Dolomite 8 Proto BIF

Fe 8 Mn-rich Dolomite Barrier-forming, clostic textured carbonates

Oolitic Barrier

Dark, Fe-poor Dolomite

Recrystallized Dolomite 8 chert

Shale

Black Reef Quartzite

Arenite

Control Sections (thickness in metres)

.Controlled 8 Projected Stratigraphic Boundaries Unconformit y

Palaeoerosional Zwartruggens / Zwortkops surfTe Penge

I rl

Fig. 2. Stratigraphic relationships in the Malmani Dolomite.


Zones B and D are truncated by chert breccias, interpreted as a chert rubble following short-lived subaerial exposure. Similar material, marking a more significant unconformity, occurs between the Olifants River Group and overlying sediments. This latter chert breccia truncates different horizons around the basin in the Transvaal. The Duitschland Formation is only locally present in the northeastern Transvaal near Penge. The Penge Formation can be traced westwards to west of Thabazimbi (Fig. l), and zone E of the Malmani Dolomite extends little further. The Penge Formation reappears near Zeerust, but traced eastwards it and zone E are again truncated (Grootenboer et al., 1974). In the Potchefstroom synclinorium and at Zwartkops the upper part of zone D has also been removed (Eriksson, 1972; Eriksson and Truswell, 1974a). Proceeding southwards in the eastern Transvaal the Duitschland and Penge Formations and zones E , D and C are all successively truncated by the unconformity.

Throughout most of the Transvaal the Black Reef can be regarded as a basal clastic member of the Malmani Dolomite (Eriksson and Truswell, 1974a); but in the northeastern Transvaal it locally thickens within an early proto-basin and is there afforded formational status (Button, 1974).

In the northern Cape the Black Reef is overlain by the Ghaap Group (Beukes, 1973). Within this the Campbell Rand Dolomite and Kuruman Iron Formation are regarded as equivalents of the Malmani Dolomite and Penge Formations, respectively. The Griquatown Jasper lies conformably above the Kuruman Iron Formation, and above that is an unconformity. The Campbell Rand Dolomite is poorly exposed beneath the Ghaap Plateau but in general contains less chert and more limestone than the Malmani Dolo- mite. The Formation commences with a clastic unit, which grades up to limestones and ferruginous dolomites, followed by a thick iron-poor dolomite unit (Visser and Grobler, 1972; Truswell and Eriksson, 1973). In the vicinity of Prieska the Campbell Rand Dolomite is similar in nature to zone E of the Malmani Dolomite, and comprises ferruginous dolomite, proto-iron- formation and shale.

This chapter is concerned with aspects of the Malmani Dolomite, Campbell Rand Dolomite, and Kuruman and Penge Iron Formations, but offers no comment on either the structurally disturbed Duitschland Formation or on the Griquatown Jasper Formation.

PALAEOENVIRONMENTAL MODELS

A number of depositional environments have been recognized in these rocks. Their distinction as facies leans heavily on Walther’s (1894) Law which recognises that unbroken vertical sequences are a reflection of lateral facies. The environments established involve uniformitarian analogy where possible, but of necessity conceptual models are more commonly applied.


The basal marine transgression

The base of the Malmani Dolomite represents a classical marine transgressive unit which frequently commences with a conglomerate veneer passing up to orthoquartzitic beach and shallow marine sands (Button, 1974). The sands grade up to and interfinger with carbonaceous shale and structureless dolomite (rich in iron and manganese), which grades in turn into dolomites containing what are believed to be subtidal domes (Truswell and Eriksson, 1975). The thickness of the basal arenaceous unit (the Black Reef Quartzite) is largely dependent on the configuration of the pre-Transvaal topography. The formation grades laterally into cobble conglomerate adjacent to what were islands during the transgression, and may wedge out altogether over the crest of an island.

The tidal-flat model

Around Zwartkops (Fig. l), an intertidal assemblage is developed within light-coloured chert-rich dolomite of Malmani zone D . In this assemblage, finely laminated sediments are ‘associated with small-scale symmetrical algal domes, flat-pebble breccia and bird’s eye and palisade structures (Eriksson and Truswell, 197413). These features are analogous to those in contemporary tidal-flat sediments in protected intertidal to marginal supratidal environments, as for example at Shark Bay, Western Australia (Davies, 1970b; Logan e t al., 1974).

Larger domes with relief ranging to more than 1 m are the most common large-scale structures of zones B and D. Increases in dome size are thought to reflect successively greater intertidal depths (Logan, 1961; Monty, 1972). The elongation of domes probably indicates the influence of tidal currents, an interpretation supported by the presence of breccia between domes, lenses of oolite and by the sporadic development of linear and ladder ripples. Intertidal conditions are also suggested by mud-cracked clay drapes over ripple marks and by small stromatolite domes swamped by rippledominated carbonates (Truswell and Eriksson, 1975).

The more steeply shelving slope

At Boetsap in the northern Cape (Fig. l), a complex of subenvironments has been related to a steeply shelving slope. The model is a modification of Irwin’s (1965) limestone shelf model. In addition, the concept that stromatolite domes increase in size outwards into the intertidal has been extended into the subtidal regime (Truswell and Eriksson, 1973).

In this model, sedimentation occurred in a telescoped range of environments, extending from the intertidal, through a highenergy agitated zone, into the subtidal. Columnar stromatolites characterize the intertidal, but


more delicate columns are also developed within domes in the shallow subtidal environment. In the agitated zone, the carbonates contain breccias, ripple marks, oolites and oncolites. Deeper subtidal carbonates are dominated by elongate stromatolitic mounds, up to 40m long and 1Om wide, with a vertical inheritance of up to 13m, and a relief of up to 2.5m. The microstructure of these mounds is a delicate crinkled lamination.

Elongate mounds of a similar order of size have been described from zone C of the Malmani Dolomite at Zwartkops and elsewhere in the Transvaal. At Zwartkops they are considered to have formed offshore from aripple-marked tidal-flat assemblage. Talus breccia marks the assumed steeper slope between the prograding tidal flat and the subtidal domes (Truswell and Eriksson, 1975).

The lagoonal environment

In the eastern Transvaal, iron formation has been shown to be an end member in a succession of carbonate sedimentary cycles (Button, 1976). It has been shown that iron formation was probably precipitated from solution in a barred basin or lagoon. The barrier behind which the precipitation took place was formed by carbonate detritus consisting of oolitic and intraclastic breccia. Transgressive and regressive shifts of the barrier resulted in sedimentary cycles in which iron formation, iron formation precursor and carbonate detritus are interstratified with limestones and iron-rich dolomites. Near the base of the Malmani Dolomite at Zwartkops a lagoon which developed behind an oolite barrier has been recognized (Truswell and Eriksson, 1975). In contrast to the lagoon described above, rock types here are confined to iron- and manganese-rich dolomites (Fig. 2).

PALAEOENVIRONMENTAL AND GEOCHEMICAL SYNTHESIS

As an introduction to a synthesis of the palaeoenvironmental models and their geochemistry it is necessary to set the general scene during the formation of these rock types in Early Proterozoic times. It is envisaged that sedimentation took place in an epeiric sea during a marine transgression across the craton. The atmosphere at. that time is considered to have been oxygen- deficient (Cloud, 197313; Garrels et al., 1973) and probably richer in COz than today. The atmosphere appears to have become more oxidizing from 2.0 b.y. onwards (Cloud, 1973b); thus, the synthesis which will be presented here has application only to sedimentary basins greater than 2.0 b.y. in age.

From the palaeoenvironmental models discussed it is possible to recognize four major lithologic associations; namely: banded iron formation, limestones with ferruginous dolomites, limestones with iron-poor dolomites, and recrystallized dolomites and chert. The limestones and ferruginous dolomites


are those which constitute the barrier to the iron formation lagoon, the carbonates which formed immediately seawards of the barrier, and also the intertidal facies at Boetsap. The deepest subtidal carbonates are iron-poor pure dolomites; those which have been interpreted as shallower subtidal sediments at Boestsap are also iron-poor and contain limestone remnants; these are grouped together as limestones and iron-poor dolomites. It is now necessary to relate these associations in space. Isopach data for the Olifants River and Ghaap groups are sparse away from the eastern and central Transvaal. Indications are, however, that the axis of the depositional basin had a roughly northeast to southwest orientation (Fig. 1). Thus, much of the original northwestern extent of the basin in the Transvaal and all of the southeastern margin in the northern Cape have been removed by erosion.

, The preserved stratigraphic thickness in the Transvaal is composed almost wholly of recrystallized dolomites with chert and iron-poor dolomites with large domes, the other two associations only appearing at the top of the succession along the northwest margin of the basin (Fig. 2). The very shallow- water structures often developed in the recrystallized dolomites and chert suggest that deposition took place close to the basin-edge while towards the depositional axis subtidal conditions prevailed in which the iron-poor dolomites with domes developed. The intertidal and subtidal domes are all elongated in a northwesterly direction (Truswell and Eriksson, 1975), providing further evidence for a northeast to southwest orientation of the depositional axis (Hoffman, 1967). Furthermore, the intertidal domes in the southeast are oversteepened to the northwest, again indicating a basin-edge to the southeast.

In contrast to the Transvaal, lithologies in the northern Cape are almost devoid of recrystallized dolomites and cherts. Banded iron formations with which substantial thicknesses of shales are associated, pass through clastic- textured carbonates into limestones and ferruginous dolomites and eventually into iron-poor dolomites. This sequence is considered to represent increasing distance from a basinedge to the northwest. A lagoon in which iron formations were formed was separated by a elastic-textured carbonate barrier bar (Button, 1976) from a more open-marine regime where limestones and iron-rich dolomites graded seawards eventually into iron-poor

Barrler-formlng N W CIastlc Textured Carbonates S E

L.W L.-

Recrystalltzed Dolomltes

Dolomites

Fig. 3. A palaeoenvironmental reconstruction.


dolomites with domes (Fig. 3), the latter being the only unifying lithology between the Transvaal and northern Cape. The four recognized lithologic associations thus represent distinct facies.

Before discussing the geochemistry of the four facies it is necessary to set some probable constraints on the composition of surficial waters draining into the Proterozoic sea. The reducing nature of the atmosphere at that time has already been mentioned and the fact that both Fe and Mn in the rocks under consideration are principally divalent strongly supports this contention (Eriksson et d., 1975). In the absence of oxidation of Fe and Mn it is considered likely that these metals would have been dissolved in river waters in their crustal abundances relative to Ca. If it is assumed that the concentration of Ca in surface waters has not changed significantly through geologic time, then the waters draining into this Early Proterozoic sea probably had Ca, Fe and Mn concentrations of around 15, 20, 35 ppm. An atmosphere significantly richer in COz would have affected the absolute amounts of metals in solution but their relative concentrations would have remained essentially the same. With these concentrations, solubility graphs predict that on increasing pH, Fe carbonates would precipitate first, followed at higher.pH’s by Fe-Ca and finally Fe-Ca-Mn carbonates (Eriksson et al., 1975), the amount of Fe in solution decreasing with time.

It is contended that these three predicted geochemical facies are represented in the northern Cape by iron formations, limestones and iron-rich dolomites, and subtidal iron-poor carbonates, respectively (Fig. 3). These are considered to have developed away from an active northwestern basin-edge along which not only substantial amounts of argillaceous sediments but also high concentrations of cations, most notably iron, were being introduced into the depository. The source of Fe, in iron formations has always been a controversial subject. Eugster and Ming Chow (1973) consider that in the absence of free atmospheric oxygen sufficient ferrous iron would be carried in solution to account for iron formations. This becomes even more plausible when the tectonic environment in which these Early Proterozoic iron formations formed is considered. The extremely stable and low-lying hinterland would have been an ideal site for deep chemical weathering with the release of Fe and other cations. A tectonic pulse would have flushed out the ground- water causing it to flow basinwards, carrying with it a suspended clay fraction and dissolved or colloidally dispersed .iron and silica (Button, 1976). Whereas the iron formations and limestones with associated ferruginous dolomites developed in and close to a barred lagoon, the limestones with iron-poor dolomites formed in a more disw subtidal environment which was least influenced by the high concentrations of Fe being washed in from the land.

The iron-rich dolomites are invariably associated with limestones and on the basis of field evidence a two-stage process for their formation is envisaged, involving: (1) the precipitation of primary limestone, and (2) the dolo-

K.A. ERIKSSON ET AL.

Bonded Iron twmatlons ( overage FeOIMnO rotlo)

No. Of analyses

Tldal flat dolomites 112 Limestones and Fe-rlch dolomites 31 Subtidal limestones and Fe-poor dolomites 32

mitization of this limestone by solutions rich in both Fe and Mg and also containing significant amounts of Mn (Button, 1976). In contrast, the deeper subtidal carbonates which are composed of pure dolomites formed either primarily or, as seems more likely, through alteration by normal epeiric sea waters. The subtidal carbonates furthest up the paleoslope contain most Mn (Fig. 4) and probably formed as that metal became saturated in solution, substantial amounts of manganese having been taken into the calcite lattice (Eriksson et al., 1975).

A major portion of the Malmani Dolomite is composed of an intertidal facies, in which pure recrystallized dolomites alternate with chert. The recrystallized dolomites have low and variable Fe and Mn concentrations and erratic Fe/Mn ratios (Fig. 4). The recrystallization and diagenetic formation of chert are considered to reflect the influence of low-pH meteoric waters (Truswell and Eriksson, 1975). Both processes are promoted by dissolution by meteoric waters, a phenomenon which also enhances dolomitization (Badiozamani, 1973). In contrast to the northwestern margin, that in the southeast appears to have acted in a passive manner with only minimal

4 -

3 .

0

0

0 0 0

0 0

0 0

2 . 0 8 O

. . . a

0 a0 0 . ..

1 0 1 2 3 4 5 6 7 8

MnO %

Fig. 4. FeO vs MnO plot illustrating the grouping of points into four geochemical facies.


amounts of clastics and cations introduced from this direction. A similar phenomenon has been observed for clastic sedimentary basins on the craton in which maximum supply of sediments was always from the northwest (Pretorius, 1974). The similar lithologic association to that of the northern Cape which occurs at the top of the succession in the Transvaal is considered to have developed during a southward overlap of the northern iron formation/iron-rich dolomite facies, the majority of which has been removed by erosion. The Fe-Mn-rich lagoonal assemblage which is present towards the base of the Malmani Dolomite is somewhat problematical. It is known that clastic sedimentation was a factor in early Malmani times and it is conceivable that cations could have been intrdduced in some abundance prior to the development of the broad tidal-flat environment.

In summary, there are considered to be three distinct environments of dolomitization, which resulted in chemically distinct dolomites. Immediately seawards of the lagoon iron-rich dolomites developed. The preservation of limestones in these rocks is considered to relate to a freshwater dilution with respect to Mg and Fe, resulting in incomplete dolomitization. In the tidal-flat regions, the presence of lower-pH meteoric waters both enhanced dolamitiz- ation, and catalyzed the diagenetic formation of chert and recrystallization of dolomites. These dolomites are characteristically poor in Fe and Mn. In the deeper subtidal environment, dolomitization was effected by the alkaline waters of the epeiric sea which had attained a critical Mg/Ca ratio.

The vertical stratigraphic relationships observed are related to transgressions and regressions in the epeiric sea. The most dramatic evidence of such a phenomenon occurs towards the top of the succession where proxi- mal lagoonal banded iron formations have come to rest on more distal marine carbonates as a result of a major regression or overlap towards the southeast. In Western Australia rocks of a similar age make up the Hamersley Group and are dominated by iron formations with only a minor development of carbonates. Carbon-isotope determinations led Becker and Clayton (1972) to conclude that those iron formations were precipitated in a basin isolated from the ocean but probably in close proximity to it, and that the carbonates developed in marine environments. The interlayering of the iron formations with carbonates was attributed to transgressions and regressions.

Many shallow-water cratonic carbonate sediments give way to a deeper- water clastic facies (e.g. Hoffman, 1974). No evidence of such a relationship was observed in the present study, with an epeiric sea being the only recognizable depositional setting. I t is feasible that deep-water conditions may have existed to the south of Prieska. The-recognition of such a facies, however, is most unlikely as a 1.0 b.y. old metamorphic belt occurs in this area. It is also presently unclear what such metamorphic belts represent in terms of Precambrian plate-tectonic movements.


Chapter 11.1

MINERAL DEPOSITS ASSOCIATED WITH STROMATOLITES

Felix Mendekohn

INTRODUCTION

Many mineral deposits are closely associated with stromatolites which occur either as bioherms (reefs), or biostromes that formed as widespread stromatolitic beds or algal mats. It is likely that there are many more than the ones dealt with here, but these are sufficiently numerous and widespread to suggest that they give a fair sampling of the class. The information available on different deposits shows a wide range in quality and quantity, probably reflecting the knowledge of and interest in the stromatolites themselves, the degree of preservation, and the effects of metamorphism. The term reef has been rather loosely used in many discussions of ore deposits to indicate any mound-, ridge-, or reef-like organic structure; since it is not always possible to be certain of their precise meaning, “reef” is used as the authors use it, and should be understood to include bioherms as well as reefs in the restricted sense of protruding wave-resistant structures. In the descriptions that follow, deposits are grouped according to the valuable element(s) they contain; a summary is presented in Table I.

COPPER

The association between stromatolites and ore deposits is well demonstrated and documented in the copper province of Zambia/Zaire. On the Zambian Copperbelt the two main groups of ore deposits are those in argillite (ore shale) and those in quartzite. Bioherms present in many of the ore-shale deposits were not recognized as stromatolitic rocks for a long time because the moderate grade of metamorphism, local shearing along the stromatolite zone, and supergene alteration to talcose cherty dolomite resulted in the partial destruction of the internal structure. However, the intersection of a number of well-preserved bodies at depth leaves little doubt that all or most are stromatolitic. Stromatolitic bioherms, forming a nearshore dolomitic limestone facies of the argillaceous Ore Formation (Garlick,

646 F. MENDELSOHN

Fig. 1. Portion of the Roan Antelope orebody, in the vicinity of the Irwin shaft barren gap, shown unrolled and restored as nearly as possible to its attitude and relations before folding, but with the vertical scale much exaggerated. The barren dolomite, considered to be part of a stromatolite bioherm, occupies the whole of the Ore Formation and rests on a hill of basement rocks; beyond, the Ore Formation rests on footwall beds. The barren dolomite thins to a feather edge and is then replaced by argillite; beyond, the widespread carbonate at the base of the Ore Formation, containing copper/iron sulphides and forming a second orebody where marked “mineralized”, is an impure dolomitic limestone representing an original algal mat. The copper-rich zone in the argillite spreads outward from the bioherm and away to the right like a plume, but decreases to normal grade in this direction.

1964), occur in many instances over ridges or hills of basement that formed promontories or islands, projecting to or slightly above the base of the Ore Formation during deposition. The bioherms themselves are barren of copper and related minerals except in shaly inclusions, but the adjacent ore shale is richly mineralized, the metal content decreasing outwards and being reflected by a mineral zoning of bomite-chalcopyrite-pyrite (Fig. 1; Garlick, 1964). A number of stromatolitic bioherms are now known for 50 km along the eastem margin of the ore shale deposits, from Roan Antelope through Nkana and on to Chambishi (Annels, 1974; Clemmey, 1974); associated layers of impure dolomite, notably at the base of the Ore Formation (Fig. l), are considered to represent original algal mats with included detritus. The

MINERAL DEPOSITS 647

Fig. 2. Stromatolite bioherms at Mufulira, in argillite immediately beneath the quartzite of the ‘B’ orebody. It can be seen how stromatolite growth started in the lower shale, but was inhibited and almost stopped by a thin layer of sand (white); in the succeeding dark shale growth again became vigorous, t o form the bioherm, which is about double the width shown. The ‘B’ orebody quartzite can be seen at the top, starting at the handle of the hammer; the quartzite immediately above the bioherm is mottled for a few feet by poorly developed algal growths.

zone of bioherms and sandy sediments marks the eastern shoreline, offshore is a zone of shale about 2 km wide that contains the copper orebodies, and beyond is a 3-8 km zone of pyritic carbonaceous shale; beyond this again are (deeper-water) carbonates (Garlick, 1961, p. 157), or terrestrial sediments (Annels, 1974). These dimensions and rock types indicate that deposition took place in an interdeltaic coastal marine lagoonal environment (Hamblin, 1973). Van Eden (1974) concludes that broadly similar conditions of deposition, with offshore shoals, existed during the deposition of the sands that now form the Mufulira C orebody, one of the major quartzite deposits. The quartzite orebodies are in general devoid of stromatolites, but the first stromatolites recognized on the Copperbelt were in a bioherm in argillaceous beds immediately below the Mufulira B orebody fringe (Malan, 1964; Figs. 2, 3); Paltridge (1968) found that an extensive algal biostrome, containing bodies of less well-developed stromatolites, spread away from the original occurrence. Malan (1964) reports a low copper content in the

648 F. MENDELSOHN

Fig. 3. Close-up of stromatolites of the Mufulira bioherm, in shale, showing shape and structure. The dark lines in the stromatolites are micaceous material similar to the adjacent argillite and represent sediment that collected on the flat tops of the stromatolites during growth; little or none is present on the steep flanks.

bioherm, proportional to the amount of argdlaceous material included in or between the stromatolites; Paltridge reports a similar low copper content in the carbonate part of the biostrome, but as much as 7% in fine-grained algal- bedded sandy sediment, which in places forms the foundation for stromatolites.

In the continuation of the copper-bearing Katanga formations in the Shaba (Katanga) province of Zdire, the Sene des Mines is considered to be the stratigraphic equivalent of the ore-bearing formations of Zambia (Franqois, 1974). The ore deposits within the Sene des Mines consist of two mineralized formations, the upper a dolomitic shale like that of the Zambian argillite deposits, and the lower a finely laminated siliceous dolomite (Oosterbosch, 1962). The intervening regular bed of dolomitic limestone in places consists wholly or partly of stromatolites, identified as Collenia undosa (Oosterbosch, 1962, p. go), forming an algal biostrome; in places bioherms protrude into the overlying dolomitic shale (Fig. 4). As in Zambia the stromatolite formation is barren of copper, except for occasional inclusions of cupriferous shale. The stromatolites here are cylindrical, and on weathering are partly replaced by silica to form hollow cylinders and the rock is known as “roches siliceuses cellulaires” (Fig. 5 ) .


Fig. 4. A bioherm in the Musonoi open pit, Shaba, Zaire, projecting about 4 m into the dolomitic shale of the upper orebody, which is in part draped over the bioherm. The underlying stromatolitic carbonate formation (roches siliceuses cellulaires, RSC) from which the bioherm grew is hidden under the broken rock on the right.

At Matsitama in Botswana, in an isolated basin of metamorphosed formations that might be correlated with the Katanga, a deposit of potential economic significance occurs in quartz-carbonate rock, dolomite, and shale. A quartz-carbonate body that is probably a bioherm is more richly mineralized with copper than the laterally adjacent shale and biostromal beds.

At Mount Isa, Queensland, the Urquhart shales of the Middle Proterozoic Mount Isa Group (1,600 m.y.) contain important lead-zinc and copper deposits; the Urquhart shales consist of dolomitic shales and siltstones, dolomite, and tuffaceous shale. Galena, sphalerite, and pyrite occur in distinct concordant layers throughout the Urquhart shales, and where sufficiently concentrated, form the fourteen known orebodies, which are arranged en echelon, plunge south, and have in places suffered folding and faulting (Bennet, 1965). There is a broad but distinct mineral zoning, from pyrite in the north into sphalerite-pyrite-pyrrhotite, then rich galena-sphalerite orebodies that terminate against silica-dolomite bodies. The silica-dolomite contains pyrite, pyrrhotite, and chalcopyrite, and in places close to the lead-zinc bodies but separate from them, are the copper orebodies. The coarse ‘disordered-looking’ silica-d olomite bodies have been interpreted as

650 F. MENDELSOHN

Fig. 5 . Siliceous tubes of the RSC stromatolitic carbonate that led to the naming of the formation the “roches siliceuses cellulaires”. The outer, presumably purer carbonate portions of the columnar stromatolites (see Fig. 3) were preferentially replaced by silica, and the interior portion was weathered away; remnants of the stromatolite layering can be seen inside the tubes.

being algal reefs and reef-breccias, the finely bedded shale being off-reef organic tuffaceous sediment (Garlick, 1964; Stanton, 1972).

In the Middle Proterozoic Belt Supergroup in Montana, a stratiform copper deposit has been found in feldspathic quartzite and siltstone of the Revett Formation, and subeconomic or unproven copper- silver concentrations are known in various places through the higher Belt formations (Harrison, 1974). In the Helena Formation of the middle Belt carbonate sequence, disseminated copper minerals are associated with algal stromatolite mounds and other structures of probably algal affinity (Fig. 6).

The Infracambrian formations of the western Anti-Atlas of Morocco, consisting of sandstones, schists, and limestones, contain a number of small copper deposits at various stratigraphic horizons, many on the flanks of basement highs (Chazan and Fauvelet, 1962). The Tamjout series, consisting of recrystallized reef dolomites containing Collenia, overlain by a layer containing lenses of white quartz, is widely but weakly mineralized with lead, zinc, vanadium, and pyrite, and has several copper deposits.


Fig. 6. Stromatolitic mounds in the Helena dolomite of the Belt series, Montana, U.S.A. In the cliff face behind the mounds are other features that are probably algal, and disseminated copper minerals occur in similar rocks nearby.

In the Bijawar formations of Madya Pradesh, in an old copper-mining district, Collenia-bearing dolomite is overlain by a more siliceous layer containing cyclindrical Conophyton (Balasundaram and Mahadevan, 1972). Chal- copyrite and galena occur disseminated and as veinlets in the formations containing the stromatolites, in places within the stromatolite zones.

LEAD-ZINC

Lead-zinc (barite- fluorite-copper-cobalt) deposits are widespread in the Mississipi Valley of the United States, the great majority in carbonate rock ranging from limestone to dolomite (Beales and Onasich, 1970), and the class of deposit represented here is known as the Mississipi Valley type. The deposits are spread through almost the complete range of Phanerozoic rocks, but over 80% of the ore lies in formations of Cambrian-Ordovician and Mississipian -Pennsylvanian age. Stromatolite reefs are associated with a number of the deposits that are present in Late Cambrian and Early Ordovi- cian formations in southeast Missouri (Wagner, 1961; Myers, 1966; Snyder

652

PLAN

F. MENDELSOHN

0 30 60 r

S ECT 10 N rn VERTICAL EXAGGERATION - l o x

/ RIDGE

Fig. 7. Stromatolite reef over topographic high in the southeast Missouri lead belt; lead minerals occur mainly on the upper and lateral margins of the reef, and also in the overlying formations (after Wagner, 1961).

and Gerdemann, 1968) and to a lesser extent in Tennessee (Kendall, 1960), but in the upper Mississippi district stromatolites are absent from the ore formations (Heyl, 1968). The southeast Missouri lead belt is situated on the flank of a basement high and here Late Cambrian reefs, now interpreted as algal banks (Larsen, 1973), of digitate or columnar Collenia or Gymnosolen, lie over ridges of either basement rocks or Early Cambrian sediments (Fig. 7), accompanied by widespread discontinuous biostromes formed from stromatolite mounds. The sulphide minerals are closely associated with the stromatolitic formations, the major concentrations commonly being along the tops or edges of the elongate reefal structures. Associated black pyritic shales interrupted in places by patch reefs represent the muddy deposits accumulated in lagoons within fringing reefs (Myers, 1966). In east Tennessee, zinc- sulphide concentrations associated with reefs in the Late Cambrian Longview


Formation form a significant proportion of the ore in some mines in this area. These reefs are 3- 5 m across and up to 2 m high and are built up of fairly even laminae, without columnar structures, but are considered to be stromatolitic (Kendall, 1960); sphalerite occurs mainly as distinct caps, up to 30 cm thick, on the reefs, with some along other margins but little internally.

On Baffin Island, Northwest Territories, Canada, the Middle Proterozoic Society Cliffs Formation is predominantly an algal carbonate that is brecciated and contains specks and massive bands of sphalerite, galena, and pyrite (Geldsetzer, 1973). The carbonate consists mainly of finely to irregularly laminated stromatolitic dolomite with occasional biohermal and nodular structures, interpreted by Geldsetzer as being of subtidal to intertidal origin.

In Bahia, Brazil, lead-zinc orebodies in the Bambui are associated with Collenia (Cassedanne, 1966,1969).

PHOSPHATE

In the Aravalli metasediments of Udaipur, India, phosphorite is.closely associated with stromatolitic dolomite overlying orthoquartzite (Banejee, 1971b), the stromatolites concerned being Collenia symmetrica, Minjaria, and Baicalia. In detail, the phosphate occurs as: concentrations in stromatolite columns, laminar algal (stromatolitic) phosphorite, reworked fragments of stromatolites forming silicified conglomeratic or brecciated phosphorite, massive-bedded phosphorite with sandy and clayey laminae, and disseminated pellets and nodules in dolomite. Phosphorites are widely distributed, the P,O, content ranges from 10% to 37%, and the major ore mineral is apatite. Bannerjee also refers to the occurrence of Collenia and oncolites in other Late Precambrian phosplioritic formations in India (Gangolihat Dolomite), Russia, and China.

IRON

Several of the major deposits of iron formation of Early Proterozoic age, around 2,000-2,200 m.y. old, are associated with stromatolites and carbonate rocks, and there is evidence that there was algal and bacterial activity in a substantial number of them (Cloud, 1973; La Berge, 1973).

The Biwabik Iron Formation in the Mesabi district of the Lake Superior iron province contains two zones of cherty columnar stromatolites (Bayley and James, 1973), and underlying dolomite also contains stromatolites. On the Ontario side of Lake Superior, the equivalent Gunflint Iron Formation consists of interbedded stromatolitic chert, bedded chert, taconite, jasper, limestone, and black shale (Hohann, 1969b; Douglas, 1970, p. 118); Moorehouse and Beales (1962) describe jaspery and carbonate stromatolites

654 F. MENDELSOHN

(Collenia and Gymnosolen ?), and also “reefs” (bioherms), one with a core of stromatolitic chert and jasper but a surface consisting mainly of calcarenite and taconite. In black shale of the overlying Rove Formation at Kakabeka Falls powerhouse a cherty reef-like structure has a capping of pyrite about 5 cm thick, and oncolites occur nearby in the shale. At Steep Rock, Ontario, iron ore in the Archaean to Early Proterozoic formations (Jolliffe, 1966; Cloud, 197313) overlies dolomite containing columnar stromatolites. In South Africa, iron formation of the Transvaal sequence overlies a thick dolomite that contains abundant stromatolites (Truswell and Eriksson, 1972), and stromatolitic limestone occurs immediately beneath the iron formation in both the Cape and Transvaal Provinces. In the Cape region, pyritic stromatolitic limestone and pyritic carbonaceous shale underlies the ferruginous chert and pyritic carbonaceous shale of the transition zone (Beukes, 1973). In the eastern Transvaal the transition zone between the two consists of interbedded stromatolitic limestone, pyritic dolomite and mudstone, and iron formation; the limestone contains laminar, columnar, and domal stromatolitic structures (Button, 1973c), and Button (1976) also emphasizes the relation between carbonate and iron formation.

MANGANESE

At Kisenge in the western part of Shaba province, Zaire, a stratiform manganese deposit in sericite schists of the Lukoshi Complex is cut by a pegmatite dated at 1,850 m.y. (Doyen, 1973). Below the oxide zone, the deposit consists of alternating beds of manganese carbonate, “gondite”, and graphite schist; the carbonate consists of almost pure rhodochrosite, and the gondite of spessartite, a manganese-bearing garnet. The carbonate ore, with about 40% manganese, contains spessartite and small amounts of nickel and cobalt sulphides. Within the carbonate ore are several layers of stromatolites that are stated by Doyen to resemble Collenia. Doyen concludes that the manganese was of sedimentary origin and that subsequent metamorphism led to the development of garnet in the more impure carbonate layers.

In the Transvaal Black Reef Formation in Botswana small stromatolites consisting of and embedded in manganese oxide, are associated with a layer of manganese nodules having a general similarity to those being formed today on the ocean floors; it is not known whether the manganese oxide replaces carbonate or is an original constituent (Litherland and Malan, 1973).

Monty (1973) found that matiganese nodules from the Blake Plateau and the south Atlantic have a very fine wavy and regular lamination, characteristic of stromatolitic structures, and consisting of fine dark brown and light- coloured laminae; in the dark laminae, 6-8 pm thick, are manganese-bearing bacterial rods and tiny filaments. The nodules result from the rhythmic


growth of filamentous bacteria concentrating manganese and iron minerals within their sheaths, and it is possible that regional variations in composition might be dependent on the biochemical activity of the bacteria. Monty considers that they can be regarded as deep-water oceanic stromatolites.

GOLD AND URANIUM

The Early Proterozoic (2,300-2,600 m.y.) Witwatersrand formations of South Africa consist of an upper arenaceous division and a lower more argillaceous division. Most of the gold-uranium deposits occur in the lower part of the upper division, in or adjacent to conglomerate layers at the bases of cycles of sedimentation (Pretorius, 1974). In the western and southwestern parts of the goldfield, the bulk of the gold and uranium is present in carbon seams at the top of certain of the cycles of sedimentation; Pretorius (quoted in Payne, 1974) estimates that in recent years, 40% of the gold produced came from carbon seams, and the proportion is probably increasing. These carbon seams are considered to represent algal mats (Pretorius, 1974) or, according to Hallbauer and Van Warmelo (1974), mats of a material resembling modem lichen and probably consisting of algal and fungal organisms. Hallbauer and Van Warmelo present evidence to show that gold and uranium were assimilated by these plants during growth, and also physically trapped as detrital grains between carbon columns.

In the Rum Jungle area of Northern Territory, Australia, uranium and associated copper, lead, and gold deposits occur in Early Proterozoic (1,800-2,200 m.y.) formations. The deposits are closely associated with silicified limestone breccias, partly reefal in origin, or occur in sediments of off-reef facies, siltstone and carbonaceous shale with intercalated limestone lenses (Condon and Walpole, quoted in Garlick, 1964). Crohn (in Walpole et al., 1968) reports both agreement and disagreement with the interpretation of a reefal environment by subsequent workers. Based on his work at Nabarlek, Cooper (1973) suggests that regional uranium mineralization took place in Late Proterozoic times (710-815 m.y.).

FLUORITE

In the western Transvaal, South Africa, fluorite is present in significant amounts at the top of the Dolomite of the Transvaal Supergroup, in what was a minor lead-zinc district in the past. The presence of abundant stromatolites in the Dolomite has already been referred to, and a good deal of the fluorite in this area is associated with stromatolites, economic concentrations being largely confined to stromatolite bioherms.

TABLE I

Summary of data on mineral deposits and associated stromatolites

Formation Age Lo cat ion Form of Elements stromatolites

Nodules Bonneterre

Longview Tamjout Unknown Unknown Aravalli

Gangolihat Bambui Katanga Katanga

Society Cliffs Belt Mount Isa Matsitama Gunflint

Biwabik Steep Rock ?

Transvaal Transvaal

Lukoshi Bijawar

Witwatersrand

Recent L. Cambrian

L. Cambrian Infracambrian Middle Riphean Sinian Precambrian (?)

Middle Riphean Middle Riphean L. Proterozoic : 900 m.y. L. Proterozoic: 900 m.y.

L. Proterozoic: 1,000 m.y. M. Proterozoic: 1,200 m.y. M. Proterozoic: 1,700 m.y. L. Proterozoic (?) E. Proterozoic: 2,000 m.y.

E. Proterozoic: 2,000 m.y. E. Proterozoic/Archaean E. Proterozoic: 1,800-2,200 m.y.

E. Proterozoic: 2,250 m.y. E. Proterozoic: 2,250 m.y.

E. Proterozoic: > 1,850 m.y. Precambrian

E. Proterozoic: 2,500 m.y.

Ocean SE Missouri, U.S.A.

Tennessee, U.S.A. Morocco Russia China Rajasthan, India

Pithoragarh, India Bahia, Brazil Copperbelt, Zambia Shaba, Zaire

N.W.T., Canada Montana, U.S.A. Queensland, Australia Botswana Ontario, Canada

Minnesota, U.S.A. Ontario, Canada Northern Territory,

Australia Cape Province, S. Africa Transvaal, South Africa; Botswana Shaba, Zaire Madya Pradesh, India

South Africa

Bacterial Collenia (Gymnosolen) ? Collenia Collenia oncolites Baicala Minjaria Collenia

? Collenia Collenia Collenia

? ? ? ? Collenia Gymnosolen (?) ? ? ?

sy m me trica

undosa

? ?

Collenia Collenia Conophy ton stratiform

Mn, Fe (Cu, Co, Ni) Pb, Zn

Zn c u P P P

P Pb (Zn) c u (CO, U) cu , co (U)

Fe Fe 9

U (Cu, Pb, P)

Fe Fe, Mn

Mn (Co, Ni) Cu, Pb

Au, U


ANALOGOUS DEPOSITS ASSOCIATED WITH OTHER BIOHERMS (Table I)

A substantial number of mineral deposits in Phanerozoic host rocks are associated with bioherms built by corals, rudistids, orbitilinids, stromatoporoids, and other animals, or with bodies loosely referred to as reefs, bio- hems, etc., and with unknown (or partial algal) affiliation. Because of their general similarity to many of the deposits in stromatolitic formations already described, some of these are listed or briefly described below.

Lead-zinc deposits in or closely associated with reefs of Devonian or Carboniferous age are those of Pine Point, Canada (Campbell, 1967); Meggen, West Germany; and Tynagh and Silvermines, Ireland (Pereira, 1967). In reefs associated with lead- zinc deposits in the Canning Basin, Australia, Monseur and Pel (1973) list the reef-building organisms as stromatoporoids (see also Ch. 10.4 herein). Maucher and Schneider (1967) show the Triassic Alpine type lead-zinc deposits to be restricted to a few distinct beds of the thick limestone-dolomite sequence; the Ladinian deposits are bound to reef complexes, occurring mainly in a few layers in the back-reef facies, with a few iron (zinc) bodies in the coral reef itself.

In northwest Africa, Monseur and Pel (1973) list the following Jurassic stratiform lead-zinc deposits that occur in reefs, many also extending into the back-reef facies: Bou Dahar (Pb), Bou Arhous (Pb), Tagount (Pb), Touissit (Zn, Pb) in Morocco; El Abed (Zn, Pb), Dominique Luciani (Zn, Pb), and Deglen (Zn, Pb) in Algeria. In addition there are a number of non- stratified deposits in reef complexes, and several deposits containing manganese and/or iron, such as Tiharatine, Youdi, and Bou Arfa in Morocco. Lower Cretaceous deposits include Quenza and Bou Khadra (Fe), and Mesloula (Pb) in Algeria, and Reocin (Zn, Fe, Pb) in Spain, as well as a number of non-stratified deposits.

The Ruby Creek deposit in Alaska is a stratiform copper deposit in interbedded dolomite, limestone, and phyllite that form part of a Middle Devonian reef complex, containing stromatoporoids, corals, brachiopods, and gastropods (Runnels, 1969). In the Jurassic Todilto Limestone, Grants, New Mexico (Perry, 1963), uranium is present within and especially on the edges of elongate to arcuate reefs. Perry considers that the reefs are of algal origin, though they are coarsely crystalline and no internal structure is preserved. Fluorite pipes occur in reef complex limestone of the Carboniferous limestones of Derbyshire (Ford, 1969).

DISCUSSION

The mineral deposits associated with stromatolites fall into the broad class of stratiform deposits in sedimentary rocks, since they are with few exceptions either stratiform in habit, or are portions of stratiform deposits. They

658 F. MENDELSOHN

occur in biohermal or biostromal stromatolitic rocks, or in closely associated rocks of both the fore-reef and back-reef environments, so that the hosts are dominantly carbonates, dolomite and limestone, or argillites , but also range to arenaceous and even coarser sediments in places; interbedded volcanic rocks are common, but intrusive magmatic rocks are conspicuously absent.

These deposits occur in rocks of Proterozoic or Early Phanerozoic age, encompassing a 2-b.y. period (2.6-0.5 b.y) during which stromatolites reached their greatest development (e.g. Glaessner, 1968). In rocks of post- Early Ordovician age, broadly similar deposits are associated with bioherms built by corals, stromatoporoids, rudistids, etc., rather than the blue-green algae that were the main builders of the stromatolites. It is not clear to what extent algae and bacteria were concerned with the building of the Phanero- ozoic bioherms; perhaps they were important but were not preserved, or their role was originally unimportant because they were destroyed by the action of grazing and burrowing animals, as described by Garret (1970b).

There have been several references to the temporal partition of elements in the broad class of stratiform mineral deposits in sedimentary (volcanic) rock sequences, such as those by King (1965), Mendelsohn (1967), and Pereira (1967). For example, King describes a broad depositional pattern in which iron concentrations appear first, then copper, followed by lead-zinc and phosphate. Turning to the biohermal environment, Monseur and Pel (1973) find that copper is frequent in Precambrian and Infracambrian reefs, whereas zinc and to a lesser degree lead prevail in the Paleozoic and Mesozoic reefs, and that the associated rock also changes gradually from dolomite to limestone over this period. A wide range of valuable elements is concentrated in the deposits associated with stromatolites, and though the data are not sufficient to provide a complete picture, there is a suggestion that the different elements tend to be confined to host rocks laid down in particular periods. The major Superior-type iron deposits of the world were laid down during the Early Proterozoic (Goldich, 1973), and it is therefore not surprising that the iron deposits associated with stromatolites are confined to host rocks of this age. The Witwatersrand gold-uranium hosts are also Early Proterozoic, slightly older than the iron deposits, as are other gold-uranium-pyrite conglomerates. The first major copper and lead-zinc deposits associated with stromatolites appear at 1,700 m.y., though there seems to be a concentration of cupriferous deposits between about 1,200 and 700 m.y.; lead and zinc hosts are confined mainly to the period from 1,000 to 500 m.y., though they again become important in Late Paleozoic and Mesozoic biohermal deposits. Phosphorus is concentrated in mid- Riphean rocks, laid down between 1,350 and 950 m.y. Manganese occurs in Early Proterozoic hosts, then again in the modem oceanic nodules.

The genesis of stratiform deposits in sedimentary rocks has been the subject of considerable debate for many years, and there has been a substantial swing away from earlier concepts of introduction of elements from late


intrusive bodies, to those concerned with syngenetic and diagenetic processes. The genesis of mineral deposits associated with stromatolites must be considered as a part of this broad picture, though they have special aspects and problems, some of which are considered below.

Banerjee (1971b) concludes that the Udaipur phosphorite was directly related to the action of stromatolitic algae, which led to the precipitation of phosphate and the trapping of phosphatic sediment in a shallow basin. Trudinger and Mendelsohn (Ch. 11.2) discuss the precipitation of phosphorus in shelf areas by chemical activity, or its accumulation by phytoplankton, though there is no direct evidence of phosphorus concentration related to stromatolite formation.

There is little doubt that the gold and uranium in the Witwatersrand formations were deposited and concentrated during the deposition of the host rocks (Pretorius, 1974). The organisms that form the “algal” mats, whether they were algae or a lichen-like form, seem to have played an active part in collecting a substantial part of the gold and uranium, both physically and organically. In the case of the uranium (gold-lead-copper) deposits of northern Australia, the evidence is conflicting (Walpole et al., 1968), and it does not seem likely that stromatolite growth contributed directly to ore deposition, though the role of the bioherms as a host during later processes could have been important.

It is generally accepted that the Early Proterozoic iron-formations are chemical sediments, though there are many differences regarding the mechanisms and conditions of derivation, transport, deposition, and diagenesis of the iron, silica, and associated elements (James and Sims, 1973). In many basins the iron formation is closely associated with stromatolitic carbonate rocks, although they may not be exactly contemporaneous, and with bacteria and blue-green algae. The ability of microorganisms to precipitate iron is discussed by Trudinger and Mendelsohn (Ch. 11.2). Cloud (1973b) suggests that, in Early Proterozoic time, the interaction of oxygen-producing photosynthetic blue-green algae and ferrous iron led to the removal of excess free oxygen, the conversion of ferrous iron to ferric iron, and the precipitation of iron as ferric oxides or hydroxides. Cloud suggests that at this time the pH was generally too low for the formation of carbonates, though Button (1975) suggests that in the Transvaal, offshore bars of clastic limestone led to the formation of barred basins within which evaporation led to iron deposition. It seems that, while microorganisms were probably associated with iron deposition, and algal mats might occur, stromatolitic carbonate rocks are unlikely to be intimately associated with iron formations, though they are an integral part of the ore-forming cycle and some forms could tolerate a substantial amount of iron in the environment.

Manganese is known to accumulate in stratiform deposits during sedimentation (Park and McDiarmid, 1964), and the work of Doyen (1973) and Litherland and Malan (1973) suggests the possibility that stromatolites were

660 F. MENDELSOHN

at least able to grow under conditions where manganese was precipitating. Evidence suggesting the direct precipitation of manganese in modern oceanic nodules by stromatolite-forming microorganisms is given by Monty (1973), and Trudinger and Mendelsohn (Ch. 11.2) show the close relation that can exist between manganese deposition and the action of microorganisms. It seems likely that some manganese deposits formed during, and perhaps related to, the growth of stromatolites.

Copper, lead, zinc, and iron sulphides, and associated minerals such as other sulphides, barite, and fluorite, provide some of the more spectacular and important examples of mineral deposits associated with stromatolites, as well as with other biogenic rocks. For the Mississipi Valley lead-zinc and related deposits, and many of the Phanerozoic deposits related to biogenic rocks that have been listed, a widely accepted idea is that the valuable elements were introduced by circulating or connate brines that were heated and moved up-dip from the depths of the depositional basins during diagenesis. The porous reef rocks provided channel ways for the solutions, and contained organic material and sulphur that created a reducing environment and led to precipitation of these elements as sulphides (Campbell, 1966; Beales and Onasich, 1970, etc.). The evidence is persuasive, and there can be little doubt that this process did operate in a number of these deposits. For the Baffin Island lead-zinc deposit, Geldsetzer (1973) attributes syngenetic dolomitization and sulphide mineralization to seawater-derived brine that circulated through karst-brecciated algal carbonates.

In the Zambia-Zaire copper-cobalt province, Garlick (1965, 1972) has summarized many of the arguments for syngenesis, while Bartholomd (1974) makes a convincing case for the diagenetic introduction of copper and cobalt at Kamoto; Annels (1974) proposes a similar process for the Zambian Copperbelt, and Van Eden (1974) suggests a combination of the two at Mufulira. Clemmy (1974) suggests a combination of tidal-flat, marine, and lagoonal condtions of sedimentation for the Ore Formation at Nkana; the copper content is related to shoreline, facies, and tidal directions, and mineralization took place during sedimentation. The bioherms in this province, which could be expected to have had the greatest permeability during diagenesis, are barren of sulphide minerals, whereas formations that represent algal mats and other stromatolitic biostromes are mineralized, the copper content commonly being related to the argillaceous content. These features suggest that mineralization was contemporaneous with sedimentation for the ore-shale deposits, and that the metal was introduced into the host rock together with the clay minerals; the stromatolitic formations and the organisms that formed them played an important role in the sedimentation of the host rock. Diagenetic processes were active in creating or enriching some deposits in this province.

The partition of metals is evident in many of 'these deposits, and this is well illustrated at Mount Isa, where copper in biohermal rocks is closely


associated with the stratigraphically equivalent bedded lead- zinc deposits. Bennett (1965) considers that all sulphides there were formed by the precipitation of metals during sedimentation, the sulphur being biogenic, and that the copper bodies formed by local re-mobilization during folding.

Trudinger and Mendelsohn (Ch. 11.2) show that the deposition of metal sulphide minerals could take place during sedimentation as a result of biogenic sulphate reduction, or of evaporitic concentration and the subsequent formation of organic complexes. Rickard (1974) considers that biogenic synsedimentary copper sulphide ores are possible, provided an adequate metal source exists; potential environments for stratiform copper-sulphide ore formation are sediments of active oceanic ridges, partly connate subsurface water systems, and near-shore organic-rich sediments near a mineralized hinterland.

CONCLUSIONS

The process of mineralization of stratiform deposits associated with stromatolites is an integral part of the whole development of sedimentary basins during sedimentation and diagenesis, and it is perhaps significant that many of the deposits discussed are in intracratonic basins. This general statement can be extended to include deposits associated with other biogenic formations, and even to a large part of the class of stratiform mineral deposits in sedimentary (volcanic) rocks. The favoured sedimentary environment is the barred basin-lagoon- tidal flat-evaporite-sabkha group of ocean-edge environments, of which stromatolitic bioherms and biostromes and algal mats often form an integral part. The valuable elements could be supplied from the hinterland by erosion, from the ocean itself, or from volcanic emanations in the vicinity; depending on the rate and nature of the supply and the conditions of sedimentation, these elements would be widely dispersed in the sediments or concentrated locally. Where mineral deposits are formed, the contribution of the microorganisms could be indirect, helping to create a favourable physical and chemical environment, or more direct in trapping enriched material, helping to precipitate the valuable elements organically, or supplying other necessary elements such as sulphur. Altema- tively the organisms could merely survive in an environment where deposition of these elements is taking place, without playing any part in the process.

Soon after sedimentation and through the diagenetic period, mineralization would take place through the agency of connate brines that leach the valuable elements from their wide dispersion or from already-formed deposits, or perhaps through brines bringing them from a source external to the sediments; the brines would move, generally up-dip, till precipitation takes place in a favourable environment. In this process, the biogenic

662 F. MENDELSOHN

formations could serve by providing channel ways for the solutions, organic material to create a reducing environment, and sulphur in the form of sulphate or biogenically reduced sulphide to precipitate metal sulphides. Later processes could be superimposed on earlier ones, leading to further enrichment or leaching of existing deposits, and also providing conflicting evidence of ore genesis. Late diagenetic and metamorphic processes could be expected to modify deposits by recrystallization or even remobilization and redeposition, so that original textures, characteristics, and processes become obscured; this could be expected to be significant in many stromatolite- related deposits, which are among the earliest known stratiform deposits.

The reason for the temporal partition of elements in these deposits is not entirely clear, but it is in part related to the changing supply of particular elements at or near the surface of the Earth. An example of this is iron, which became abundantly available when the evolving conditions at the surface reached a particular stage in Early Proterozoic times, to form a group of major deposits that has not been duplicated since. Lead does not appear in any stratiform deposit before about 1,700 m.y., and indeed in few known major deposits formed before this time. Copper and zinc are common in Archaean volcanogenic deposits, and copper occurs in Early Proterozoic stratiform deposits, but both appear, with lead, for the first time in stromatolite-related deposits in the Middle Proterozoic at around 1,700 m.y. In addition to the changing supply of the elements to the environment in which stromatolites formed, perhaps there was an evolution of the microorganisms or changes in their ability to survive in or react with concentrations of these elements.


Chapter 11.2

BIOLOGICAL PROCESSES AND MINERAL DEPOSITION

Philip A . nudinger and Felix Mendelsohn

INTRODUCTION

Mendelsohn (Ch. 11.1) has discussed the frequent association between mineral deposits and stromatolites and one may enquire whether this association is fortuitous and whether stromatolite-building organisms play some role in mineralization.

There are two ways in which reefs and biostromes can affect sedimentation: the building of a reef or the creation of an algal mat affects the physical and chemical environment of sedimentation, and the organisms can participate directly in chemical processes, either actively or passively.

Reef building has the major effect of creating lagoons in back-reef areas or within patch reefs, which would result in quiescent, anaerobic conditions favourable for the accumulation and preservation of organic material and metals. It is also possible that a fringing reef could protect the fore-reef area from wind and wave action and thus create similar conditions there. A reef could also form a front between waters of different composition, where reaction could take place and metal become precipitated (Hofmann, 1973). The trapping of sediment by an algal mat is also a significant factor in sedimentation, and it has been shown at Mufulira, Zambia, that the amount of copper present in a stromatolite body is proportional to the amount of sediment present (Malan, 1964).

Along the Zambian Ore shale deposits, offshore from the zone of reefs and sandy sediments, a zone of shale about 2 km wide contains the copper orebodies, beyond which is a zone 5-10km wide that contains pyritic carbonaceous shale; beyond this again are (deeper-water) carbonates (Garlick, 1961, p. 157). These are the dimensions and types of formations (apart from the copper?) that could be expected in an interdeltaic coastal marine lagoonal environment (Hamblin, 1973). Van Eden (1974) concludes that broadly similar conditions of deposition, with off-shore shoals, existed during the deposition of the sands that now form the Mufulira ‘C’ orebody. It is perhaps significant that these major orebodies are associated with such an environment, in which conditions similar to those described for back-reef lagoons existed over wide areas.

664

TABLE I

P.A. TRUDINGER AND F. MENDELSOHN

Examples of element concentration by marine plants (from Bowen, 1966)

Element Concentration in Concentration in Enrichment factor sea water marine plants ( P P d (PPm dry wt)

Ca 400 10,00CL300,000 2 5-7 50 c u 0.003 11 3,700 Fe 0.01 7 00 70,000 Mn 0.002 53 26,500 Ni 0.005 3 600 P 0.07 3,500 50,000 Pb 0.00003 8 267,000 Zn 0.01 150 15,000

Direct contributions to mineralization by organisms may take place by concentration of mineral-forming elements within living cells or by active metabolic transformation of elements. The ability of marine plants (principally algae) to accumulate elements is illustrated in Table I. It is apparent that, following death and sedimentation of cellular material, these organisms are potentially capable of inducing considerable enrichments of a number of elements in sediments. The proportion of cellular material actually reaching the sediment surface, however, will depend on the rates of sedimentation versus those of cell degradation. In some instances, the proportion may be quite low: Deuser (1971), for example, calculated that only about 4% of primary organic carbon is incorporated into the sediments of the Black Sea. On the other hand, in shallow-water environments where growth of microflora may take place primarily at the sediment- water interface, it might be reasonable to conclude that a high proportion of intracellular metals and other elements would be available for mineralization at the sediment surface. As well as “active” accumulation of elements, “passive” formation of metallo-organic complexes has a profound influence on the mobilization, transport and fixation of metals (Berge, 1950; Saxby, 1969). The frequently observed positive relationship between organic content and metal enrichment in sediments (Krauskopf, 1955) is probably the result of active and passive biological processes.

Metabolic reactions leading to the fixation of mineral-forming elements have been discussed by Silverman and Ehrlich (1964). The reactions are primarily oxidations or reductions and fall into two groups: (1) those which cause direct precipitation, e.g. the oxidation of soluble ferrous and manganous ions to ferric and manganic oxides or hydroxides; and (2) those in which the products react with other elements to form precipitates, e.g. metabolic production of the metal precipitant, hydrogen sulphide.

In this chapter examples of both direct and indirect contributions of


biological activity are discussed with reference to the formation of minerals of the types associated with stromatolites.

PHOSPHORITES

Large-scale deposition of phosphate minerals frequently occurs in regions where deep, cold ocean waters containing up to 0.3 ppm PO$- upwell along a shelf and merge with warmer surface waters which have an average PO:- concentration in the order of 0.01 ppm (McKelvey, 1967; see discussion by Cook, 1975). Phosphate precipitation has been attributed to purely chemical factors such as loss of C02 and increases in pH resulting from warming of the phosphate-rich waters (Kazakov, 1973). On the other hand, due to the increased availability of nutrients, upwelling regions are also zones of intense planktonic productivity (Ryther, 1963) which may be as much as 4 g C rn-' day-' (Steemann-Nielsen and Jensen, 1957). This has given rise to the hypothesis that organisms, particularly phytoplankton, may be instrumental in accumulating phosphorus which, after deposition and degradation of the organisms, is then released and fixed in surface sediments (see discussions by Bushinskii, 1966; Gulbrandsen, 1969). Based on a phytoplankton productivity of about 1-1.5gC m-2 day-', Gulbrandsen (1969) and Cressman and Swanson (1964) concluded that large phosphorite deposits, such as those of the Permian Phosphoria Formation, could have been derived from planktonic sources in the relatively short time of 2 l o 5 years. A diagrammatic representation of the upwelling-biochemical hypothesis for phosphorite formation is shown in Fig. 1.

Increasing temperature, pH salmlty S H O R E

High organic

Organic degradation and deposition

Apatite precipitation

Warm bottomwater high pH high saltnity

O C E A N

Cold water low pH High phosphorus ' /

Fig. 1 . Diagrammatic representation of the biochemical-upwelling hypothesis for phosphorite deposition (after Gulbrandsen, 1969).

666 P.A. TRUDINGER AND F. MENDELSOHN

As far as the authors are aware there are no known instances of modern phosphoritic stromatolites. Nevertheless, with some reservations, the arguments advanced for biochemical deposition of phosphate could be equally applicable to stromatolite-building organisms. The reservations are introduced by lack of agreement of some aspects of the mechanism of phosphorite formation and the conditions under which it takes place. Proponents of “biochemical accumulation’’ hold that shallow-water conditions are necessary to allow for rapid deposition of dead organic matter and a minimum of degradation, which could lead to premature release of phosphorus, during sedimentation (Bushinskii, 1966). Gulbrandsen (1969) proposed that the environment must also be highly aerobic (see also Dietz et al., 1942; Berge, 1972) to facilitate rapid decomposition of sedimented organic matter. These conditions are similar to those under which calcareous stromatolites are formed (Bathurst, 1971). On the other hand uranium in sea-floor phosphorites is frequently in the reduced tetravalent state (Altschuler et al., 1958), which, together with the frequent association of bedded phosphates with organic matter, pyrite and black shales, points to an essentially reducing environment for phosphate deposition (Youssef, 1965). A requirement for lo’w oxygen tension in phosphorite formation had been suggested earlier by Blackwelder (1916) and by Mansfield (1927).

A second point of contention is whether phosphorites are composed of primary or secondary minerals. Ames (1959), Youssef (1965) and Pevear (1966, 1967), for example, suggest that at least some phosphorites are formed by secondary phosphatization of carbonate minerals rather than by direct precipitation. In such situations a relationship between stromatolite- building organisms and phosphate mineralization could be coincidental, or it might reflect the permeability of many stromatolites.

MANGANESE NODULES

Monty (1973) concluded that manganese nodules from the Blake Plateau are oceanic stromatolites and several others have alluded to the association of organisms with manganese nodules (Crerar et al., 1972; Crerar and Barnes, 1974). Graham (1959) detected organic matter in nodules from the Blake Plateau and suggested that organisms were important in nodule formation (Graham and Cooper, 1959). Greenslate (1974) described structures resembling tubes of Saccorhiza (foraminifer) in nodules from the Pacific Ocean and also reported that Saccorhiza and other, unspecified, organisms on nodule surfaces build structures containing large amounts of manganese and iron compounds. Oborn (1964) reported the presence of abundant Micrococcus (bacterium) in a manganese nodule from deep Pacific Ocean water but, as there was a lapse of nine years’between collection and microbiological examination of the nodule, this finding must be viewed with


reserve. More reliable are the results of Ehrlich (1972 and references quoted therein) who isolated both manganese-oxidizing and -reducing organisms from Atlantic Ocean nodules.

The ability to promote manganese oxidation appears to be quite widespread amongst bacteria, fungi and algae (see, for example, Silverman and Ehrlich, 1964; Alexander, 1965; Schweisfurth, 1970). Deposition of manganic oxides from fresh waters has been particularly associated with chlamydobacteria (Zapffe, 1931; Wolfe, 1960), Hyphomicrobium (Tyler, 1970) and Metallogenium (Perfilev and Gabe, 1965; Sapozhnikov, 1970). These bacteria are characterised by the presence of sheaths, stalks or filaments around which metal oxides are deposited, and this property could possibly contribute to the development of structure within the oxide precipitates. Filamentous and stalked bacteria are frequently found in ferromanganese nodules in lake sediments (Sapozhnikov, 1970) and in metalliferous deposits in filters (Wolfe, 1960) and pipelines (Tyler and Marshall, 1967a). In the latter instances thick deposits containing 1-50%Mn and 3-25%Fe may be derived from waters containing trace amounts of metals, and the Mn:Fe ratio of the deposit may be the reverse of that in the waters (see Trudinger, 1975).

Tyler and Marshall (196713) showed experimentally that biological factors are necessary at least for the initiation of manganese deposition in pipelines. The means by which microorganisms promote oxidation of manganous ions is, however, not known for certain. Three main mechanisms have been proposed each of which might operate in particular circumstances. Sohngen (1914) concluded that oxidation is catalysed by hydroxy-acids such as malate and citrate and suggested that microorganisms contribute by creating an alkaline environment (and, presumably, by producing hydroxy-acids) which favours the chemical reaction. Mose and Brantner (1966) and Brantner (1970) proposed that manganese precipitation is due to microbial metabolism of the organic moieties of soluble organo-manganese complexes. A similar proposal was advanced by Sorokin (1972) for the formation of marine manganese nodules.

There is evidence, however, that some organisms possess specific manganese-oxidizing enzymes. Johnson and Stokes (1966), for example, reported that Sphaerotilus discophora developed the ability to oxidize manganous ions only when grown in the presence of manganese and Ehrlich (1968) prepared a protein fraction, which catalysed manganese oxidation, from an Arthrobacter isolated from a marine manganese nodule. Ehrlich (1963, 1966) noted that preformed manganese oxide, or certain other solid materials, were necessary for manganese oxidation by Arthrobacter. He concluded (Ehrlich, 1968, 1972) that manganous ions were initially absorbed on preformed manganese oxide and subsequently oxidized by bacteria with the creation of new absorption sites:


chemical H2Mn03 + Mn2+ -MnMn03 + 2H+

biological MnMn03 + 2H20 + 402 (H2Mn03)2

Whatever the precise oxidation mechanism may be, the expected slow and periodic growth of microorganisms and localized variations in microfloral communities might account for the observed characteristics of nodule growth and for differences in shape, structure and composition of manganese nodules from different locations (Greenslate, 1974).

IRON MINERALS

Species of chlamydobacteria, Hyphomicrobium and Metallogenium (loc. cit.) are associated with the precipitation of ferric iron at near neutrality. This property is shared with Gallionella spp. which are kidney- shaped bacteria characterized by the presence of long, slender, twisted ribbons or stalks (Choldony, 1924). The latter structures appear to be composed largely of oxidized iron compounds bound to an organic matrix. The general properties of the neutral iron bacteria have been discussed by Pringsheim (1949a, b). Acidophilic bacteria of the genera Thiobacillus and Ferrobacillus oxidize ferrous iron enzymically for the production of energy but because of their highly acidic habitats they are probably not relevant to the present discussion.

The mechanism of bacterial iron precipitation by the neutral iron bacteria is unknown: there is at present no evidence for enzymic catalysis of ferrous oxidation. Ellis (1907) and Pringsheim (1949b) have suggested that the bacteria may possess specific binding sites for preformed ferric hydroxides while others (e.g., Aristovskaya, 1961; Clark et al., 1967) propased that ferric-organic complexes are absorbed into the bacterial cell and that iron is fixed in capsular material after metabolism of the organic moieties.

The iron bacteria are widely distributed in ferruginous waters and they precipitate, ferric hydroxides from solutions containing less than 1 ppm of iron (Starkey, 1945). They are frequent inhabitants of environments favouring the formation of bog-iron ore (Harder, 1919). Kuznetsov et al. (1963) suggest that microorganisms play a dual role in bog-iron ore formation: ferric iron is reduced by the microflora in subsoils to the mobile ferrous form which migrates towards the oxidizing zone and becomes reoxidized by iron bacteria. Puchelt et al. (1973) observed Gallionella ferruginea in ferric hydroxide sediments presently forming in bays of the Kameni Islands of the Aegean Sea. They concluded (Puchelt et al., 1973, p. 767) that the bacteria “occur in such masses, that there is no doubt concerning the presence, activity and share of these bacteria in the process of iron deposition”. Periodic mass development of iron bacteria in Lake


Krasnoe of the Isthmus of Karelia is accompanied by a decrease in both ferrous and total soluble iron again indicating a role for these organisms in iron oxidation and deposition (Drabkova and Stavinskaya, 1969).

Algae have also been implicated in iron deposition. Pringsheim (1946) noted incrustations of iron oxides on the surface of algal cells and Choldony (1926) suggested that algae participate indirectly by creating an oxidizing environment which facilitates oxidation of ferrous iron. A similar suggestion was advanced by Cloud, 1968a, b, 1973b) who equated the deposition of banded iron formations prior to 2 * l o 9 years B.P. with the development of blue-green algae (see also L.Y. Kodyush, 1969, cited by Alexandrov, 1973). Lepp and Goldich (1964) believe that changes in pH, Eh and C02 concentration brought about by algae and bacteria may have been important factors in the genesis of Precambrian iron formations.

The possible contributions of iron bacteria to the formation of massive iron formations has been discussed theoretically by Harder (1919) and by Zapffe (1933) and, more recently, structures resembling modern iron bacteria have been detected in stromatolitic cherts from the Gunflint Iron Formation (Barghoorn and Tyler, 1965). Cherts from a number of iron formations also contain alga-like structures (Gruner, 1922; Tyler and Barghoorn, 1954; Cloud and Hagen, 1965; Cloud and Licari, 1968; LaBerge, 1967, 1973), and on the basis of such fossil evidence it has been suggested that microorganisms were possibly instrumental in precipitating the original constituents of iron formations (Moorhouse and Beales, 1962; LaBerge, 1973). In this respect the similarities noted by Walter (197213) between siliceous stromatolites of the Lake Superior iron formations and those formed by algal and bacterial action in the hot-spring and geyser effluents of Yellowstone National Park, Wyoming (Walter et al., 1972), take on added significance.

SULPHIDE MINERALS

The formation of metal sulphides as the result of bacterial reduction of sulphate is now welldocumented (Miller, 1950; Baas Becking and Moore, 1961; Rickard, 1969; Hallberg, 1972). The responsible organisms are the so- called dissimilatory sulphate-reducing bacteria belonging to the genera Desulfovibrio and Desulfotomaculum. These organisms are strict anaerobes which oxidize organic matter at the expense of sulphate reduction (Postgate, 1959, 1965; Trudinger, 1969), and the overall synthesis of metal sulphides (MeS) may be depicted as follows:

organic + SO:- - organic + H2S -+ MeS matter matter

Me reduced oxidized C


Rates of biogenic sulphide production in modem euxinic environments appear to be adequate to account for the formation of a number of stratiform mineralized deposits (Temple, 1964; Trudinger et al., 1972; Rickard, 1973; Goldhaber and Kaplan, 1974). Some support for the concept of biogenic sulphide deposits is the fact that patterns of sulphur isotopic composition in such deposits often resemble those predicted from the known isotopic fractionation effects during bacterial sulphate reduction (e.g., Wedepohl, 1971; Sangster, 1968, 1971b; Schwarcz and Burnie, 1973). Recently, microfossils of bacterial dimensions have been detected in mineralized zones of the McArthur River lead-zinc-silver deposit, Northern Territory, Australia (Hamilton and Muir, 1974).

The major factor controlling the extent of bacterial sulphate reduction 5n natural environments is the availability of organic matter (Sorokin, 1962; Ivanov, 1968; Berner, 1970; Chukrov, 1970) derived from decaying algae and other flora and fauna. It is perhaps not surprising therefore, that mineralized zones of copper, lead, zinc and iron sulphides are frequently associated with stromatolites (Mendelsohn, Ch. 11.1).

Renfro (1974) considered that the depositional environments of Kupferschiefer, Roan and certain other evaporite-associated stratiform metalliferous deposits may be likened to evaporite flats known as coastal sabkhas. He proposed that, during periods of regression, oxidized terrestrial- formation waters a t low pH mobilize metals from underlying rocks and pass upward through H,Scharged layers of decomposing algal mats where metals are immobilized as sulphides. Jackson and Beales (1967) in a discussion on the genesis of Mississippi Valley-type ores, proposed that in the presence of organic matter such as oil or bituminous sediment, gypsum and anhydrite in, or adjacent to, carbonate horizons, are reduced by sulphate-reducing bacteria to hydrogen sulphide which reacts with lead and zinc transported as metal- chloride complexes from adjacent basinal sediments.

An alternative role for algae in sulphide mineralization was proposed by Roberts (1973) in relation to the genesis of the Woodcutters lead-zinc prospects in the Northern Territory, Australia. The host rocks of this prospect are highly dolomitic and contain abundant stromatolitic structures (W.M.B. Roberts, 1974, personal communication). Roberts concluded that the initial stages of the formation of the Woodcutters deposit took place in a shallow restricted basin where dense algal growth and evaporitic conditions caused co-precipitation of lead and zinc with precursors of the host-rock dolomite. During subsequent dolomitization the metals, in the form of complexes with products of algal decomposition, were released into pore solutions ih ich then migrated to the sites of metal-sulphide deposition.

Roberts (1973) also concluded that the concentration of lead and zinc in the basin waters from which the Woodcutters deposit was derived, would need to have been in the order of 0.5-lppm. Heavy metals are often considered to be extremely inhibitory to microorganisms but there are


several reports of reasonably high tolerance amongst algae (Table 11). In many instances the precise levels of tolerance listed in Table I1 are probably overestimates because of complications introduced by partial precipitation of metals by anions, such as phosphate, in the growth media. Nevertheless, it does not seem unreasonable that vigorous algal growth could take place in the presence of the metal concentrations required by Roberts’ (1973) model, particularly since tolerance to metals may be increased by such factors as the presence of organic matter, ion interactions and adaptation (Sadler and Trudinger, 1967).

TABLE I1

Toxicity of metals to algae

Alga Level of tolerance Reference (PPm) Pb c u Zn

Chlorella vulgarus 110 1 400 Den Dooren de Jong (1965)

Scenedesmus quadricauda - 0.5-6* - Khobotev et al. (1969)

Variety of fresh-water 0.1-5** 0.1-1** 3-60** Whitton (1970)

Variety of fresh-water - >25*** - Mendelsohn (1973)

(Chlorophyta)

(Chlorophyta)

Chlorophyta

Cyanophyta

*Cu-organic complexes used.

***Growth at pH 1 1 : major part of Cu in ionic form as determined polarographically Geometric mean of “just non-inhibitory ” and “just-lethal” concentrations **

CONCLUDING REMARKS

The association of stromatolites and other fossil structures with mineral deposits has often been taken as evidence for a possible role of biological activity in mineral formation. In the presentday environment there are several examples of biological processes which contribute, directly or indirectly, to the fixation of mineral elements and there is evidence that analogous processes may have operated during the genesis of some mineral deposits in the past. Should there have been great “blooms” of algae and bacteria, as seems likely in the presence of sufficient nutrients and the probably higher carbondioxide content in the atmosphere in Precambrian times (Rutten, 1966), the effect on the accumulation of metals could have been substantial.

I t is considered that, given an adequate supply of metals or other valuable elements and the appropriate conditions, stromatolite formation during


sedimentation could be accompanied by the accumulation of these elements to form ore deposits.

ACKNOWLEDGEMENTS

The Baas Becking Geobiological Laboratory is supported by the Bureau of Mineral Resources, Commonwealth Scientific and Industrial Research Organization and the Australian Mineral Industries Research Association Ltd.

12. CONTRIBUTIONS TO THE HISTORY OF THE EARTH-MOON SYSTEM

Chapter 12.1

GEOPHYSICAL INFERENCES FROM STROMATOLITE LAMINATION

Giorgio Pannella

INTRODUCTION

In the preceding pages different aspects and uses of stromatolites have been emphasized. Potentially, if not practically at the moment, the use of stromatolites as recorders of geophysical and astronomical conditions of the Earth during its evolution could be even more important. Since growth of stromato!ites is a direct outcome of environmental dynamics, it has been suggested that growth structures could be used to check geological changes in the speed of the Earth’s rotation (Runcom, 1966; McGugan, 1967; Pannella et al., 1968; Pannella, 1972a) and to determine the amplitude of paleotides and polar shifts (Cloud, 1968a; Nordeng, 1963; Vologdin, 1963, 1964; Cloud, 1968a). Fossil organic structures which may provide the first type of information have been called paleontological clocks (Runcorn, 1966). In order to separate the effects of changes in the Earth’s moment of inertia from those due to tidal dissipation, a paleontological clock must provide, from continuous sequences, two sets of data: one on the length of the sidereal day, the other on the length of the synodic month (Runcom, 1964).

The crucial and controversial question with regard to the use of stromatolites as clocks is with what precision and continuity stromatolites record, during growth, the astronomically driven environmental fluctuations and how much of this record is left for us to decipher. For several reasons, which will be mentioned in the following pages, stromatolites, in general, are not as reliable paleontological clocks as other fossils. Hofmann (1973) attributes to them either zero or a very low degree of reliability. In this paper it is not intended to take an opposite and over-optimistic view of these applications because stromatolitic growth records in most instances are so incomplete as to justify Hofmann’s conclusion. However, in exceptional instances and with the help of statistical techniques, stromatdites are going to provide precious insights into the largely unknown Precambrian. The strongest reservation to these applications of stromatolites depends on the danger of extrapolating from one localized situation to world-wide generalizations. Another reservation is the uncertainty in respect to determining the absolute age of Precambrian material.

674 G. PANNELLA

Notwithstanding the many difficulties and the rather tenuous results of the first naive attempts, there is ample room for improvement and, perhaps, the ultimate goal of tracing in broad lines the Precambrian evolution of the Earth-Moon system, is not too far in the future. The geologist and the paleontologist with the help of stromatolites will undoubtedly contribute.

STROMATOLITES AS RECORDERS OF PERIODICAL ENVIRONMENTAL CHANGES

The modem analogs of stromatolites provide insight and models of their growth. Light stimulation, by controlling metabolism, plays an essential role in the growth of stromatolites. Short-term experiments, extending from a few days to two weeks, have shown that laminae of many modem stromatolites form with a daily frequency because of light stimulation of algae or bacteria. Monty’s experiments, in which Bahamian algal oncolites formed laminae consisting of organic-rich (deposited during the day) and inorganic- rich lamellae (deposited at night), suggest daily growth is a basic feature of some stromatolites, but add little to the question of the accuracy and continuity of stromatolitic sequences (Monty, 1965b, 1967). Monty noted that there is size limitation of growth for the oncolites which are carried away after about two months growth.

Gebelein (1969) described daily layers of about 1 mm thick from subtidal algal oncolites from Bermuda and attributed them to the day-night algal growth cycle. He also emphasized the role of currents, and of velocity and rate of sediment movement in the building of stromatolitic structures. Currents over 20 cm/sec do not permit the formation of algal mats, and only bottom sediment movement of less than 60-80 g/h per foot allows mats and domes to accrete. The thickness of lamination is inversely proportional to current velocity. Tidal fluctuations, which control current velocity and direction, can thus leave a record in sequences of laminae. In Bermuda, for example, more sediments are transported at flood tides than at ebb tides (Gebelein, 1969). In intertidal algal mats high-tide flooding creates sediment- rich lamellae, the thickness of which is proportional to the period of flooding, whereas organic-rich lamellae are related to low tides (Gebelein and Hoffman, 1968). The relationships between tidal fluctuations and stromatolitic deposition in the intertidal zone could produce, in areas of semidaily tides, two laminae per day. The possibility that this could reduce the accuracy of stromatolites by producing more laminae than days is not relevant here since intertidal stromatolites, limited and discontinuous in their growth, are unlikely candidates for geochronometry. But the tidal mechanism of stromatolitic growth must be kept in mind when interpreting sequences of laminae. Whereas flooding and air exposure are the controlling parameters of

HISTORY OF THE EARTH-MOON SYSTEM

TABLE I

675

Hypothetical frequencies of periodical growth sequences in modem stromatolites

Growth-pattern frequencies Remarks

Tidal bands/year Daily increments/tidal band

2 variable 1- 3 (only exceptionally laminae are rarely daily; Y 0 higher) discontinuous sequences ; U algal mats are common

forms Extreme high water of spring tides I 78 - - - - - - - - - m - - < 5- - - - - - - - - -

< 12 a

3-7 the frequencies are highly variable depending on tidal types, diurnal inequality

d

< 20 4-10 columnar stromatolites and algal mats are common forms

laminae in most instances are deposited with daily Extreme low water of spring tides a - - - - - - - - - - - - - - - - - - - - -

; 2 <25 z 1 rn

< 14 frequency ; continuous sequences may form; large dome and digitate stromatolites are the most characteristic forms

lamina formation in the intertidal zone, subtidal stromatolites, unfortunately still poorly known, are affected by both light-dark cycle and tidal effects on sediment supply and are the only promising paleontological clocks. In addition the relatively deep-water ones are less subject to erosion.

Other works show that not all laminations are daily. Supratidal algal mats grow only during floodings, the frequency of which depends on many factors: tidal amplitude, height of the algal mats above mean low water, and storms (Hardie and Ginsburg, 1971). Generally the higher is the position of the mats the more discontinuous is the record. Theoretically, and apart from any post-depositional disruptive action, the number of laminae per tidal cycle (Pannella, 1972b) is a direct function of tidal floodings, which, in turn, are a function, at any given point, of tidal amplitude and tidal type. Table I represents the theoretical frequency ranges of periodical growth patterns, based on these assumptions: (1) a lamina records a 24-h dark-light cycle, or an episode of tidal flooding (which can occur with frequencies ranging from 12.42 h up to six lunar months), or a mixture of both tidal and solar cycles; (2) deposition occurs when algal mats are covered with water, when there is a proper sediment supply (in allochemical growth), and during metabolic

G . PANNELLA 676

activity of algae or bacteria (in allochemical and orthochemical growth). The Table has obvious limitations because it applies only to instances of a constant accretion, which are extremely rare in nature, and does not include the very common random noise and disruptions, but it provides consistency to the apparently contradictory and spotty data on lamina formation. One of the weaknesses of any generalized model of stromatolitic growth lies in the fact that, in the shallow-water marine environment the same parameters involved in growth patterns may also be the cause of their destruction. This fact, together with the tendency, in the intertidal zone, to form incomplete sequences, led to the conclusion-at the time when stromatclites were thought to be only intertidal - of the unsuitability of stromatolites as paleontological clocks.

From Table I it is clear that ideal growth sequences can only form in a subtidal environment, where the light-dark cycle is still effective and where physical disruption is less likely. In the modem situation, however, even a subtidal environment cannot produce ideal sequences because of biological disruption. One has to look for a special habitat where metazoan life is impossible. Hence the importance of data on growth patterns of hot-spring stromatolites and stromatolite-like deposits. From the experimental data obtained by marking stromatolites in the Yellowstone hot springs it appears that algal and bacterial stromatolites there grow by daily addition of laminae as in the Bermudan stromatolites. The growth record, however, does not appear to be continuous or similar in all columns even in the same spring. The change of flow patterns gives to each column an individual flood and, thus, growth history. In conical forms the laminations appear to be “crudely diurnal” (Walter et al., Ch. 6.2). Laminations in nonbiogenic “stromatolites” (geyserites) have periodicities that appear to be related more to geyser activity than to diurnal environmental variations. Their number is highly variable as shown by Walter’s figures in this book (p. 106). It will be interesting to confront these figures with continuously monitored variations in the pools and to compare lamination periodicities in stratiform geyserites in the same pools.

All these works on modem stromatolites demonstrate that while diurnal laminae are common the continuity of the growth record is restricted to short time intervals. For the formation of ideal sequences an environment must meet the following conditions: (1) be aqueous, quiet, and with some physicochemical parameters oscillating regularly and periodically with daily fixed frequencies; (2) be affected by tides but not subjected to strong storm waves; (3) be populated by algae or bacteria with diurnal growth and which cause precipitation of a preservable layer; (4) be lacking in other biological systems which cause disruptions. Such an environment containing stromato- Iites, with perhaps the exception of hot-spring pools, is not found in today’s world. In the Precambrian, however, this favourable environment may have existed.

HISTORY OF THE EARTH-MOON SYSTEM

FOSSIL STROMATOLITES

677

Considering that we are living in unusual geologic-climatic conditions, extrapolations from modem analogs to Phanerozoic stromatolites are danger- ous. From modem analogs to their Precambrian ancestors, however, there is a “Quantum jump”, because of the time involved and of the biological and geological evolution of our planet. The Precambrian geologic record with its impressive sequences of rhythmites and stromatolites creates the inevitable feeling that in pre-metazoan time environmental rhythms were written undisturbed in the rocks, that the Precambrian is the ideal locus of rhythmometric studies. Strangely enough, no in-depth study like that of Sander (1951) exists for the Precambrian.

Paleontological clocks

A survey of many fossil stromatolites has led to the conclusion that most of them, in fact all of those studied except one, could not provide reliable data for reconstructing the length of the day and of the synodic month in the Precambrian (Pannella, 1972a, and unpublished data). Sequences are incomplete and show unmistakable signs of interruptions and erosion (Fig. 1). In their general morphology most of the studied stromatolites are similar to modern Shark Bay stromatolites. Pannella (1972a) has pointed out, and further observations have confirmed, that small, digitate columnar stromatolites show generally longer sequences and a higher number of laminae per periodical band than large columnar ones. Much evidence strongly suggests that they were formed in subtidal environments and they are similar to others reported as subtidal (Playford and Cockbain, 1969; MacGregor et al., 1974).

Very small digitate types are present in stromatolitic beds of the 2,000 m.y. old Biwabik and Gunflint formations. One type, found on the western flank of Mink Mountain in the Thunder Bay area in Ontario, has yielded the most promising sequences of laminae, and if further study confirms the finding, is going to provide the first clues on the length of the synodic month and the year of Huronian times (Pannella, in prep.)

Using morphological criteria Walter (1972b) has suggested that some Mink stromatolites are nonbiogenic geyserite-like deposits very close to those found in Yellowstone. This hypothesis does not contradict my results nor change the interpretation of the periodicities in the lamination sequences. Harmonic analysis of the lamination spacing indicates strong tidal control of the lamina deposition. The possibility of a tidal effect on geyser activities is mentioned by Walter (Ch. 3.3 herein). Earth tides are more strongly felt nearshore because of the direct contribution of marine tides (Melchior, 1966). In the nearshore setting where the Gunflint stromatolites grew tides must have represented the strongest periodical force.

678 G. PANNELLA

Fig. 1. Polished surface of Great Slave Lake stromatolites showing distinct seasonal changes in color and composition of growth bands. Light bands contain thicker laminae than dark bands and were deposited during periods of fast growth (possibly warm and/or dry season). Dark bands show high organic content and are preferentially deposited over micro-unconformities indicating erosion (due, perhaps, to “equinoctial” storms) of the light bands. The filling between columns contains large fragments of light-band laminae further supporting the idea that the energy in the environment increased after the fast- growing periods. The numerous interruptions in the sequences make this type of stromatolite (similar to the modem ones found in Shark Bay) unsuitable for geochronometry. w = slow-growth bands; s = fast-growth bands; d = major discontinuities in growth sequences. Scale 1 cm.

While the analytical procedures and the detailed discussion of the sequences will be the object of another work, some results must be mentioned here for they justify the optimism in the use of stromatolites for deriving geophysical information. Preliminary results of the study of the Mink stromatolites and of the Biwabik Formation which contains stromatolites simjlar in morphology but quite different in microstructure have been published elsewhere (Pannella, 1972% b). The method used to obtain the periodicity of recurrent growth bands and zones consisted simply of counting laminae between repetitive patterns. There are, at least, two objections to this method: it relies only on the observer to solve ambiguities of the record

HISTORY OF THE EARTH-MOON SYSTEM 679

S I L I C A nicn

Fig. 2. a. Tidal amplitude modulations at Barnstaple Harbor, Massachusetts, from September 1965 to April 1967. b. Curve obtained by using running mean of lamina thicknesses in Mink Mountain stromatolites (Gun flint Iron Formation). The dashed line in A is related to the lunar changes in declination, in B is only a hypothetical reconstruction since the noisy record does not allow a resolution of the ambiguities. The dotted area coincides with the silica-rich, fast growing band. The use of several curves of this type and coherence tests will possibly resolve the ambiguities. The important fact remains that the two curves are comparable, supporting the conclusion that, besides changes in periodicities, the Earth-Moon system has not changed much in the last 2,000 million years.

and it erases many useful details recorded in the sequences. The method of measuring the thickness of the laminae, though by no means foolproof, is better in that it provides quantitative figures that can be subjected to mathe- matical analysis.

Both methods were applied to the study of Mink stromatolites. But, while the former supplied values for the frequency of what have been interpreted as monthly and annual bands, of 39 laminae (the highest count) and 448 laminae, respectively, the latter allows the harmonic analysis of the variations in spacing, providing a more precise way to evaluate the frequency of the bands and other additional data that may help the interpretation of their meaning.

Fig. 2 compares the amplitude modulations of modern tides in Boston Harbor (U.S.A.) (uniformly semidiumal tides with the principal variations following the changes in the Moon's distance and phase) with the spacing modulations of the laminae of the Mink stromatolites, obtained by using the running mean to smooth the noise (Dolman, 1975). The similarity of the two curves is striking and implies the same shaping cause. We know from data of modem analogs that tidal amplitude can effect stromatolitic growth and thus the lamina spacing. The most noticeable matching is between the modern six-month amplitude changes and the lamina thickness and the alternation of iron-rich and silica-rich bands. These bands, characteristic of the iron formation and the associated stromatolites, have been attributed to seasonal chemical changes in the environment of deposition. The Mink periodogram appears to be the result of lunar

680 G . PANNELLA

(spring-neap periodical changes) and solar (six-month chemical changes reflected in the Si/Fe ratio) effects.

There are some differences between the modem curve and the Huronian peridogram. Basically the modem curve, being a calculated prediction, is much more regular than the real one which records (as stromatolites do) all the noise spectra of the environment, but the Mink spectrum seems to contain not only more noise but some, perhaps fundamental, differences in tidal modulations. The similarity of the two curves supports the original interpretation of the tidal bands and of the annual band (Pannella, 197213). Still open is the question whether the laminae are daily or semidaily (in which case the length of the synodic month would not have changed in the last 2,000 m.y.). But statistical and coherence tests, hopefully, will answer the question. It must be pointed out that stromatolites can only provide minimum values and thus any recorded number of the days per year and per synodic month in the Precambrian higher than that at present indicates a definite deceler- ation in the rotation of the Earth around its axis, but leaves open the question of the rate. Moreover, only negative accelerations could be inferred from stromatolite data. Qualitatively, if not quantitatively, stromatolites have already provided the information that the Moon has been associated with the Earth for at least three aeons (the age of the oldest stromatolite with developed tidal bands) (Pannella, 197213). This information obviously eliminates all hypotheses implying a late lunar capture (Olson, 1966, 1968; Niini, 1969). Laminae can also be used to determine the amplitude of paleotides, as discussed in the following pages.

Stromatolites are tidal gauges

The discovery of Shark Bay stromatolites, in many aspects similar to those found in great abundance in Precambrian rocks, has influenced many of the current ideas and speculations on stromatolites. One of them, that the height of stromatolites could be a reflection of tidal amplitude, is based on the highly suggestive coincidence of the height of the Shark Bay stromatolites with the tidal amplitude of the area. Logan stated that their height is controlled “by the tidal range and the position of the structure in the intertidal zone” (Logan, 1961, p. 531). The way stromatolites grow forbids upward extension above the highest water level and it is likely that in a system in which this level remains unchanged for a long time, stromatolites will grow up to that level. Thus, Cloud’s (1968a) suggestion of using the height of Pre- cambrian stromatolites to infer the variations of tidal amplitude through time, and thus the moment of closest approach of the Moon to the Earth, is logical. By checking the literature and his notebooks, Cloud found that stromatolites supported the idea of a maximum tidal amplitude, and thus of a closest approach of the Moon, in the Middle Precambrian. Cloud’s data

HISTORY OF THE EARTH-MOON SYSTEM 681

were by no means complete or even representative and his argument was rather tenuous.

Walter (1970a) has stressed several points that make Cloud’s conclusion untenable. First, not all stromatolites with relief grew in intertidal areas; many were probably subtidal. The trend of increasing stromatolite height with increasing age is not at all evident from the stromatolites reported in the literature. If anything, they seem to indicate that the tallest stromatolites formed during the Paleozoic. Also tidal amplitude is the result, at any one locality, of the complicated interaction of many variables and cannot be used in any way even qualitatively, to estimate the distance from the Earth to the Moon. What amplitude could we use today as indicative of this distance: the 60-cm amplitude of Shark Bay or the 14 m of the Bay of Fundy?. The hopelessness of the task has been pointed out by Hohann (1973). To all these difficulties one can also add Monty’s suggestion that Precambrian gigantism may be related to a sequence of environmental circumstances and intrinsic algal factors rather than to any change in the amplitude of the tides (Monty, 197313). While Cloud’s generalization is untenable, the fact that stromatolite growth is governed by tidal fluctuations remains and. can be used as a method to infer tidal amplitude, at least locally. One way of doing it is to check the periodicity, in terms of laminae, of the tidal bands. As suggested in Table I, the number of laminae per band should change progressively from subtidal through intertidal to supratidal. Along the same bioherm, and perpendicularly to the shoreline, it is possible to check these variations and determine quite accurately the extent of the intertidal zone, and the tidal amplitude. So far stromatolites have been studied individually rather than collectively; studies that consider lateral as well as vertical morphological variations are needed and could provide useful insights and information. Along the same synchronous bed from deep to shallow water the number of laminae per periodical band should decrease. Vertical variations are often easier to study than lateral ones. Examples of morphological vertical evolution from columnar types to stratiform types are well developed in the Nash Formation, Medicine Bow Mountains (Wyoming). Many bioherms are capped by a finely laminated dolomitic bed. The rhythms: columnar stromatolites-stratiform stromatolites-dolomite are repeated several times and many indicate a change from intertidal (or slightly subtidal) to supratidal environment. The thickness of this cycle is consistently 1.5 m and represents, assuming the sea level was not changing significantly, the amount of growth necessary to reach a supratidal environment, that is, the tidal amplitude of the locality.

Bioherm evolution from columns to stratiform types lasted from 150 to 400 2.5 cm thick bands (Fig. 3). The preservation of the laminae is so poor that study of the vertical frequency changes is not possible. The thickness of the bands and the number of laminations still recognizable within each band decreases upward.

682 G. PANNELLA

In one type of the Mink stromatolites, the upward evolution of the lamina frequencies has been checked. Fig. 4 shows vertical morphological variations in the columns. As the number of laminae in the tidal bands decreases so also does the convexity, and the shape of the columns changes from erect to inclined, uniform to coalescent and linked. The evolution is also accompanied by the increasing abundance upwards of oolites between the columns. All seems to indicate an upward increase in the energy of the environment. The number of laminae per tidal band follows the trend and decreases from 14-17 to 3-9. This change occurs in a vertical distance of 15 cm. It is possible that this was the tidal amplitude. At least locally stromatolites may be able to provide accurate records of tidal range.

HISTORY O F THE EARTH-MOON SYSTEM 683

Fig. 3. A. Early Proterozoic stromatolites from the Nash Formation, Medicine Bow Mountains, Wyoming. Morphologically the bioherms show a vertical change from large columns to stratifonn stromatolites, to dentate laminae. The upper part is capped by laminated dolomite. These cyclic variations are successively repeated several times. Scale in inches. B. Lower part of a bioherm showing large columns and seasonal bands consisting of light dolomite and dark silica zones. Scale in inches.

CONCLUSIONS

The use of stromatolites as paleontological clocks presents many difficulties, but before concluding that these are always insuperable much more work is needed. The odds of determining the number of days per year with any

684 G. PANNELLA

Fig. 4. Periodical growth bands in Mink Mountain stromatolites (Gunflint Formation). Photographic print from a thin section used as negative. The number of laminae per tidal band ( t ) below the level A ranges from 14 to 17 , above the level from 3 to a maximum of 9. Boundaries between annual bands are marked by X. Scale 1 cm.

HISTORY O F THE EARTH-MOON SYSTEM 685

degree of accuracy are low indeed. But if not quantitatively, then qualitatively stromatolites can be useful. Figures should always be considered to represent minima and thus can only suggest decelerations (if indeed slowing down has occurred). Reconstruction of paleotide types and amplitude is a promising application. It could provide constraints on the evolutionary models of the Earth-Moon system.

Not all laminae have the same chronological meaning: they represent episodes of organic and inorganic deposition occurring with different frequencies ranging from tidal (12.4 h, 24.8 h, fortnight) to solar (24 hr, for one organic and one inorganic lamella) to random (i.e., storm laminae). These frequencies are characteristic of different environments from subtidal to supratidal. Using the criteria for the recognition of stromatolites from these different environments one can select the most promising samples for geophysical analysis. This selection is critically important. Subtidal stromatolites are more promising than intertidal ones. Only in an absolutely calm environment, where post-depositional erosion is unlikely to occur, however, may long and uninterrupted sequences of laminations survive. The environment most likely to provide the ideal sequences for reliable data must have the following characteristics: be absolutely calm, and be affected by seasonal, fortnightly, tidal and daily changes. On a pure numerological basis, when seasonal tidal bands are present, it is possible to recognize whether the sequence is continuous or interrupted (Pannella, 1972a, b). The degree of incompleteness can be used for separating subtidal from intertidal stromatolites. Most columnar stromatolites, morphologically similar to Shark Bay analogs, cannot be used as paleontological clocks because of their discontinuous growth. Small columnar types appear to contain longer periodical sequences. Harmonic analysis of the variations in thickness of the laminae should be used to support the interpretation of the periodical bands.

Keeping in mind that one or even several attributes (Hofmann, 1969a) do not provide unambiguous clues, the data from stromatolitic sequences should be critically evaluated and corroborated by as many other sources of information as possible. Their validity will be a direct outcome of a total and still faraway understanding of the physico-chemical and biological parameters at the basis of each attribute.

ACKNOWLEDGEMENTS

While assuming the full responsibility for the ideas expressed here, the writer wishes to thank those who have kindly helped with their comments, suggestions and materials, namely: J.D. Weaver, M.R. Walter, H.J. Hofmann, P.F. Hoffman, S.K. Runcom, G. Rosenberg, J. Dolman and Cristian Jones.

The research of this paper was made possible by the financial support of Minna-James-Heineman Foundation of Hanover (Germany).

APPENDICES

Appendix I

GLOSSARY OF SELECTED TERMS

M.R. Walter

There is not yet any wholly satisfactory system of terminology in use for stromatolites. The terms given below are mainly those used by Dr W.V. Preiss and myself (as published in Walter, 1972a) and by Hofmann (1969a). See also Fig. 1, Ch. 2.1 (p. 6) and the tables on pp. 688-689 and 692. Defi- nitions of other terms may be found in the text. Abiophoric - Of stromatolites, lacking microfossils. Banded microstructure - One in which the laminae are very continuous and

have abrupt, distinct, more or less parallel boundaries. Bioherm - A circumscribed organo-sedimentary structure whose minimum

width is less than or equal to one hundred times its maximum thickness, embedded in rocks of different lithology.

Biophoric - Of stromatolites, containing microfossils. Biostrome - A stratiform organo-sedimentary structure whose minimum

width is more than one hundred times its maximum thickness. Note: When the dimensions are unknown, the term “stromatolitic bed” may be used.

Branching - The division of a column into new, discrete columns. Note: columns become discrete when they are first separated by an interspace. Markedly divergent: Branching in which the axes of the new columns

diverge at more than 45’. Multiple: Branching at approximately the same level into more than two

new columns. Parallel: Branching in which the axes of the new columns are parallel.

Note: Most commonly, the axes of the new columns are also parallel to the axis of the original. a-parallel: Branching in which the width of the individual remains con-

P-parallel: Branching in which the original column widens gradually

y-parallel: Branching in which the original column widens abruptly

stant.

before branching.

before branching.

688 APPENDICES

OTHER ATTRIBUTES STROMATOLITES

Growth foclor A & leretc

5 Cyllndrcol

a Iurbinole

_L bulbous

A nodulor

- strollform

spheroid01

Thicknessof laminoe ( T I p . m m , conslonl. lapcrlnq

Relief of lominoe I h)

Height of slromotolile I H p , mm, cm, dm.m, elc

oiirnulh of elonqolion

azimuth ond omounl of incllnoliin of occrelion vector

Orientofion

Al lmde I rlroiqht I erect / inclined - horizonlo1

( curved /ncvmbml /decvmbenl ( U ~ W Y

@ cenlrifugol Pelroloqy , mineroloqy corbonole, colcble. dolomile.elc

51I1co

s~licolei

oxides. hydroxlder I Fe. Mnl

orqonic molter

open pore space

f l " l h

Iolilude. longitude. elevollon

Branching style W furcote

y" umbellok

diqifols

dendroid

8 coalescent

# onorlornosed

Geoqrophic position

Geologic sellinq

Geoloqic ope

rlrohgrophc ml. envaonmenl

year5

SURFACE

ORNAMENTATION nome of oriqinolinq wqonirmlsl qenvs. species

nomd of slromololib qroup. form

TAXONOMIC

NOMENCLATORIAL

INTERNAL FEATURES

Fig. 1. Graphical summary of attributes of stromatolites. (From Hofmann, 1969a; reproduced with the permission of the author and of the Geological Survey of Canada.)

APPENDICES 689

LAMINAE

COMPONENTS

r T

u - Cellvlor bofobnc

Curvolure lype and order T--

Profble :r 9 Q , 3

0

Plan outline

L L I N K A G E

S P A C I N G

RELIEF Irslolivel

DEGREE OF M M T A

lominoe liqhl ond dark lomelloe

even

'7 wavy I order

x 2 orders vv\ g corruqote - crenate

CM dentole & 151 3 orders

0 round circular. eIIIp1IcoI. ovole

0 oblong

0 scutote

w crescenllc

8 loxilobole

@ 5 densllobele

0 _o brevilobole

0 polygonal

0 IonceoIote

690 APPENDICES

Slightly divergent: Branching in which the axes of the new columns

Bridge - Stromatolitic lamina or set of laminae linking adjacent columns. Catagraphia - Microscopic carbonate problematica. Many are probably

grapestones, botryoidal lumps, and other sedimentary structures. Chemocline - The boundary between the circulating and the non-circulating

water masses or layers of a lake; specifically, the boundary separating the mixolimnion and the monimolimnion in a meromictic lake (q.v.).

Clone - The descendants produced asexually from a single organism (as genetically uniform population as is possible).

Clot - A microscopic segregation of pigment. Column - Discrete stromatolitic structure with the dimension in the direc-

I tion of growth usually greater than at least one of the transverse dimen-

Cornice - Overhanging lamina or set of laminae elongated transversely to its

Crestal line - Line joining the crests of successive laminae. Crestal zone - The environs of t he crestal line. Cryptalgal - Of sedimentary rocks or rock structures, those believed to

originate through the sediment-binding and/or carbonate-precipitating activities of non-skeletal algae.

Cryptalgalaminate - Of stromatolites, those that have more or less planar laminae (synonym: stratiform).

Cyanophyte - “Blue-green alga”. Endolithic - Of organisms, living within rock; specifically, boring micro-

organisms. Epilimnion - The uppermost layer of water in a lake, characterised by an

essentially uniform temperature that is generally warmer than elsewhere in the lake and by relatively uniform mixing caused by wind and wave action.

Eucaryote (Eukaryote) - Nucleated protists and all higher organisms (i.e. including all algae but not cyanophytes) .

Euphotic zone-That part of the ocean in which there is sufficient penetration of light to support photosynthesis (q.v.).

Gently convex - In stromatolite laminae, with a ratio of height to width less than or equal to 0.5.

Gliding - In procaryotes, slow movement not involving flagella. Grumous - Mineral texture in which fine-grained patches are surrounded by

Heliotropism - Tropism (q.v.) in which the stimulus is sunlight. Heterotrophism - In organisms, the requirement of a source of organic mat-

Homeostasis - Maintenance of constancy of internal environment. Individual - A group of columns arising from a single basal column or a dis-

diverge at 45’ or less.

sions.

column axis.

coarser grains.

ter (food) from their environment.

crete stromatolite within which the laminae are continuous.

APPENDICES 691

Interspace - The space between columns, usually filled with sediment. Lamina - The smallest unit of layering. Laterally linked stromatolite - With wavy laminae continuous between

Macrolamina - A set of laminae. Meromictic lake - A lake that undergoes incomplete mixing of its waters

during periods of circulation: specifically, a lake in which the bottom, non-circulating water mass (monimolimnion) is adiabatically isolated from the upper, circulating layer (mixolimnion).

Microphytoliths - Oncolites and catagraphia. Mixolimnion - See Meromictic lake. Monimolimnion - See Meromictic lake. Naked column - Column without walls. Oncolite - Unattached stromatolite with encapsulating laminae. Organelle - In organisms, a specialised part of the cell. Photoheterotrophism - Light-driven heterotrophism (q.v.). Photophobotaxis - Phototaxis (q.v.) in which the organism reacts to a gradi-

Photosynthesis - The conversion of COz to organic cell materials, using light

Phototaxis - Taxis (q.v.) in which the stimulus is light. Phototopotaxis - Phototaxis (q.v.) in which the organism detects the direc-

Phototrophism - In organisms, the obtaining of energy from light. Phototropism - Tropism (q.v.) in which the stimulus is light. Procaryote (Prokaryote) - Protist in which the genetic material is never

separated from the cytoplasm by a nuclear membrane (specifically, bacteria and cyanophytes).

Pseudocolumnar stromatolite - Laterally linked stromatolite in which successive crests are superimposed forming column-like structures.

Selvage - Unlaminated coating on column margins. Steeply convex - In stromatolite laminae, with a ratio of height to width

greater than 0.5. Stratiform stromatolite - Non-columnar stromatolite with flat continuous

laminae (cryptalgalaminate sediment). Streaky microstructure - One in which the laminae are moderately distinct

and continuous; the darker are usually the most distinct and they are set in a pale matrix into which they frequently grade vertically.

Striated microstructure - One in which the laminae originally formed as chains of lenses.

Synoptic profile - In stromatolites, the morphologic aspect at an instant of time during formation of the stromatolite.

Synoptic relief - In stromatolites, the relief of a stromatolite above its substrate at an instant of time during formation of the stromatolite.

crests.

ent in light intensity.

as a source of energy.

tion of the light source.

692 APPENDICES

Taxis - Locomotory movement of an organism or cell in response to a directional stimulus, the direction of movement being oriented in relation to the stimulus.

Thermocline (lake) - The horizontal plane in a thermally stratified lake located at the depth where temperature decreases most rapidly with depth.

Thrombolite - “Cryptalgal structures related to stromatolites, but lacking lamination and characterised by a macroscopic clotted fabric” (Aitken, 1967). I consider thrombolites to be a category of stromatolites.

Tropism - In an organism, response to stimulus by growth curvature, the direction of curvature being determined by the direction from which the stimulus orginates.

Vermiform microstructure - One in which narrow, sinuous, pale-coloured areas (usually of sparry carbonate) are surrounded by darker, usually fine- grained areas (usually carbonate).

Wall - Structure at the margin of a column, formed by one or more laminae from within the column bending down and coating the column margin for at least a short distance.

Wavy lamina - With flexures of wavelength greater than 2 mm. Wrinkled lamina - With flexures of wavelength less than or equal to 2 mm.

TABLE I

Classification of carbonate rocks according to depositional texture (from Dunham, 1962)

DEPOSITIONAL TEXTURE RECOGNIZABLE

Original components not bound together during deposition

Contains mud (particles of clay and fine silt size)

Mud-supported Grain-supported grain-supported

Less than More than 10% grains 10% grains

Lacks mud and is

Mudstone Wackestone Packstone Grainstone

Original components were bound together during deposition . . . as shown by intergrown skeletal matter, lamination contrary to gravity, or sediment-floored cavities that are roofed over by organic or questionably organic matter and are too large to be interstices

Boundstone ~~~ ~~~ ~

DEPOSITIONAL TEXTURE NOT RECOGNIZABLE

Crystalline Carbonate

(Subdivide according to classifications designed to bear cr. physical texture or diagenesis)

APPENDICES

Appendix II

TABLE OF TIME-RANGES OF THE PRINCIPAL GROUPS OF PRECAMBRIAN STROMATOLITES*

I .N. Krylov and M.A. Semikhatov

The compilation of a table of time-ranges of Precambrian stromatolite

(1) There are different concepts of many of the taxa (see Chapter 2.4). (2) Some students of Precambrian stromatolites refer to the taxa without

sufficient documentation (e.g. “graphical reconstruction’’ of gross morphology, investigating thin-sections, etc.), using only a general similarity of the constructions. That is the case especially for the pre-Riphean (Aphebian) stromatolites with a Middle or Late Riphean aspect. The absence of paleontological descriptions of the great majority of such stromatolites makes it impossible to elucidate the problem of whether these stromatolites represent convergent evolution or whether they are the same as the Middle and Late Riphean ones (see Chapters 7.1 and 7.2).

(3) The isotopic data on the stromatolite-bearing sections are rather scarce and in some cases contradictory (e.g. Adelaide Supergroup). Moreover, some of the well-known localities with rich stromatolite assemblages are not isotopically dated.

In order to reduce the influence of the difficulties mentioned above, we accept the following:

(a) the concepts of the taxa should in general be as close to the original as possible. In the case of Gymnosolen and other groups described long ago, we follow the most widely held concept (e.g. for Gymnosolen -that of Raaben, 1964; see Chapter 2.4).

(b) In general, only fully described taxa. should be listed on the Table. (c) The succession of assemblages of taxa does have time significance, even

if it is not supported by available isotopic age determinations. The experience of interregional and intercontinental sorrelations using Precambrian stromatolites supports this assumption.

taxa is a rather difficult task because of at least three reasons:

* See also the Figure on p. 347.

694 APPENDICES

." Cambrian

Lerva

Linella

Limollenia

Madiganifes

Microstylus

Mtijaria

Nouafila

Nucleella

Omachtenia

Cambrian

Acaciella

Alcherrnga

Aldonia

Anobaria

Baicalia

Basisphaero

Boxonia

Colleniella

Collumnocoilenia z ? Colonnella

Collumnaefacta

Compacfocallenia

Conophflon

Dgerbia Eucapsiphora

Gongylina

Gruneria

Gymnosoten

Jacufophyton :? Gearginia

Jurusania

Inzerio

Zrregulario

Uasaia

Uafavia

Uaternia

Uafuikania

Uulparia

Yusiella

Aphebion

Paniscallenia

Putomia

Parmifu Pitella

Plonocollha

poludia Pilbaria

Sacculia

Schancharia

Serizia

Sfratifera

Svetliella

Tarioufetia

T ~ m s i n a

TiIbunkeia

Timia

Tlmgusria

Tunic&

Vetella

The points mentioned above have to be kept in mind when studying the Table. The Table must be taken as preliminary - we well realize how much still has to be done before it will be possible to compile a solid chart of time-ranges of Precambrian stromatolite taxa. We hope the present one will help to solve the problem.

APPENDICES

Appendix 111

LIST OF AVAILABLE TRANSLATIONS OF MAJOR WORKS ON STROMATOLITES

M.R. Walter

The library of the Geological Survey of Canada has a collection of translations into English of many geological publications. Complete lists can be found in Papers 70-62 and 72-16 of the Survey. The translations are available as Xerox copies. Address orders to Library Services, Geological Survey of Canada, 601 Booth Street, Ottawa K1A OE8, Ontario, Canada. Those of interest here are listed below.

Komar, V.A. Upper Precambrian stromatolites in the north of the Siberian Platform, and their stratigraphic significance. 336 pp. (Stromatolity verkhnedokembriyshikh otlozheniy severa Sibirskoy platformy i ikh stratigraficheskoe znachenie. Tr. Geol. Inst. Akad. Nauk S.S.S.R., 154: 122 pp. (1966).)

Komar, V.A., Raaben, M.E. and Semikhatov, M.A. Conophytons in the Riphean of the U.S.S.R. and their stratigraphic importance. 151 pp. (Konofitony Rifeya S.S.S.R. i ikh stratigraficheskoe znachenie. Tr. Geol. Inst. Akad. Nauk S.S.S.R., 131: 72 pp. (1965).)

Krylov, I.N. Riphean and Lower Cambrian stromatolites of Tien-Shan and Karatau, 125 pp. (Rifeyskie i nizhnekembriyskie stromatolity man’-Shanya i Karatau. D. Geol. Inst. Akad. Nauk S.S.S.R., 171: 88 pp. (1967).)

Menner, V.V. e t al. Paleontological basis of the stratigraphy of the upper Precambrian.

(Paleontologicheskoe obosnovanie stratigrafii verkhnego Dokembriya. In : Verkhniy Dokembriy, Stratigrafiya S .S .S .R , 2 . Moscow, pp. 475-505 (1963).)

Raaben, M. Upper Riphean stromatolites (gymnosolenides). Translated from Russian, Translation Bureau, Dep. of the Secretary of State, Ottawa, Ont. 1970. (228) 1. (Stromatolity verknego Rifeya (gimnosolenidy). Tr. Geol. Inst. Akad. Nauk S.S.S.R., 203: 100 pp. (1969).)

Two translations from Russian into ‘French are available from the

86 PP.

B.R.G.M. Service de documentation, BP 555,45 018 Orleans (France): Krylov, I.N., 1963. Stromatolites columnaires ramifies du Ripheen de 1’Oural sud e t leur

importance pour la stratigraphie du Precambrien supkrieur. Traduction No. 4531. Maslov, V.P., 1960. Les stromatolites. Gknese, methodes d’etudes, lien avec les facies et

valeur geologique, d’aprhs I’example de I’Ordovicien de la plate-forme siberienne. Traduction No. 4518.

696 APPENDICES

The following review and condensed translation from English to Russian is available :

Semikhatov, M.A., 1973. Korrelyatziya verkhnego dokembriya Avstralii i Severo-Vostoka S.S.S.R. Ekspress-informatziya. Seriya: Obshchaya i regionalnaya geologiya: geologicheskoe kartirovanie, No. 10. VIEMS, 1973: 1-39. (Review and translation of: Walter, M.R.., 1972. Stromatolites and the biostratigraphy of the Australian Precam- brian and Cambrian. Palaeontol. Assoc. Lond., Spec. Pap., 11: 190 pp.; and Walter, M.R. and Preiss, W.V., 1972. Distribution of stromatolites in the Precambrian and Cambrian of Australia. Znt. Geol. Congr., 24 th Sess., Montreal, 1972, Sect. 1 , Pre- cambrian Geology : 85-93.)

APPENDICES

Appendix IV

SELECTIVE SUBJECT INDEX TO THE BIBLIOGRAPHY

S.M. A wramik

The bibliography includes 1456 references dealing with stromatolites. Broad categories are used in the index. The numbers refer to the references listed in the bibliography.

GEOLOGIC AGE

Recent

11,12,13, 31,32, 37,48,58,59,67,87, 96,97,98

101,102,103,104,109,154,159,160, 165,171,172,183,199

206,207,208,211,212,213,214,215, 216,217,218,219,220,235,237, 239,272,274,280,281,286

300,320,321,336,347,356,364,382 383,388

406,407,409,411,417,447,495,497, 49 8

513,524,528,568,570,580,581,582, 583,584,585,586,587,588,589, 590,591,592,594,595

611,612,617,618,619,620,621,622, 623,624,625, 626,627,628, 638, 641, 642, 643, 644, 645, 646, 647, 648,649,650,659,660,661, 677, 679, 683, 685, 686, 687, 688, 689, 690,693,696

711,712,723,750,751,752,753,754, 755,756,764,779,782,788,796, 799

800,807,808,818,828,846,847,852, 895,896,897,898,899

903,934,936,937,938,955,956,959, 960,962,963,969

1044,1049,1050,1051,1079,1083 1103,1125,1126,1127,1128,1130,

1131,1132,1133,1134,1140,1141, 1156

1281,1282,1283,1284,1285,1286, 1287,1288

1304,1305,1307,1327,1330,1360, 1364,1365,1387,13&9,1394

1402,1411,1425,1438,1444,1464, 1495

1507,1525,1534,1535,1554,1577, 1582,1592,1594,1595

1615,1620,1622,1645,1651,1652, 1659,1661,1663,1677,1679

1702,1703,1729,1734,1741,1742, 1753,1766,1778,1781,1782,1783, 1790,1791,1792

1223,1239,1256,1263,1264,1280,

1880,1894,1895,1896,1897 1900,1901,1902,1904,1905,1914,

1915,1916,1920,1925,1946,1975, 1982,1984

2027,2029,2032,2034 2019,2020,2021,2022,2023,2025,

Cenozoic

41 145,187,188,189,191,192,194 233,240,241,243 37 3

69 8

445,488 515,525,578,579,597 635 721,781 869,871,874,887 1063,1088 1201,1203,1286 1305 1515,1527,1547,1553,1589 1732,1736 1881 1918,1919,1926,1947,1960

APPENDICES

530, 531, 534, 535, 538, 539, 540,551,

600,614,615,640,666,667,695 714,715,726,727,734,735,738,747,

760,761,767,773,774,776 808, 816, 817,821,824,829,830,831,

839,870,873,874,876,877,879, 881,888,890,892

902,905,907,908, 935,941,942,943, 971,985,996,998,999

1001,1002,1003,1004,1005,1009, 1010,1012,1013,1015,1016,1018, 1019,1020,1021,1025,1026,1030, 1032,1033,1034,1035,1036,1038, 1040,1065,1066,1067,1068,1069, 1071,1077,1082,1097,1098

1100,1101,1102,1111,1114,1124, 1137,1138,1143,1153,1164,1175, 1178,1181,1182,1183,1186,1189, 1190,1191,1194,1195,1196,1197, 1199

1222,1245,1250,1253,1254,1279, 1286,1299

1363,1374,1384

1428,1429,1430,1437,1439,1453, 1455,1468,1482,1484,1485,1486, 1496

1560,1573,1579,1580

1653,1666,1669,1680,1681,1689, 1690,1691,1696,1699

1700,1701,1720,1721, 1725, 1739, 1754,1755,1756,1757,1764,1765, 1768,1772,1785,1788,1790

1809,1854,1855,1856,1857,1861, 1864,1869,1877,1884,1888,1890, 1892,1893,1899

1939,1942,1943,1950, 1951,1953, 1957,1958,1959,1962,1966,1968, 1978,1980,1994,1995,1999

2000,2001,2003,2004,2005,2007, 2008,2009,2012,2015,2026

556,565,571

1200,1202,1212,1213,1220,1221,

1324,1325,1342,1356,1361,1362,

1415,1416,1417,1424,1426,1427,

1526,1538,1539,1540,1541,1546,

1609,1623,1624,1626,1628,1650,

1900,1901,1903,1907,1909,1921,

Mesozoic

8,9,70 119,180,181 290 348,381,390,391,393,394 403,466,471,472,493 506,510,527,555,557,572,577,578,

579,596 616,632,669 709,730,740,775 837,866,874,893

1083 1167 1206,1286 1335,1338,1398 1434,1463,1491,1492 1503,1504,1533,1543,1545,1556,

1625 1717,1744,1745,1746,1748,1749,

1851,1852,1885,1886 1952

944,945,977

1565,1566

1751

Paleozoic

14,15,16,17,20,27, 33,40,69,94,95 108,131,141,142, 150,156,168,169,

170,196,198 201,203,226,242,261,284,285,288,

297 307; 310,316,322,327,338,360,368,

374, 380,384,385,386, 389,393, 39 5

411,412,426,427,430,431,454,462, 463,464,465,491

Proterozoic

1,7,19,24,30,42,46,49, 52, 54,55, 56,60,61, 62,63,65,72,75,76,77,

APPENDICES

80,81,82,83,84,85,86,88,90,92 100,110,111,112,113,114,115,116,

117, 122,123,127,128,130,132, 133, 134,135,136,137,138,139, 140,143,144,146,147,149,151, 157,163,173,174,175,178,179, 18 5

229,230,232,234,249,250,251,252, 253,254,255,256,257,258,259, 263,264,265,266,271,275,276, 277,278,287,291,292,293,294, 296,298,299

304,305,306,308,309,314,327,328, 329, 330,331,332, 333,334,335, 337,339,340,342, 344,345,350, 355,365,366,372,378,385,389, 396,397,398,399

400, 401,402,413,414,415,416,421, 424,425,428,436,438,439,440, 441,442,443,444,448,449,450, 451,452,453,455, 456,457,458, 459,460,461,467,468,473,474, 475,476,483,492,494,496

509,514,516,518,519,520,521,522, 523,529,532,533, 536,538,539, 540, 541,542,543, 544,545,546, 547,548,549,550,551, 552,553, 554, 556,562,563, 564,569,575, 598,599

601, 602,603,606,608,612,613,621, 624,626,627,630,634,639,642, 643,647,648,650,651,653,654, 655,656,657,658,663,670,671, 672,673,674,675,685,690,698, 699

701, 702,703,704,705,706,710,716, 718,719,720,722,728,729,731, 732,733,741,742,748,759,762, 766,770,772,777,789,790,791, 792,793,794,795,796,797,798

801,802,803,804,805,806,807,808, 809,810,811,816,820,825,826, 832,833,834,835,836,851,855, 859,860,863,864,865,868,874, 877,889,890,891

900,904,911,914,915,916,917,918, 919,920,921,922,923,924,925, 926,927,928,929,930,931,932, 933,946,947,948,950,951,952, 953,957,964,965,966,967,968, 972,973,974,975,978,979,980, 981,982,983,984,985,986,988, 989,990,991,992,993,994,995, 997,998,999

699

1000,1003,1008,1010,1011,1013, 1014,1015,1016,1017,1019,1022, 1023,1024,1027,1029,1031,1032, 1033,1034,1036,1037,1038,1041, 1046,1047,1048,1052,1053,1054, 1055,1056,1057,1058,1059,1060, 1061,1062,1064,1065,1066,1067, 1068,1069,1070,1071,1072,1073, 1074,1075,1076,1077,1078,1080, 1099

1119,1120,1121,1123,1130,1135, 1136,1144,1146,1147,1148,1149, 1150,1152,1154,1155,1158,1159, 1160,1161,1162,1163,1168,1173, 1174,1175,1177,1179,1180,1181, 1182,1183,1184,1185,1186,1187, 1188,1190,1193,1194,1195,1197, 1198,1199

1214,1215,1216,1217,1218,1222, 1224,1228,1229,1238,1240,1241, 1244,1245,1246,1249,1251,1252, 1257,1258,1259,1260,1261,1262, 1266,1267,1269,1270,1271,1274, 1275,1276,1277,1278,1280,1286, 1287,1290,1292,1293,1294,1297,

1317,1318,1319,1320,1321,1328, 1331,1337,1340,1341,1342,1343, 1344,1345,1346,1347,1348,1349, 1353,1354,1355,1366,1377,1378, 1379,1380,1381,1382,1383, 1392

1401,1408,1409,1410,1431,1433, 1436,1442,1443,1448,1449,1450, 1451,1452,1453,1454,1455,1465, 1466,1467,1468,1469,1470,1471, 1472,1473,1474,1475,1476,1477, 1478,1479,1480,1481,1483,1487, 1488,1489,1494,1495,1496, ? 497, 1498,1499

1519,1522,1524,1529,1530,1531, 1532,1540,1541,1544,1546,1555, 1558,1563,1564,1569,1576,1581, 1584,1585,1586,1587,1588,1597, 1598,1599

1613,1614,1619,1627,1628,1629, 1630,1631,1632,1633,1634,1635, 1636,1637,1638,1639,1640,1641, 1642,1643,1644,1645,1646,1647,

1105,1110,1115,1116,1117,1118,

1200,1202,1207,1208,1209,1213,

1301,1302,1313,1314,1315,1316,

1505,1508,1509,1510,1511,1513,

1600,1602,1603,1604,1605,1608,

700

1648,1649,1655,1656,1657,1664, 1665,1668,1674,1675,1682,1683, 1686,1687,1688,1689,1690,1691, 1692,1693,1694,1695,1696,1697, 1698

1725,1726,1730,1752,1760,1761, 1763,1769,1771,1785,1786,1787, 1789,1793,1794,1795

1806,1809,1810,1814,1817,1818, 1819,1820,1821,1822,1823,1824, 1832,1833,1835,1836,1837,1838, 1839,1841,1842,1843,1844,1845, 1846,1847,1848,1849,1850,1853, 1854,1855,1856,1857,1858,1859, 1860,1861,1862,1864,1865,1866, 1867,1868,1869,1870,1871,1872, 1873,1874,1875,1876,1877,1878, 1883,1884,1887,1889,1891,1899

l930,1931,1932,1933,1934,l948, 1949,1955,1964,1966,1969,1970, 1971,1972,1973,1974,1976,1977, 1979,1981,1982,1983,1984,1985, 1986,1987,1989,1990,1991,1992, 1993,1998

2002,2003,2004,2005,2006,2007, 2008,2010,2011,2012,2013,2014, 2015,2016,2017,2018,2024

1708,1711,1712,1713,1714,1724,

1800,1801,1802,1803,1804,1805,

1900,1901,1903,1927,1928,1929,

APPENDICES

Archean

164,882,883,1151,1350,1400,1560, 1607

GEOGRAPHIC AREA

Africa

11,37,55,98 130,131, 132, 133, 134, 135, 136, 137,

138,139,140,141,142,143,144, 147,164,165,169,170,178,179, 185

254,255,256,257,258,259,263,264, 265,266,286

304,305,306,307,308,309,310,344, 372

428,436,475,492

518, 519,520,521,522,523,575 606,608 710,719,721,728,732 836,863,864,865,869 1040,1041,1048,1099 1110,1123,1150,1151,1152,1162,

1244,1245,1246,1252,1275,1276,

1313,1314,1315,1331,1350,1377,

1400,1431,1436,1448,1464,1494 1507,1519,1527,1587,1588,3.595 1600,1607,1613,1614,1619,1659,

1734,1761,1763,1769,1786,1793,

1800,1801,1802,1803,1832,1833,

1955,1981,1982,1983,1984,1985,

1177

1277

1378,1386,1389

1661,1682,1683

1794,1795

1838,1881

1986,1987

An tarct ica

198

Australia

60,67,90 165 320, 344,355,365,384, 385,386,397 406,407 532 655,656,657,658,687 742,767,799 800,824,825,826,833,834,835 1119,1121,1125,1126,1127,1128,

1129,1130,1131,1132,1133,1134, 1154,1155

1222,1223,1224,1275 1301 1424,1425,1426,1427,1428,1429,

1430,1449,1450,1451,1452,1453, 1454,1455

1524,1597,1598,1599 1600,1602,1603,1604,1608 1760,1766,1789 1900,1901,1902,1903,1905,1957 2026,2031,2034

APPENDICES

Eurasia

1, 7,11, 12,13,19,27,40,42,46,52, 54,60,70,87

104,115,116,117,119, 143,144,145, 146, 150,151,166, 180, 181,183, 199

240,241,243,246,247,249,250,251, 252,253,271,272,290,296,297, 298,299

300,314,344,347,348,350,381,390, 391,392,393,394, 395, 396

403,426,427,430, 431,444,445,449, 450,451,452,453,462,463,464, 465,466,471,472,473,474,483, 49 3

509, 510,515, 516, 524, 528,529, 533, 534, 535,536,555, 565, 572,576, 577,578,579,580,581,582,583, 587,588,590,591,592,594,596, 598,599

601,602,603,613,614,615,616,621, 624, 630, 632, 634, 635, 651, 658, 659,669,670,671,672,673,674, 675,677,679,683,685,686,687, 693,695,696

701,702,703,706,709, 715,716,718, 723,726,727,730,731,733,740, 759,771,773,775,779,781,782, 783,788

816,817,837,846,847,851,852,866, 888,889,890,891,893,895,896, 897,898,899

900,902,903,904,905,911,914,915, 916,917,918,919,920,921,922, 923, 924,925, 926, 927, 928,929, 930,931,932,933,936,937,938, 950,951,952,953,959,960,962, 963,964,965,966,967,968,972, 973,977,978,979,980,981,982, 983,984,985,986,988,989,990, 991,992,993,994,995,996,997, 998,999

1000,1001,1002,1003,1004,1005, 1008,1009,1010,1011,1012,1013, 1014,1015,1016,1017,1018,1019, 1020,1021,1022,1023,1024, 1025, 1026,1027,1029,1030,1031,1032, 1033,1034,1035,1036,1037,1038, 1047,1050,1051,1052,1053,1054, 1055,1056,1057,1058,1059,1060, 1061,1062,1063,1064,1065,1066, 1067,1068,1069,1070,1071,1072,

701

1073,1074,1075,1076,1077,1078, 1079,1083,1088,

1143,1144,1146,1147,1148,1149, 1156,1158,1159,1160,1161,1163, 1167,1175,1178,1179,1180,1181, 1182,1183,1184,1185,1186,1187, 1188,1189,1190,1191,1192,1193, 1194,1195,1196,1197,1198,1199

1206,1211,1213,1242,1251, 1253, 1256,1257,1258,1259,1260,1261, 1262,1266, 1267,1275,1279,1297

1320,1321,1325,1335,1337,1338, 1342,1343,1344,1345,1346,1347, 1348,1349,1354,1364,1366,1398

1439,1442,1463,1465,1466,1467, 1468,1469,1470,1471,1472,1473, 1474,1475,1476,1477,1478,1479, 1480,1481,1482,1483,1484,1485, 1486,1491,1492,1497,1498,1499

1538,1539,1540,1545,1546,1547, 1556,1560,1565,1566,1576, 1577, 1581,1582

1631,1632,1633,1634,1635,1636, 1637,1638,1639,1640,1641,1642, 1643,1644,1645,1646,1648,1649, 1655,1656,1657,1664,1665,1668, 1669,1674,1679,1685,1686,1687, 1688,1689,1690,1691,1692,1693, 1694,1695,1696,1697,1699

1700,1701,1708,1717,1724,1725, 1730,1739,1741,1742,1744,1745, 1746,1748,1749,1751,1772,1778, 1785,1787,1791,1792

1847,1848,1849,1850,1853,1854, 1855,1856,1857,1858,1859,1860, 1861,1862,1864,1865,1866, 1867, 1868,1869,1870,1871,1872,1873, 1874,1875,1876,1877,1878,1883, 1884,1885,1886,1887,1894,1895, 1896,1897,1899

1962,1964,1968,1969,1970,1971, 1972,1973,1974,1975,1976,1977, 1989,1990,1991,1992,1993, 1994, 1995,1998

1101,1102,1103,1114,1115,1116,

1200,1201,1202,1203,1204,1205,

1304,1307,1316,1317,1318,1319,

1401,1415,1416,1417,1434,1437,

1503,1504,1505,1511,1515,1533,

1615,1625,1627,1628,1629,1630,

1804,1805,1835,1836,1837,1846,

1925,1926,1939,1949,1950,1952,

2002,2003,2004,2005,2006,2007,

702

2008,2009,2010,2011,2012,2013, 2014,2015,2016,2017,2018,2027, 2032,2033

India

56,77,80,81,82 287, 291,292,293,294 442 704,705,766 860,946,947,948 1046,1080 1135 1207,1208,1209,1238,1269,1270,

1302,1379,1380,1392 1433,1487,1488,1489 1544,1564,1584,1585,1586 1650 1711,1712,1713,1714 1817,1818,1819,1820,1821,1822,

1271

1823,1824,1839,1841,1842,1843, 1844,1845

North America

8,9,11,14,15,16,17,20,24, 31,32, 33,41,48,49,59,60,62,63,65,69, 72,75,76,83,84,85,86,88,92,94, 95,97,98

100, 101,102,103, 108, 109,110,111, 112,113,114,122,123,127,128, 154,156,157,159,160,163,171, 172,173,174,175,187,188,189, 191,192,194,196

201,203,206,207,208,209,211,212, 213,214,215,216, 217,218,219, 220,226,229,230,232,233,234, 235,237,239,242,261,274,280, 281,284,285,288

316,321,322,327,329,332,333,335, 336, 338,340,342, 344,345,356, 360,364,366,368, 373, 374,378, 380,382,383,388,389,398,399

409,411,412,413,417,421,424,425, 438,439,440,441,443,447,454, 455,456,457,458,459,460,461,

497,498 506,513,514,525,527, 530,531,538,

539,540,541, 542, 543,544, 545,

46.7,468,476,488,494,495,496,

APPENDICES

546,547, 548, 549, 550, 551,552, 553,554,560, 562, 563, 564, 568, 569,570,571,595,597

600, 611,612,620, 623, 625, 626, 627, 638,639,640,641,642, 643, 644, 645,646,647,648,649,650,660, -661,663,666,667, 685,687, 688, 689,690,698,699

714,720,722,729,734,735,733, 741, 747,750,751,752,753,754,755, 756,760,761,762,764,770,772, 774,776,777,789,790,791,792, 793,794,795,796,797,798

801,802,803,804,805, 806,807, 808, 809,810,811,820,821,828,829, 830, 831,832,839, 855, 868, 870, 871,872,873,876,877, 879,881, 882,883,887,892

907,908,934,935,941,942,943,955, 957,969,971,974,975

1044,1082,1097,1098 1100,1105,1111,1117,1118,1120,

1212,1214,1215,1216,1217,1218, 1124,1136,1137,1138,1153,1164,

1220,1221,1228,1229,1240,1241, 1249,1250,1254,1263,1264,1280, 1281,1282,1284,1288,1290,1292, 1293,1294,1299,

1353,1355,1356,1360,1361,1362, 1363,1374,1381,1382,1383,1384, 1387,1394

1444,1496

1526,1529,1530,1531,1532,1540, 1541,1543,1553,1554,1555,1569, 1573,1578,1579,1589,1592,1594, 1598

1623,1624,1651,1652,1653,1662, 1663,1666,1675,1677,1680,1681

1736,1752,1753,1754,1755,1756, 1757,1764,1765,1768,1771,1781, 1782,1783,1788,1790

1890,1891,1892,1893

1918,1919,1920,1921,1927,1928, 1929,1930,1931,1932,1933,1934,

1953,1960,1966,1978,1979,1980,

1305,1324,1327,1330,1340,1341,

1408,1409,1410,1411,1438,1443,

1508,1509,1510,1513,1522,1525,

1600,1605,1608,1609,1620,1622,

1703,1720,1721,1726,1729,1732,

1806,1809,1814,1880, 1888,1889,

1904,1907,1909,1914,1915,1916,

1942, i943,1946,1947,1948,1951,

APPENDICES 703

1984,1999

2025,2029 2000,2001,2019,2020,2021,2022,

Oceans and islands

48,58,97,98 151,154,171,172 239 373 417 617,618,619,623,645,649,687,689 708,712,750,751,752,754,755,764 1031,1044,1053,1072 1140 1239,1280,1281,1282,1284,1285,

1288 1327,1330 1402 1535 1618,1620,1651,1652,1663 1702 1984

South America

149,168 275,276,277,278 339 400,401,402,414,415,416,491 1174 1274 1698 1851,1852

OTHER CHARACTERISTICS

Microfossils, in stromatolites

61,62,63,65,82,88,89,90,92 131,150,187 329,340,345 803,804,805,809,811,836 1039 1117,1118,1119,1120,1121,1152 1313,1314,1315,1350,1380 1560,1597,1598,1599 1600,1602,1603,1604,1605,1608 1814,1869,1873,1877

1931

Non-carbonate

31, 32, 58, 59, 62, 63, 65, 82, 88, 90,92 187 206,207,208,209,211,212,213,214,

329,336,340,345 412,421,496 698,699 7 20 803,804,805,809,811 1292,1293,1294 1360 1597,1598,1599 1600,1602,1603,1604,1605 1752 1806,1814 2019,2020,2021,2022,2023

215,216,217,218,219,220

Non-marine

11,12,13, 31, 32,87 101,102,103, 104, 108, 142,145, 154,

159, 160,183,187, 188,189, 191, 192,194

206,207, 208,209,211,212,213,214, 215,216,217, 218, 219, 220,233, 235,237,240,241,280,281

320,321,327,328,336,347,364 447,462,488,495,497,498 513, 525, 578, 579, 585, 586, 587, 588,

590,591,592,593,594,597 635,659,660,661,683,685,688 723,782 828,846,847,852,870,871,887 956,973 1103,1156 1256,1263,1264,1280,1281,1282,

1286,1288 1304,1307,1360,1364,1394 1438,1439,1444 1525,1527,1553,1554,1577,1582,

1589,1594 1615,1622 1729,1732,1736,1741,1742,1753,

1778 1894,1895,1896,1897 1902,1904,1914,1915,1916,1918,

1919,1920,1947,1957,1960,1975

704 APPENDICES

Oncolites

1 146 241 314,390,391,393,394 445,449,451,452,471,472 515,557,562,563,564 600,603,659,675,677 870,879,889,890,891 902,950,951,952,953,956,971 1031,1048,1059,1071,1075,1083 1175,1190,1194,1197,1199 1223,1257,1258,1259,1260,1261,

1262,1297

1304,1307,1317,1318,1319,1320,

1444,1476,1477,1483,1497,1498,

1505,1525,1534,1536,1537,1540,

1619,1675,1677 1725,1744,1745,1746,1748,1749,

1785,1788 1809,1874 1909,1918,1919,1960,1968,1971,

2002,2003,2004,2005,2006,2007,

1321,1324,1327

1499

1541,1579,1595,

1989,1990,1991,1992,1993,1998

2008,2009,2010, 2011,2012,2013, 2014,2015,2016,2017

BIBLIOGRAPHY’

S.M. Awramik, H.J. Hofmann and M.E. Raaben (compilers)

INTRODUCTION

The Bibliography cites all the references mentioned in the authors’ texts. We have included several hundred additional references on stromatolites, making this the most comprehensive collection of stromatolite references yet published. The bibliography, though comprehensive, is by no means complete. Broad gaps exist, particularly where there is only a casual mention of a stromatolite in a paper or where works are in languages employing non-Roman alphabets - notably Chinese. In some cases, Russian and other references lack page numbers and volume numbers. These references were unavailable to the compilers but are included for the reader’s information.

1 Abdullaev, R.N.. Akhmedzhanov. M.A. and Borisov, O.M., 1972, Occurrence of oncolites in the Precambrian of northern Nuratau. U z b . Geo l . Z h . . 1972 (1): 111 (in Russian).

2 Abelson, P.H.. 1967. Conversion of biochemicals to kerogen and n-paraffins. In: P.H. Abelson (Editor). Researches in Geochemis try , 2. Wiley. New York. N.Y.. pp. 63-86.

3 Abelson. P.H. and Hare, P.E.. 1969. Recent amino acids in the Gunflint chert. Carnegie Inst. Wash , Yearb . . 67: 208-210.

4 Abelson. P.H. and Hare, P.E., 1970. Uptake of amino acids by kerogen. Carnegie Inst. Wash.. Yearb . , 68: 297-303.

5 Abelson, P.H. and Hare, P.E.. 1971. Reactions of amino acids with natural and artificial humus and kerogens. Carnegie Inst. Wash., Yearb . . 69: 327-334.

6 Abelson, P.H. and Hoering. T.C.. 1961. Carbon isotope fractionation in formation of amino acids. Proc. Nat l . Acad . Sci. U S . . 47: 623-632.

7 Ablizin. B.D.. Kurbatskiy. A.M. and Krylov, I.N., 1969. Upper Precambrian stratigraphy in the north Urals western sector. I z u . A k a d . Nauk S .S .S .R. , Ser. Geol . , 1969 (9): 108-113 (in Russian).

8 Achauer, C.W.. 1967. Petrography of a reef complex i n Lower Cretaceous James Limestone. Bull. A m . Assoc. Pet . Geo l . , 51: 452 (abstract).

9 Achauer. C.W. and Johnson, J.H., 1969. Algal strornatolites in the James Reef Complex (Lower Cretaceous). Fairway Field, Texas. J. Sediment . Pe tro l . . 39: 1466.

10 Ackman, R.G.. Tocher, C.S. and McLachlan. J.. 1968. Marine phytoplankter fatty acids. J. Fish. Res . Board Can. , 25: 1603-1620.

11 Adolphe. J.-P.. 1970. Etude comparde des calcins en choux fleurs du Moyen Nord Canadien, d e France e t d u Liban. C . R . Acad . Sci. , Paris, 270: 1080-1083.

12 Adolphe, J.-P., 1973. The carbonate encrustations of the Pont du Gard Aqueduct, France. C . R . Acad . Sci. , Paris, Ser. D . 277(21): 2329-2332 (in French).

13 Adolphe, J.-P. and Rofes, G.. 1973. Calcareous concretions from t h e Levri6re. a tributary of the Epte River, subtributary of the Seine. Assoc. Fr. Etud. Quaternaire. Bull . . lO(35): 79-87 (in French).

14 Ahr, W.M.. 1967. Origin and Paleoenvironment of some Cambrian Algal Ree f s , Mason County Area , Texas . Thesis, Rice University, Houston, Texas, 104 PP. (UNv. Microfilms, Ann Arbor. 67-13.049.)

15 Ahr. W.M.. 1969. Paleoenvironment. algal structures. and fossil algae i n Upper Cambrian of central Texas. Bull. A m . Assoc. Pet . Geo l . . 53(3): 703 (Abstract).

See also additional references on p. 771.

706 BIBLIOGRAPHY

16 Ahr. W.M.. 1971. Paleoenvironment. algal structures and fossil algae in the Upper Cambrian of Central Texas. J. Sediment . Pe tro l . , 41: 205-216.

17 Aitken, J.D.. 1967. Classification and environmental significance of cryptalgal limestones and dolomites, with illustrations from the Cambrian and Ordovician of southwestern Alberta. J. Sediment . Pe tro l . . 37: 1163-1178.

18 Akiyama,M. and Johns. W.D.. 1972. Amino acids in the Cretaceous Pierre Shale of eastern Wyoming. North America. Pac. Geo l . , 4: 79-89.

19 Akulshina. E.P.. Davydov. Y.V., Pisarev, V.D. and Pisareva, G.M., 1969. Lithological and geochemical peculiarities and conditions of the Middle Riphean carbonate strata formation in the Mayan Depression. In: A.P. Vinogradov, Atlas o f Lithological and Paleogeogmphical Maps o f the U.S.S.R.. I . Tr . S ib . Nauchno-lssled. Ins t . , Geo l . , Geophys . Miner. Ser., 98.

20 Alberstadt. L.P. and Walker, K.R.. 1974. Patch reefs in the Carters Limestone (Middle Ordo- vician) in Tennessee, and vertical zonation in Ordovician reefs. Bull. Geol . SOC. A m . , 85(7): 1 171-1 182.

21 Albrecht. P. and Ourisson, G., 1969. Diagdnhse des hydrocarbures saturds dans une sdrie se'di- mentaire e'paisse (Douala, Cameroun). Geochim. Cosmochim. A c t a . 33: 138-142.

22 Albro, P.W. and Dittmer, J.C.. 1970. Bacterial hydrocarbons: occurrence, structure and metabolism. Lipids , 5: 320-325.

23 Albro. P.W. and Huston. C.K., 1964. Lipidsof Sarcina lutea 11. Hydrocarbon content of the lipid extracts. J. Bacteriol.. 88: 981-986.

24 Alc0ck.F.J.. 1938. Geology of Saint John regions. New Brunswick. Geol . Surv. Can. M e m . 216: 65 pp.

25 Alderman, A.R. and Von der Borch, C.C.. 1960. Occurrence of hydromagnesite in sediments in South Australia. Nature , 188: 931.

26 Alderman, A.R. and Von der Borch, C.C.. 1961. Occurrence of magnesite - dolomite sediments in South Australia. Nature , 192: 861.

27 Aleksandrova. G.I.. 1964. Stromatolites from Upper Permian deposits of the Aktyubin-Ural region, Tr . Molodykh Uch. Sarat. Ins t . , V y p . Geo1.-Geogr. , Sarat. . 1964: 87-88 (in Russian).

28 Alexander, M.. 1965. Introduct ion t o Soil Microbio logy . Wiley New York, N.Y.. 472 PP. 29 Alexandrov, E.A.. 1973. The Precambrian banded iron-formations of the Soviet Union. Econ .

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41 Andrews. J.T.. Guennel, G.K., Wray, J.L. and Ives, J.D., 1972. An Early Tertiary outcrop in northsentral Baffin Island, N.W.T., Canada: environment and significance. Can. J. Earth Sci..

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43 Annels. A.E.. 1974. Some aspects of the Zambian Copperbelt and their genetic significance. In: P. Bartholomd (Editor), Gisements stratiformes e t provinces cupr i f i res . Socidte' Gdologique de Belgique. LiAge. pp. 235-254.

44 Anonymous, 1916. Explorations and Field-Work o f the Smithsonian Insti tution in 191 5 . Smi thson. Misc. Coll . . 66(3): 119 pp.

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62 Awramik. S.M.. 1973a. Environmental and biological controls on stromatolite morphology: stromatolites of the Gunflint Iron Formation. In: Sympos ium o n Environmental Biogeo- chemis try , 1 9 7 3 , Logan. U tah , pp. 3 4 (abstracts).

63 Awramik. S.M., 1973b. The Origin and Evolut ion o f Stromatolites w i th Special Reference t o the Gunflint Iron Formation. Thesis, Harvard University, Cambridge. Mass., 262 pp.

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ADDITIONAL REFERENCES

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2030

2031

2032

2033

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Brock, T.D.. 1975a. Salinity and the ecology of Dunaliella from Great Salt Lake. J. Gen. Micro- biol.. 89: 285-292. Brock. T.D.. 1975b. Effect of water potential o n a Microcoleus (Cyanophyceae) from a desert

Brock, T.D., 1975c. The effect of water potential on photosynthesis in whole lichens and in their liberated algal components. Planta. 124: 13-23. Cardosa, J., Brooks, P., Eglinton, G., Goodfellow, R., Maxwell, J.R. and Philp, R.P., 1975. Lipids in recently-deposited algal mats a t Laguna Mormona. Baja California. Environmental Biogeochemistry Conf . . Proc., 2nd, Burlington, Canada, 1975 . in press. Doemel. W.N. and Brock. T.D., 1976. Vertical distribution of sulfur species in benthic algal mats. Limnol. Oceanogr., 21, in press. Eichmann, R. and Schidlowski. M., 1975. Isotopic fractionation between coexisting organic carbon-carbonate pairs in Precambrian sediments. Geochim. Cosmochim. Acta. 39: 585-595. Gebelein. C., 1975. Holocene sedimentation and Stratigraphy, Southwest Andros Island, Bahamas. IXth Int. Congr. Sedimentol. (I .A.S.)v Nice, 1975 , Theme V, pp. 191-197. Haslett, P.G., 1975. The Woodendinna Dolomite and Wirrapowie Limestone - two new Lower Cambrian formations. Flinders Ranges, South Australia. Trans. R . SOC. S . Aust . . 99(4):

Kinsman, D.J.J. and Patterson, R.J.. 1976. Hydrologic framework of a coastal sabkha in the Persian Gulf. In press. Mosser. J.L. and Brock. T.D., 1976. Temperature optima for algae inhabiting cold mountain streams. Arct . Alp. Res.. 8, in press. Philp, R.P. and Calvin, M., 1975. Kerogen structures in recentlydeposited algal mats at Laguna Mormona. Baja California: A model system for the determination of kerogen structures in ancient sediments. Enuironmental Biogeochemistry Conf . , Proc.. Znd, Burlington, Canada, 19 75 , in press. Sarjeant, W.A.S.. 1975. Plant trace fossils. 1n:R.W. Frey (Editor), The Study of Trace Fossils. Springer-Verlag. Berlin. Sweet, I.P.. Mendum. J.R., Bultitude, R.J. and Morgan, C.M.. 1974. The geology of the southern Victoria River region, Northern Temtory. Aust. Bur. Min. Resour.. Geol. Geophys., Rep. 167: 143 pp. Trichet, J., 1967. Essai d’explication d u ddp& d’aragonite sur des substrats organiques. C.R. Acad. Sci. Paris, 265: 1464-1467. Tsao Rui-chi and Liang Yu-zhou. 1974. On the classification and correlation of the Sinian SYS- tem, based on a study of algae and stromatolites. Mem. Nanking Inst. Geol. Palaeontol,, Acad. Sin., no. 5: 1-16 (in Chinese). Western Australia, Department of Environmental Protection, 1975. Conservation o f Hamelin Pool, Western Australia. Western Australia, Dept. Environmental Protection, Rcport, 21 pp.

crust. J. Phycol., 11: 316-320.

211-219.

INDEX*

Abenab Formation, 364, 367 Abiophoric, 687 Abu Dhabi, 122,125, 212

, see Persian Gulf Acaciella, 33, 40, 369, 694 - angepena, 255

- australica, 49, 191 -- , organic geochemistry, 174-176 Acetabularia, 124, 265, 396, 403, 405,

Adelaidean, 350 Adelaide gzosyncline, 565 - Supergroup, 693 Aimchan Group, 341,615 - Series, 366 Akademikerbreen Group, 341 Akaitcho Formation, 602 Akinetes, 129, 136 Akitkan Series, 366 Aladin Formation, 341 Aladyin suite, 366 Alberta, 561 Alcheringa, 40, 694 - narrina, 191, 362 Aldan Anteclise, 361 - River, 613 - Shield, 361, 613, 614, 615, 621 Aldania, 39, 362, 368, 369, 694 - sibirica, 361 Alenia, 603, 604, 605 Algal mats, distribution, 124-125 -- , internal growth, 204-205 -- , morphology, 113, 115-126, 146,

Alignment of grains in stromatolites,

Alps, 559 Alternating lamination, 194 Alternella hyperboreica, 255 Altyn Limestone, 585-597

--

-augusta, 256

407

158

217,218, 239

Alunite, 490 Amadeus Basin, 362 Amelia Dolomite, 362 Amino acids in stromatolites, 163, 164,

Ammatoidea, 140 Ammonia beccarii, 418, 419 Ammonoids, 551, 553 Anabaena, 128, 136,137.138 - cylindrica, 184 A na baenopsis, 136 Anabaria, 39, 327, 347, 361, 362,.368,

373,375,694 - radialis, 49, 340 Anabar massif, 325,328,329,341 - Region, 360, 361, 366 Anacystis cyanea, 171,172 - montana, 171,172 Andros Island, 197, 198, 199, 201, 203,

204, 217, 218, 219, 221, 222, 225, 237,245, 246,248,385,386,447- 477

169-170,182-183

Anhydrite, 430, 432, 506, 670 Animikiea septata, 316, 317, 318 Animikie Group, 365 Apache Group, 364,367 Aphanizomenon, 128, 136,137,138 Aphanocapsa, 128,131, 132 - thermalis, 143 Aphanothece, 128, 130, 131,422 Aphebian, 348,353,372-380,523,693,

694 - see also Proterozoic, Early Appia, 352 Aragonite, 56, 61, 65, 66, 68, 77, 82, 83,

123,182,183, 214-215,217,219, 227, 229, 231, 232, 235, 237, 247, 263, 297, 405, 411, 413, 417,418, 419, 435, 437, 438, 439, 440,441, 443, 444, 445, 460, 463, 464, 465, 467,469,471,472,477

* Numbers in bold type refer either t o illustrations or t o definitions.

774 INDEX

Aravalli, 653, 656 Arbuckle Mountains, 217 Archaeorestis, 316, 317 A rchaeosphaeroides barbertonensis, 508 Archaeozoon, 15, 48, 49 - acadiense, 15, 49 Archean, 150, 153, 156, 160, 364, 500,

501, 502, 503, 508, 509, 510, 515, 654

Arthrobacter, 667 Arumbera Sandstone, 367 Arymas Formation, 341 - suite, 366 Asha Group, 341 Ashin suite, 366 Assabet el Hassiane Group, 517, 518 Atar Group, 367, 517, 518, Athapuscow Aulacogen, 365, 599-611 Atmosphere, Precambrian, 500, 501,

502, 503, 508, 509, 512, 639, 641, 660, 671

Auloporq 551 Autobioturbation, 247, 248, 249 Avzyan suite, 366 Ay suite, 366

Bacteria, 3, 24, 25, 26, 27, 28, 75, 83, 88, 99, 110, 114, 115, 116, 119, 123, 126, 127, 151, 153, 154, 156, 157, 159, 162, 165, 166, 176, 180, 183, 189, 194, 198, 208, 222, 231, 233, 235, 239, 241, 242, 247, 253, 315, 492, 493, 502, 503, 505, 507, 508, 509, 513, 515, 547, 558, 562, 655, 656

- environmental limits, 141-148 - in stromatolites, 273-310 -, iron precipitation, 668 -, Mn precipitation, 3, 667, 668 -, organic chemical composition,

170-173, 179-180, 181, 185-187 -, sulphate-reducing, 669-670 Bacterial stromatolites, 2, 3, 397, 507,

513 , see also Bacteria --

Bacteriochlorophyll, 24-26, 29, 276,

Baffin Bay, 166 Baffin Island, 653, 660 Bahamas, 117, 121, 153, 212, 253,

-, see also Andros Island

277

256, 382, 384, 385, 413

Bahia, 653 Baicalia, 37, 39, 42, 325, 327, 329,

330, 331, 332, 333, 335, 347, 351, 352, 361, 362, 363, 364, 365, 368, 369, 373, 374, 375, 376, 505, 608, 609, 625, 627, 653, 656, 694

- anastomosq 518, 519 - baicalica, 42, 340, 614 - burrq 49, 191, 256, 362, 363

, organic chemistry, 176 - capricorniq 256, 362 - ingilensis, 340, 614, 626, 629 - lacerq 253, 340, 614, 626, 629 - maicq 340, 363, 614, 629 - mauritanica, 252, 253, 518, 519 - minuta, 340 -, organic chemistry, 175, 176, 178 -prima, 340 - rarq 258, 340, 364 - uncq 340 Baja California, 190 Bakal suite, 366 Bambui Group, 365, 653, 656 Banded iron formations, see Iron

- microstructure, 253, 279, 285, 687 Bangemall Group, 362, 367 Barbados, 57, 67, 70 Barite, 490, 528, 588, 589, 590 Basisphaera, 40, 362, 369,' 694 - irregularis, 49 Bathurst aulacogen, 599, 600 - Inlet, 523 Batrachospermum haematites, 120 Bear Province, 523 Beck Spring Dolomite, 157, 161, 182,

Beggiatoq 119 Belcher Group, 372, 373, 374, 375,

376, 377, 378, 379, 509 - Islands, 117, 156, 157, 311 Belt Supergroup, 187, 364, 365, 367,

Bermuda, 153, 327, 381, 383, 413,

Bezymen suite, 366 Bhander Limestone, 364 - Series, 367 Bibra Formation, 393 Bijawar formations, 651, 656 Bioherm, 687 - series, 41-42, 43, 334, 336, 355

- -

formations

510

585-597, 650, 656

676

INDEX 775

Biological methods, 21-29 Biophoric, 687 Biostrome, 687 Bioturbation, 70, 567 -, see also Burrowing Biryanian horizon, 350 Bitter Springs Formation, 157, 161,

165, 167, 168, 169, 170, 172, 182, 191, 311, 362, 363, 367, 369, 504, 510

Biwabik Iron Formation, 188, 364, 653, 656, 677, 678

Black Reef Formation, 654 - - Quartzite, 636, 637 - Sea, 664 Blake Plateau, 666 Blister mat, 264, 265, 267 Block diagram, 9-10 Blue-green alga, see Cyanophyte Boetsap River, 156, 188 Bonneterre Formation, 656 Boring cyanophytes, 132, 134, 139,

150, 399 - of carbonates, 217, 221 Borings, 70 Boxonia, 32, 39, 40, 48, 347, 361,

- allahjunica, 340 - divertata, 255 - gracilis, 255, 364 - grumulosa, 258, 340, 621, 630, 631 - ingilica, 258, 621, 631 - lissq 49, 258, 340 Brachiopods, 553, 555, 557 Brachytrichia, 128, 139 Branching, 583-584, 594-596, 603,

687, 688 - of columnar stromatolites, cause of,

158 - styles, 6 Brazil, 365, 653 Bridge, 690 Bulawayan Group, 156, 167, 187 - System, 169 Bulawayo, 150, 156, 160, 184, 500,

- Dolomite, 364 Bunyeroo Gorge, 191 Burdur suite, 366 Burovaya Formation, 341 - suite, 329 Burovoy suite, 366

362, 368, 373, 375, 694

508

Burra Group, 191, 362, 363, 367 Burrowing, 215, 425, 453, 460 - metazoans, 381, 382, 385, 386,

507, 512, 513, 514 -, see also Bioturbation Bushimay Group, 367 Bushveld Complex, 635 Burzyan Group, 341 - Series, 361, 366

Calcite, 56; 59, 61, 63, 64, 66, 69, 77, 82, 83, 165, 182, 186, 189, 199, 201, 205, 208, 231, 247, 255, 413, 417, 418, 419, 455, 480, 482, 483, 484, 485

-, high-Mg, 56, 61, 63, 66, 183, 217, 417, 465, 467, 469

465, 469, 476, 484 -, magnesian, 417, 419, 451, 453;

463, 465, 467, 472, 475, 477 Calcrete, 55-70, 56, 417 Calc Zone of Pithoragarh, 364 Caliche, 55 Callanna Beds, 367 Calothrix, 116, 128, 137, 138, 143,

- coriaceq 288, 303, 304, 495, 496 - crustacea, 121 - parietina, 483 Calyptra, 136 Cambrian, 154, 158, 159, 167, 180,

255, 259, 347, 353, 354, 355, 369, 501, 565-584, 651, 652, 656, 694

-, low-Mg, 56, 201, 207, 449, 455,

207, 290, 291, 305, 494

Campbell Rand Dolomite, 637 Canning Basin, 384, 514, 543-563,

Cape Sable, 383, 385, 386, 387 Carbla Oolite, 393 - Point, 124, 215, 218, 219, 221,

2’69 Carbonate crystal form, 451, 453,

454, 455, 463, 467, 483, 484, -485

- - - in algal mats, 207, 208, 209, 249

- precipitation by algae and cyanophytes, 66, 83, 115-116, 117, 120, 182, 199, 200, 201, 202, 204, 205-208, 212-213, 221, 439, 445, 451, 475, 485, 580, 581, 582, 589-590, 596, 611

657

776 INDEX

--- bacteria, 75, 83, 123, 231,

Carboniferous, 158, 559, 657 Carbon isotopes, 156, 163, 185-188,

189, 455, 473, 501, 508, 643 Cardisoma, '4 59 Carlsbad Caverns, 75, 78, 80, 81, 82 Carnic Alps, 559 Castle Roads, 327, 383 Catagraphia, 256, 257, 258, 630, 690,

Caulerpa, 116, 151 - fastigiata, 122 Cavan Limestone, 167, 175, 191 Cave coral, see Speleothems Cave popcorn, see Speleothems Cave pearls, see Speleothems Caves, 73-86, 123 Cenozoic, 513, 515 Cerion, 459 Chalcedonic quartz, 77 Chalcedony, 506 Chamaesiphon, 128, 132, 133 Chamaesiphonaceae, 128 Chamaesiphonales, 128, 130, 132, 133,

Chantmnsia incrustans, 211 Cham, 482, 483, 485 Char Group, 367, 517, 518 Charlton Bay Formation, 602 Chemocline, 690 Chemotaxis, 507, 508, 509 Chert, see Silica Chikhan Formation, 341 Chironomids, 205, 222, 224, 226, 227 Chlorella, 183 - pyrenoidosa, 171, 172 - vulgarus, 671 Chlorellopsis coloniata, 158, 180 Chlorococcum infusionum, 158 Chloroflexus, 88, 116, 119, 198, 275,

- aurantiacus, 288, 297, 299, 300,

Chlorogloea, 123, 131, 132 Chlorophyll, 24-26, 29 -, see also Pigments, photosynthetic Choctawhatchee Bay, 166 Christie Bay Group, 602 Chroococcaceae, 128, 131, 156, 158 Chroococcales, 128, 130-132 Chroococcidiopsis, 128, 133, 134

233, 235, 239, 451

69 1

134

276, 281, 301

493

Chroococcus, 128, 131, 132 Chrysogloea, 151 Churochnaya Group, 341 Cladophom, 151 Classification of stromatolites, 31-43,

339, 355-356, 359-360, 405, 407

- - - , columnar, 39-40 - - - , columnar-nodular, 39 --- , columnar-stratiform, 38

, nodular, 38-39 , stratiform, 38 , subtype, 41 , supergroup, 41

Clastidium, 128, 132, 133 Clones, 119, 690 Clot, 690 Coccogoneae, 128, 129, 130-134,

Coelosphaerium, 128, 130, 131 Collenia, 33, 38, 333 - baicalica, 364 - buriatica, 364 - columnark, 364 - kussiensis, 364 - symmetrica, 364 - undosa, 648, 656 Collenia, 519, 650, 651, 652, 653,

- symmetrica, 653, 656 Colleniella, 38, 347, 694 - singularis, 340, 621, 631 Colloform mat, 264, 265-266 Collumnncollenia, 33, 34, 35, 37, 39,

Collumnaefacta, 38, 347, 694 Colonnella, 33, 34, 49, 309, 327,

329, 330, 347, 605, 626, 627, 694

--- --- --- - - -

154, 156, 157

654, 656

694

- cormosa, 340 - discreta, 340 . - frequens, 340 - kylachii 340, 614 - lineata, 340 - ulakia, 340, 614, 629 Colorado, 74, 79, 84, 85 Column, 690 Compactocollenia, 40, 694 Computer methods, 15, 17, 20 - processing, 45-53 Conchocelis, 154 Conestoga Creek, 123

INDEX 777

Conocolleniq 38 Conodonts, 551, 557, 558, 561 Conophyton, 5, 12, 34-36, 37, 38,

39, 40, 41, 42, 52, 159, 191, 253, 258, 329, 330, 331, 333, 338, 347, 351, 352, 362, 363, 365, 369, 376, 377, 4518, 452, 453, 517, 518, 519, 523-534, 604, 605, 606, 608, 611, 625, 627, 651, 656, 694

-, axial zone, see -, crestal zone -, crestal line, 690 -, crestal zone, 12, 34, 35, 41, 285,

286, 293, 295, 307, 308, 327, 527, 528, 532, 690

- cylindricum, 340, 361, 364, 365, 368, 626, 629

- garganicum, 35, 253, 310, 340, 361, 362, 368, 614

- gaubitza, 361 - lituum, 309, 340, 361, 614, 626,

629 - metulum, 340, 361, 368, 614, 629 - miloradovici, 340, 361 -, organic chemistry, 175, 176, 178 - ressoti, 252, 253, 364, 518, 519 - weedii, 273-310, 494 Coorong Lagoon, 158

area, 413-420 Copper in stromatolites, 645-651,

Coppermine River Group, 524, 525 Corals, 543, 545, 551, 553, 556, 563 Cornice, 690 Coronation Geosyncline, 374 - Gulf, 523, 524 - sills, 524, 525 Correlation, intrabasinal, 331, 334 Cowles Lake Formation, 601 Coxiella confusq 417, 419 Cretaceous, 558 Crinoids, 551, 553, 556, 557, 563 Cryptalgal, 690 - laminite, 194 - structures, 194 Cryptalgalaminate, 690 Cryptozoon, 33, 38, 49 Crystalline laminated fabric, 208, 209,

Crystal Springs Formation, 364 Culturing, 275, 277 Cumulate stromatolite, 6 Cyanidium caldarium, 145, 186

- -

656

210, 211

Cyanophanon, 128, 132, 133 Cyanophyte, 64, 66, 73, 690 -, culturing, 24 -, environmental limits, 141-148 -, organic chemical composition, 170-

173, 179-180, 181, 182-187 -, taxonomy, 24, 127-140, 275,

300-301 Cyanostylon, 116, 117, 118, 128, 131,

132, 151 Cyanothrix; 130

- sedimentation, 600-605, 624 Cyclothemic lamination, 194 Cylindrospermum, 128, 136, 137, 138 Cymbellq 116, 117, 151 Cymodoceq 396

’ Cyclic deposition, 477

Dadykta Formation, 341 Dadyktin suite, 366 Dampier Formation, 393 Dashkin suite, 366 Debengda Formation 341 - suite, 324 Debengdin suite, 366 Decomposition, 147, 149, 155, 157,

- of microbial mats, 28, 115, 122-124

Deep-sea environments, 507

Dema Formation, 341 Deoban Limestone, 364 Depot Creek, 191 Derevnin suite, 366 Derevnya Formation, 341 - suite, 330, 332, 333 Dermocarpa, 128, 133, 134 Dermocarpaceae, 128, 134 Desulfotomaculum, 669 Desulfovibrio, 669 - desulfuricans, 170, 185 Devonian, 167, 182, 249, 384, 514,

Dgerbia, 38, 694 Diagenesis, 3, 56, 64, 65, 68, 70, 149,

217, 244, 247, 251, 256, 258, 311, 313, 354, 465, 473, 477, 527, 565, 590

182

173, 174, 176, 178, 179, 185, 188

160, 166, 178, 218, 313

organic matter, 453

sediments, 186

--

--

543-563, 559, 561, 657

-, effect of amino acids on carbonates,

- of organic matter, 168-169, 170,

778 INDEX

Diapir, 565, 566-567, 569, 571 Diatom, 110, 116, 117, 149, 151, 154,

224, 381, 444, 450, 483, 491, 496 - mats, 242-244 Dichothrix, 128, 137, 138 - bornetiana, 237 Discovery Point, 313 Dismal Lakes Group, 523, 524, 527-

Djura Formation, 341 Dolomite, 56, 69, 77, 82, 86, 183,

186, 189, 323, 405, 411, 413, 417, 418, 419, 472, 477, 504. 506, 587

532

- Series, 364, 367 Dolomitization, 165, 166, 189, 216,

Dook Creek Dolomite, 362 Douik Group, 517 Duhamel Formation, 602 Duitschland Formation, 637 Dunaliella salina, 144, 145 Durnomys suite, 366 Dzhur suite, 366

Earthquakes, 107 Ecotone, 225, 229, 233 Ediacara fauna, 155, 362 Egan Formation, 367 Elgee Siltstone, 362, 367 Elimberrie Spring, 544, 548, 549,

559-561, 563 El Mreiti Group, 517 Elongate stromatolites, 265, 266, 267,

268, 270, 305, 323, 324, 335, 382, 400, 401-403, 551, 552, 553, 569, 570, 575, 592, 593, 594, 596, 597, 603, 605, 606, 611, 638, 639, 640

217, 430, 431, 475, 639-643

Endolithic, 690 - microorganisms, 123, 124 Endosporangium, 134 Endospores, 129, 130, 133, 134 Ennin suite, 366 Enteromorpha, 237 Enterosphaeroides amplus, 315 Entophysalidaceae, 128, 131, 157 Entophysalis, 116, 123, 228, 131, 132,

- granulosa, 124 - major, 117, 118, 124, 125, 235,

- riuularis, 483

151, 157, 227, 241, 242

264, 405, 424, 450

Envelopes, around cyanophytes, 129 -, see also Sheath, Mucilage Eoastrion bifurcatum, 313 - simplex, 313 Eocene, 156, 158, 180, 535-541 Eo m icrh ys trid iu m barghoo rn i, 3 1 8 Eomycetopis, 154 Eosphaera tyleri, 316, 317 Epilimnion, 690 Epworth Group, 372, 374, 523 Etching, 17 Etina Formation, 191 Eucapsiphora, 39, 369, 694 -paradisa, 362 Eucapsis, 128, 131, 132 Eucaryote, 690 Eucaryotic algae, 115, 116, 221, 122,

124, 127, 140, 151, 154, 156, 157, 158, 159, 162, 165, 201, 211, 237, 381, 385, 405, 483, 512, 514

- _ , environmental limits, 141-148 -- , organic chemical composition,

170-1 73, 179-1 80, 183-187 Euphotic zone, 690 - -, see aJso Photic zone Everglades, 447, 451, 456, 457, 459,

Exosporangium, 132 Exospores, 129, 132, 133

, advent of, 182 - -

475

Fabric, 194-249, 263-267 Facing, 7, 8 False branching (cyanophytes), 136 Famennian, 544, 545, 557, 561 Fatty acids in stromatolites, 163, 164,

Faure Sill, 391, 394, 395, 409 Fawn Limestone, 364 Fenestrate fabrics, 64, 68, 70, 92, 96,

104, 105, 107, 194, 205, 208, 217, 224, 225, 226, 227, 230, 231, 233, 235, 237, 264, 265, 266, 267, 386, 388, 403, 408, 410, 411, 437-439, 471, 484, 485, 486, 487, 545, 546, 547, 562, 563, 603, 638

169, 177, 179-180, 189

Ferrobacillus, 668 Ferruginous stromatolites, see Iron in

stromatolites Field and laboratory methods, 5-13 Fig Tree Group, 156, 170, 311, 504,

50 8

INDEX 779

Filament, 129 Film mat, 264 - microstructure, 253, 259 Fischerella, 128, 139 Fission, 129 Flinders Ranges, 565 Flint Point, 313 Florida, 57, 63, 67, 70, 74, 212, 383,

Flowstone, see Speleothems Fluorite in stromatolites, 655, 656 Fluviatile environemnts, 120, 122, 123 Foraminiferids, 384, 393, 396, 405,

411, 418, 460, 461, 464, 551 Fortescue Group, 191, 364, 367 Fragum hamelini, 393, 396, 399, 405 Frasnian, 544, 557 Frazer's Hog Cay, 384 Fresh Creek, 199, 201, 241, 242, 243,

Freshwater environments, 132, 134,

385, 447, 449

244

136, 138, 140, 146, 199, 202, 203, 207, 211, 222, 231, 232, 244, 245, 246, 255, 258, 447-477, 479-487, 502, 505, 507, 513, 535-541, 599, 608-611

Frustration Bay, 313 Frutexites, 547, 554, 555, 556, 558 Fungi, 66, 75, 115, 123, 154, 165,

-, environmental limits, 141-148 -, Mn precipitation, 667 -, organic chemical composition, 170-

Fyn Sd Dolomite, 365

Gaia, 39 Gallionella ferruginea, 668 Gangolihat Dolomites, 364, 653, 656 Gardnerula, 116, 128, 137, 138 - corymbosa, 120, 124 Garwoodia, 244 Gastropods, 409, 417, 419, 439, 441,

Geitleria, 116, 140 - calcarea, 121 Gel, 129 Gelatinous mat, 264, 265, 267 Geochemistry, 3, 635-643 Geopetal fabrics, 553, 555, 557 Georginia, 34, 40, 41, 369, 694 Geotropism, 547, 551, 557

166, 189

173, 179-180, 183-184

459, 475, 483, 538

Geyserite, 87-112, 180, 676, 677 Ghaap Group, 636, 637, 640 Gibraltar Formation, 602 Girvanella, 158, 180, 227, 231, 255,

439, 453, 457, 547, 554, 555, 556, 558

Glacial sediments, 362, 363, 365, 367, 512, 518, 519, 520

Glacier National Park, 585-597 Gladstone Embayment, 392 Glebulella, 40 Gliding, 690 -, see Movement Gloeocapsa, 128, 131, 132, 140, 450 Gloeothece, 128, 130, 131, 132 Gloeotrichia, 128, 137, 138 Gold in stromatolites, 655, 656, 659 Gomphonema, 117 Gomphosphaeria, 128, 130, 131, 132 Gonam Formation, 341 - suite, 615 Gongrosira, 116, 151 - incrustans, 121 Gongylina, 38, 41, 323, 347, 694 - differenciata, 340, 360, 614 -mixta, 340 - nodulosa, 340,621, 631 -zonata, 340 Goniatites, 557 Gorbilok suite, 366 Goulburn Group, 372 Graphic representation, 15-20 Grazing and burrowing of microbial mats,

115, 116, 146, 261, 289, 308, 381, 382, 384, 385, 409, 418, 429, 507, 512, 513, 514

-, see also Burrowing Great Bear Lake, 523, 524 - Estuarine Series, 188 -Salt Lake, 413,435-445 - Slave Lake, 500, 502, 678 ---.Supergroup, 372, 374 Green Lake, 123,231,232, 479-487 Greenland, 365 Green River Formation, 156, 158, 180, "'479, 486, 535-541

Griquatown Jasper, 637 Growth rates, 27, 386, 407,411, 557,

674,675 Grumous, 690 Gruneria, 38, 347, 363, 368, 694 - biwabikia, 312, 314, 318, 319, 364

780 INDEX

Gulf of Aqaba, 182 -- Batabano, 166 Gunflint Group, 372 -Lake, 313 Gunflintia grqndis, 316, 317, 318 -minuta, 315,316, 317,318, 319,320 Gunflint Iron Formation, 111, 156, 157,

161,167,168, 169-170, 173,180, 181,182,188,311-320,364,367, 509, 510, 653, 654, 656, 669, 677, 679,682,684

Gymnosolen, 32, 33, 34, 35, 37, 39, 335, 347, 350, 351, 352, 361, 362, 365, 368, 373, 374, 375, 605, 652, 654, 656,693,694

-altus, 340 - asymmetricus, 340 - directus, 518, 521 - furcatus, 49, 340 - hankii, 518, 519 - levis, 340 - ramsayi, 41-42,191,256,340,363,

518, 519 - tungusicus, 340 Gypsum, 427, 429, 430, 432, 480, 504,

506, 531,670

Hadrynian, 523 Halite, 506, 531, 539, 623, 624 Hamelin Pool, see Shark Bay Hamersley Group, 367, 643 Hapalosiphon, 128, 139 Harbor Island, 166, 171, 172 Hawker Group, 191,367 Hearne Formation, 606 Helena Formation, 650, 651 Helikian, 365, 523 Heliotropism, 690 -, see also Inclined columns Heterocysts, 129, 134, 135, 136, 137,

138,139,143,173 Heterotrophism, 690 Himachal Pradesh, 364 Hinde Dolomite, 362 Homeostasis, 690 Homocystous filamentous cyanophytes,

Homoeothrix, 140 Hormathonema, 128,131,132 - violaceo-nigrum, 123, 124, 264 Hormogoneae, 128, 129, 134-140, 154,

Hormogonia, 129

134

156,157

Hornby Bay Group, 523, 524 - Channel Formation, 602 Hot Springs, 21, 25, 87-112, 180, 186,

Hoyoux River, 207, 208, 209, 211, 232,

Huntite, 77, 82, 86 Huntsman Limestone, 156 Huroniospora, 313, 315, 316, 317, 318,

Hurwitz Group, 372 Hutchison Embayment, 262, 265, 269,

391,392,395,396 Hyalophane, 589, 590 Hydrocarbons in stromatolites, 163, 164,

Hydrococcus, 128, 133, 134 Hydrocoleum, 128, 135, 136 Hydromagnesite, 77, 83, 86, 413, 417,

Hyella, 128, 133, 134 Hyphomicrobium, 667, 668

Iceland, 288 Ignikan Formation, 341 -suite, 614, 615, 622, 627, 629, 632 Ilicta, 39, 347 Ilyin suite, 366 Image-analysis, 45 Inclined columns, 267, 269, 270, 304,

305, 324, 402, 403, 532, 533, 592, 593, 594,605

187,273-310,489-498, 507

238

319,320

168-169,174,175,177) 189

418,419

Individual, 690 International Code of Botanical Nomen-

Interspace, 691 Intertidal environment, colonization by

microorganisms, 143, 144-145, 153, 156

clature, 32, 43

Intrabasinal correlation, 379, 5 17-522 Inzer Formation 341 - suite, 366 Inzeria, 39, 48, 49, 327, 335, 347, 351,

352, 361, 362, 363, 368, 313, 376, 318, 505,694

- confragosa, 340,350,614,629 - djejimi 340 - intia, 256 - nimbifera, 340 - nyfrieslandica, 253 - tiornust 35, 42, 340, 350, 362, 614,

- variusata, 340 629

INDEX 781

Ireland, 559 Iron Formation, 87, 110-111, 501-502,

504, 508, 509,635443,653,654, 659,669

-- , see also Biwabik Iron Formation, Gunflint Iron Formation

- in stromatolites, 122, 126, 132, 209- 211, 555,558-559,561,562, 563, 635-643,653-654,656,668-669

- River Formation, 187 Irregularia, 35, 38, 41, 347, 614, 627,

628,629,694 Irus irus, 403 Isocystis, 299, 300 Isometric diagrams, production of,

Isotopic dating, 340, 366, 367, 368,

-- in the U.S.S.R., 348, 352 -- of glacial rocks, 365

15-20

312

Jacutophyton, 34-36, 31, 40, 41, 42, 48, 49, 330, 331, 332, 333, 347, 351, 353, 361, 362, 363, 365, 369, 313, 374, 376, 377, 518, 519, 608, 609, 614,625,631,694

627,632 -cycle, 324, 329, 330-335, 625, 626,

- multiforme, 340 - ramosum, 340 Jay Creek, 255 -- Limestone, 255 Johannesbaptistia, 128, 130, 131, 132 -pellucida, 200, 201, 450 Johnnie Formation, 364, 367 Joulters Cays, 382, 384 Ju . . ., see also Yu . . . Judoma Formation, 341 Jurassic, 188,558, 559,657 Jurusania, 39, 347, 351, 361, 362, 363,

- burrensis, 256 - cylindrica, 42, 256, 340, 350 - derbalensis, 518, 520 - judomica, 256, 340, 621, 630, 631 - lissa, 518, 520 - nisvensis, 256, 518, 520 - sibirica, 340, 621 - tumuldurica, 340

364,313,375,694

Kahochella Group, 602 Kaimur Series, 367

Kakabeka Falls, 313 Kakabekia umbellata, 313 Kameni Islands, 668 Kandyk suite, 366, 615, Kaniapiskau Supergroup, 372, 374 Karatau, 360 Karatav Group, 341 - Series, 366 Karatavian Series, 352 Kartochka Formatior, 340 Kartochki suite, 366 Kasaia, 39, 694 Kasegalik Formation, 117 Katanga, 648,656 Katav Formation, 341 -suite, 366 Katauia, 39, 352, 361, 362, 368, 374,

Katavian horizon, 350, 351, 352 Katelysia, 417, 419 Katernia, 39, 347, 694 - africana, 364, 368 Kentucky, 74, 79, 81, 84, 85 Kerogen in stromatolites, 167, 168, 169,

170,177,180,181,183-185,187, 188-189,190

694

Kerpyl Group, 341, 615 Kessyusin suite, 366 Key West, 121 Khaipakh Formation, 341 Kharaulakh Highlands, 361 Khatyspyt suite, 366 Khaypakh suite, 366 Khorbusuonka Group, 341 Kimberley Block, 543 Kingston Peak Formation, 365 Kleberg Point Lagoon, 166 Kluziai Formation, 602 Knob Lake Group, 374 Korea, 350, 351 Kotuikania, 39, 40, 49, 328, 347, 361,

362,694 - torulosa, 49, 340 Kotuikan River, 34 Kotuykan suite, 325, 329, 360, 366 Krasnogor suite, 366 Krasnya Gora Formation, 340 Kukhan Formation, 341 Kulpara, 191 Kulparia, 40, 362, 369, 694 - kulparensis, 191 Kupferschiefer, 6 7 0

782 INDEX

Kuruman Iron Formation, 637 Kussiella, 34, 39, 40, 41, 48, 327, 329,

335, 347, 362, 364, 368, 373, 374, 614,694

- kussiensis, 41-42, 49, 325, 329, 340, 360,361,614

-superiors, 312, 314, 318, 319 - vittata, 340, 360 Kutingda Formation, 341 Kyrtuthrix, 128, 131 Kyutingdin suite, 361, 366

Labaztakh suite, 366 Laboratory methods, 9-12

* Labrador Geosyncline, 374 Lacustrine environments, 435-445,

-- , see also Freshwater environments Laguna Atacosa, 171,172 - Madre, 247 Lakes, see Lacustrine environments Lake Gosiute, 535, 536 - Krasnoe, 668,669 -Superior, 157, 313 Lakhanda group, 341,351,615 Lakhandian horizon, 350, 351, 352 Lakhandin suite, 363, 366 Lamina, 691 - coefficient of thickening in

-convexity, 11, 12, 51-52 - morphometry, 48, 51-52 -shape, 11, 115, 321, 322, 323, 327,

-- , in taxonomy, 38-40 -, temporal significance, 674-680, 685

-thickness, 12, 52-53, 96, 103, 111, 285-286,298-299,679,680,686

Laminated fabric, 194 Lamination, production by unicellular

cyanophytes, 156 Laminoid boundstone fabric, 194 - fabric, 194 - fenestral fabric, 194 Laterally linked stromatolite, 691 Layered fabrics, origin, 197-227 Lead and zinc in stromatolites, 559,

651-653,656 Lena River, 35 Lenia, 40, 347, 694 Lennard Shelf, 543

479-487

Conophyton, 12

582,587-590

- -- , , see also Growth rate

Lenticular laminoid fabric, 194 Lesser Himalayas, 364 - Karatau, 361 Lewis Overthrust, 585 Lichen, 145 Light intensity for photosynthesis in

-- , effect on cyanophytes, 204, 241,

- in morphogenesis, 301-309 -penetration in algal mats, 23, 114 Linella, 39, 347, 362, 363, 368, 313,

-avis, 361, 362 - simica, 258, 340,361,369, 621 - ukkq 42, 258,340,361,364,369 Linocolleniu, 39, 347, 694 - angaricq 309 Linok Formation, 341 - suite, 366 Lithophyllum, 151 Little Dal Formation, 365 Longview Formation, 652, 653, 656 Lukoshi Complex, 654,656 Lunzer Untersee, 204 Lyngbya, 116, 128, 136, 136, 151, 153,

211, 237,247,429 - aerugineo-coeruleq 120 - aestuarii, 114, 118, 119, 120, 124,

- conferuoides, 114, 120, 171, 172 - lagerhaimii, 171, 172 - majuscula, 120 - semiplena, 120

cyanophytes, 145-146

247

376,378,694

126,171,172,264,425

Maastakh suite, 366 Mackenzie Mountains, 365 Macrolamina, 691 Madiganites, 40, 369, 694 - mawsoni, 264,255, 578 Madya Pradesh, 651 Magnesite, 77, 419 Maia River, 36 Malgina Formation, 341 -suite, 366, 615, 628 Malginellq 617 - malgica, 340, 352, 614, 628 - zipandica, 340, 614, 628 Malmani Dolomite, 635-643 Manganese in stromatolites, 209-21 1,

513, 558-559, 561,562, 563, 636-643,654,655,656

INDEX 783

-nodules, 3, 153, 209, 558, 561, 562, 654,655,656,658,659,660,666- 668

Manitounuk Group, 372 Mantle, see Selvage Margin structure, 7, 11 Marsh deposits, 447-477, 491, 496 Massive fabrics, 235-244 Mastigocladaceae, 128, 139 Mastigocladus, 128, 139 - laminosus, 143, 180 Mastigocoleus, 128, 139 Mauritania, 517-522 Maya depression, 615, 621, 622 - River, 361 - Series, 366 Mayak Group, 341 McArthur Basin, 670 - Group, 191 McLean Formation, 599, 600, 606, 608 McLeod Bay Formation, 602 McNamara Formation, 367 Meiosis, 160 Merismopedia, 128, 130, 131 Meristematic zones, 136 Merokhsan Formation, 341 Meromictic lake, 691 Mescal Limestone, 364 Mesozoic, 514, 515 Metallogenium, 667, 668 Metazoans in algal mats, 450, 459 Michigan Basin, 559 Microbial destruction of stromatolites,

Microchaete, 128, 137 Microcodium, 64 Microcoleus, 116, 128, 135, 145, 151,

211, 215, 239, 248, 450, 454, 455 - chthonoplastes, 114, 119, 125, 126,

171, 172, 264, 424 - corymbosus, 121 - tennerrimus, 124, 264, 405 Microcystis, 128, 131, 132 Microfossils, 11, 64, 70, 150, 154-158,

181, 311-320 Microorganisms, culturing, 24 -, taxonomy, 24 Microphytolites, 327, 338, 339, 348,

Microphytoliths, 691 Microspherulite, see Spherulite

122-124

622

Microstructure, 11, 70, 189-190, 194- 249, 251-259, 334, 335, 336, 339, 349, 350, 352, 354, 355, 356, 385, 437-445, 484-485, 509, 587-590, 628-632

-, in taxonomy, 38-42 -, interpreting, 159 -, morphometry, 48, 52-53 - of geyserite, 96-107 -, relation to distribution of organic

matter, 165 Microstylus, 40, 694 - perplexus, 360 Micro-unconformities, 8 Migoedikha Formation, 340 Minjaria, 33, 39, 347, 351, 352, 361, 362,

-procera, 255 - sakharica, 340, 614, 629 - uralica, 42, 340, 350 Mink Mountain, 11, 313 Minyar Formation, 341 - suite, 366 Minyarian horizon, 350 Miroyedikhin suite, 366 Mississippian, 651 Mississippi Valley, 651, 660, 670 Missoula Group, 587, 590 Missouri, 651, 652 Mistassini Group, 372 Mixolimnion, 691 Monimolimnion, 691 Mono Lake, 232, 241 Moonlight Valley Tillite, 367 Moonmilk, see Speleothems Moora Group, 389 Morocco, 650 Morphogenesis, 2, 261-310, 382, 487,

613, 625-628 Morphometrics, 45-53 Morphotype, 157 Moss, 115, 132, 479 Motility, see Movement Mount Bruce Supergroup, 362 - Herbert, 191 - Isa, 649, 656, 660, 661 -- Group, 649 Movement of microorganisms, 115, 116,

364, 368, 627, 653, 656, 694

119, 129, 136, 152, 160, 197-198, 201, 211, 275, 690

Mucilage, 129 -, organic chemistry, 184

784 INDEX

Mud mounds, 249 Mukhchan Formation, 341 Mukun Group, 341 Mundallio Creek, 191 Murky Formation, 599, 600, 608-611 Murrumbidgee Group, 191 Muskeg evaporites, 561 Muskox Complex, 524 Myxobacteria, 299, 300 Myxosarcina, 128, 133, 134 - chroococcoides, 171, 172

Nannocytes, 130 Napier Formation, 544, 547, 557, 558 Nash Formation, 364, 367, 681, 682,

Nautiloids, 553, 557 Nemakit-Daldyn Formation, 340 Neruen Formation, 341 -suite, 326, 330, 331, 615, 622, 625,

Nilemah Embayment, 262, 391, 392,

- Sands, 395 Nizva Formation, 341 Nodularia, 128, 136, 137, 138 Nomenclature of stromatolites, 2, 31-32,

43, 45, 339, 359-360, 525 -- parts of bioherms, 36, 41-42, 43 Non-layered fabrics, 227-245 Normandy, 237 Worthwest Territories, 523, 524 Nostoc, 128, 136, 131, 138, 139, 211 Nostocaceae, 128, 136, 137, 138, 154,

Nostocales, 128, 129, 134, 136-138,

Nostochopsidaceae, 128, 139 Nostochopsis, 128, 139 Nouatila, 40, 363, 694 - frutectosa, 518, 521 Nucleella, 39, 347, 694 - figurata, 340, 360, 614 - inconformis, 340

Oceans, Precambrian, 503, 509, 512 Odjick Folimation, 601 Olenek Uplift, 324, 341, 361, 366 Olifants River Group, 635, 636, 640 Omachtenia, 34, 38, 41-42, 329, 360,

362, 368, 313, 375, 603, 605, 606, 608, 609, 694

683

626, 627, 629, 631, 632

395, 396

156, 160, 509, 510

156

- omachtensis, 340, 614

Omakhta Formation, 341 -suite, 366, 615, 622, 623, 624, 631,

Omnia Uplift, 614, 615, 621, 622 Dmnin suite, 366 Omolon Massif, 341 Oncolites, 67, 69, 70, 86, 120, 123,

632

194, 328, 382-384, 487, 531, 545, 555, 556, 557, 563, 578, 603, 614, 624, 639, 653, 656, 674, 691

-, see also Pisolites Onondaga Lake, 487 Ontario, 111 Onverwacht, 156 - Group, 501, 508 Oocardium, 116, 151 - stratum, 117, 121 Ooids, 59, 66, 58, 88, 89, 97, 98, 107,

111, 217, 221, 229, 258, 262, 318, 396, 405, 411, 421, 425, 437, 506, 530, 531, 538, 543, 567, 568, 569, 571, 603, 605, 636, 639

Ordovician, 154, 217, 337, 501, 513, 558 558, 651

Ore Formation, 645-647, 660 - Shale, 663 Organelle, 691 Organic carbon content of Recent

stromatolites, 166-1 6 7 algal mats, 166

---- - compounds in stromatolites,

syngeneity, 168-170, 188 - matter in stromatolites, 150, 159,

163-191 - pigments in stromatolites, 164, 173,

181-182 Orientation of columns, 267, 269, 270,

304, 335 Orthonella, 244, 469 Osagia, 49, 384 Oscillatoria, 119, 128, 135, 136, 151,

221, 382 - okenii, 143 - terebriformis, 143 - williamsii, 171, 172 Oscillatoriaceae, 128, 135-136, 140,

154, 157, 158, 160 Oscillatoriacean stromatolite micro-

structure, 159 Oscillatoriales, 128, 129, 135-136 Oslyan series, 366 Oslyanka Group, 341 Ostracods, 418, 439, 441, 450, 475

INDEX 785

Ostreobium quekettii, 121 Ostrov formation, 340 Ottonosia, 49 Oued Cheikha Group, 517 Oyster Bay, 171, 172 Ozone, 182, 502

Padre Island, 171, 172 Palaeoanacystis vulgaris, 172 Palaeobiology, 3 Palaeolyngbya barghoorniana, 171, 172 Paniscollenia, 38, 347, 694 - emergens, 340, 621, 631 Paradise Creek Formation, 82, 157, 161,

168, 169, 311, 319, 362 Parmites, 38, 347, 363, 369, 627, 694 - aimicus, 340, 614, 629 - concrescens, 49, 340 Parry Bay Formation, 523 Patomia, 39, 347, 362, 368 - ossica, 49, 361 Pearson Formation, 602 Peat, association with stromatolites,

Peels, 11-12 Pekanatui Point Formation, 606 Penge Formation, 635, 637 Pennsylvanian, 651 Periodic calcification, 205-208 - lamina production, 120, 124, 194-195,

251, 277, 281, 301, 303, 304, 307, 382

- deposition of geyserite, 103-107, 109 Peron Sandstone, 393 Persian Gulf, 117, 120, 122, 158, 413,

-- , microbiology of stromatolites,

Pertatataka Formation, 362, 363, 367 Pethei Group, 602, 605, 606, 607,

Phaephila, 151 Phormidium, 116, 119, 128, 135, 136,

146, 151, 207, 209, 211, 224, 232, 239, 243, 247, 275, 276, 277, 281, 288, 301, 307, 493

429, 456, 457

421-433, 465

125-126

608

- crosbyanum, 120 - hendersonii, 118, 120, 125, 126, 197,

- incrustatum, 120, 205, 206, 207, 208,

- laminosum, 143

422

209, 211, 224, 226, 227

- tenue, 197, 299, 300, 305, 307, 308 - truncatum, 297, 300 Phosphate, 567 - in stromatolites, 653, 659 -, see also Phosphorites Phosphoria Formation, 665 Phosphorites, 66 5-666 Photic zone, 146, 152, 507, 509, 510,

- -, see also Euphotic zone Photography, 8-9, 25, 24 Photoheterotrophism, 691 Photophobotaxis, 691 Photosynthesis, 26-27, 114, 122, 123,

127, 150, 151, 152, 153, 218, 502, 506, 508, 691

512

-, advent, 187 -, measuring, 276-277, 299, 301 Phototaxis, 197-198, 211, 224, 275,

' 277, 305, 306, 307, 308, 309, 382, 507, 508, 509, 691

Phototopotaxis, 691 Phototrophism, 691 Phototropism, 201, 403, 547, 551,

Phytems, 348, 350 Pigments, analysis, 276 -, photosynthetic, 24-26, 29, 114,

-, seasonal variation, 289 Pilbaria 40, 694 - perplexa, 362 Pillara Limestone, 543, 546 Pillingini Tuff, 191 Pine Point, 657 Pisolites, 57, 58, 59, 60, 61, 67, 68, 69,

557, 691

153

70, 83-86, 87, 88, 89, 97, 98, 103, 105, 107, 111, 538

-, see also Oncolites Pitella, 39, 694 Plankton, 136, 138, 151, 152, 186, 313,

Planocollina, 35, 31, 39, 694 Plants, element concentration by, 664 Platella, 324, 327 -protensa, 340 Platonovka Formation, 341 Platonov suite, 323, 366 Plectonema, 140 Pleurocapsq 128, 133, 134, 143 Pleurocapsaceae, 128, 134 Pleurocapsales, 128, 130, 133, 134

508, 514, 665

786 INDEX

Pogoryuy suite, 366 Pokegama Quartzite, 311 Polarisbreen Group, 341 Poludia, 40, 694 - polymorpha, 253 Polud Ridge, 341 Pool accretion, see Speleothems Porphyrins in stromatolites, 163, 164 Port Aransas, 171, 172 Posidonia, 396 Pound Quartzite, 362 Procaryote, 691 Proterozoic, Early, 348, 352, 353, 365,

368, 369-370, 500, 501, 502, 503, 509, 510, 515, 599-611, 635-643, 653-654, 655, 656, 658, 677, 680

-- , see also Aphebian Protococcus, 143 Protosystems, 348 Protozoa, 75 Pseudanabaena, 135, 136, 299, 300 Pseudocolumnar, 8, 691 Pseudokussiella, 39, 42 Purcell Supergroup, 166, 167, 365 Pustular mat, 264, 265, 266, 267 Pyrite, 490

Quartz, detrital, in stromatolites, 253, 411, 505, 555, 581-582, 587-590

Radial fabrics, 244-245, 246, 253-255 Rae Group, 523, 524, 525, 532, 533, 534 Rassolnaya Formation, 341 Receptaculites, 55, 559 Rechkin suite, 366 Recluse Formation, 601 Reconstruction, 9-10, 15-20, 39-40,

Reefs, 401, 402, 403, 506, 512, 513,

Renalcis, 547, 563, 575 Repetative lamination, 194 Reticulate fabric, 245 Revett Formation, 650 Rhine River, 123, 235 Rhizophorq 449, 459 Rhodesia, r56 Rhodothamniella, 237 -floridula, 122 Rhythmic processes, see Periodic

processes

53, 337, 591

514, 543-563, 645, 663

Rivularia, 116, 128, 137, 138, 207, 208,

- haematites, 120, 204, 207 Rivulariaceae, 121, 128, 136, 137, 138,

140, 154, 156, 157, 160, 510 Roan Antelope, 646, 670 Rocknest Formation, 599, 600, 601,

Root molds, 61, 63, 64, 65, 66, 68 Rove Formation, 654 Ruisseau du bois d’Haumont, 205, 207 Rum Jungle, 655 Ruppia maritima, 417

237, 255

604, 605

Saanick Inlet, 166 Sabkha, 122, 125, 421-433, 476, 561,

Saccorhiza, 666 Sacculia, 39, 341, 694 Sadanu Formation, 341 Sadler Limestone, 544, 559 Salinity tolerance of organisms, 144-

Sample preparation, 16-17 Sampling, 5-9, 16, 33-40 -, living stromatolites, 23-24 Satkin suite, 366 Scenedesmus quadricauda, 671 Schancharia, 38, 347, 694 Schisto-Calcaire, 363 Schizothrix, 124, 128, 135, 136, 151,

199, 200, 201, 202, 203, 206, 207, 208, 209, 210, 211, 215, 218, 221, 223, 224, 225, 226, 227, 229, 230, 231, 232, 233, 239, 241, 243, 255, 256, 382, 460, 463, 464

199, 253, 450, 451, 454, 483

670

145

- calcicola, 120, 171, 172, 183, 197,

- coriacea, 119 - cresswellii, 119 - fasciculata, 119 - gracilis, 119 - helva, 264, 405 - lateritia, 119 - penicillata, 119 - splendida, 118, 119, 124, 125, 126 Schorikha Formation, 341 Scytonema, 116, 128, 137, 138, 198,

199, 200, 201, 202, 203, 215, 217, 219, 221, 222, 223, 224, 225, 229, 230, 231, 232, 233, 237, 244, 245,

INDEX 787

246, 248, 255, 258, 386, 453, 458, 459, 460, 461, 462, 463, 464, 466, 47 1

- crustaceum, 121, 256 - hoffmanni, 450 - julianum, 121 - mimbile, 121 - myochrous, 121, 449, 451, 454, 455,

- stad. crustaceum, 467, 468, 469 Scytonemataceae, 128, 136, 137, 138,

Scytonemine, 114, 153 Seasonal cycles, 22, 678 Selvage, 582, 691 Semri Series, 367 Serial sectioning, 9-10, 53 Serie des Mines, 648 Serizia, 40, 363, 694 - radians, 518, 520, 521 Serogo-Klyucha suite, 366 Seryi Kluch Pormation, 341 Seton Formation, 602 Shaler Group, 523 Shali Formation, 367 - Series, 364 Shark Bay, 57, 58, 60, 63, 67, 117, 120,

121, 123, 124, 125, 153, 166, 213, 214, 218, 219, 221, 223, 224, 227, 229, 235, 261-271, 318, 359, 385, 389-411, 413, 465, 502, 505, 506, 513, 538, 546, 561, 562, 563, 677, 678, 680, 685

_ - , microbiology of stromatolites, 124 Sheath, 129, 506 -, calcification of, 451, 467, 468, 469,

-, organic chemical composition, 182-

Shorikhin suite, 366 Shuntara Formation, 341 Shuntar suite, 366 Sicily, 559, 561 Silica, 56, 87-112, 413, 419 - in stromatolites, 160, 180, 186, 187,

465, 467, 468, 469

140, 158

47 5

183

189, 273-310, 311-320, 491-496, 498, 603, 635-643

Silurian, 559 Silvermines, 657 Sinter, 87 Siphononema, 128, 133, 140 Siphononemataceae, 128, 134

Siyeh Limestone, 364, 367 Skillogalee Dolomite, 167, 175, 178, 191,

Slabbing, 9-10, 16-17 Slave craton, 599-61 1 - Province, 500 Smooth mat, 264, 265, 266-267 Snare Group, 372 Society Cliffs Formation, 653, 656 Solentia, 128, 131, 132 Sosan Group, 602 South Africa, 152, 156 Spain, 158 Species concept in cyanophytes, 127 Speleothems, 73-86 Sphaerocodium, 547, 551, 557, 558,

Sphaerotilus discophora, 667 Spherulites, 229, 231, 233, 235, 236, ,237, 240, 241

SpiriNum, 2 99, 3 00 Spirochetes, 299, 300 Spirulina, 128, 135, 136, 143 -platensis, 171, 172 Spitsbergen, 341 Spokane Shale, 367 Sponges, 483 Spongiostro ma maeandrin um, 5 2 Sporopollenin, 181, 183 Springs, 117, 120 -, see also Hot springs Staining methods, 11, 17, 217, 248,

Stalactites, see Speleothems Stalagmites, see Speleothems Staraya Rechka Formation, 341 Stark Formation, 602 Starorechen suite, 366 Statistical methods, 11, 12 Steamboat Springs, 100 Steep Rock Lake, 160 Stein am Rhein, 235 Stephanodkcus, 117 Sfichosiphon, 128, 132, 133 Stigonema, 128, 139, 140 Stigonemataceae, 128, 139-140 Stigonematales, 128, 134, 138-140,

Stiriolite, 111, 112 Stmtifem, 34, 35, 38, 41, 42, 323,

363, 413, 510

559

565

157

329, 330, 336, 347, 603, 605, 614, 627, 628, 694

788 INDEX

- biwabikensis, 312, 314, 318, 319 - flexurata, 340, 360 - irregularis, 340 - undata, 340, 360 Stratiform, 691 Streaky microstructure, 285, 691 Striated microstructure, 279, 691 Striped fabrics, 209, - laminated fabric, 194, 204 Stromatactis, 57, 68, 69 Stromatolites, 1, 149, 194, 195 Stromatolitic cyclicity, 328-329, 331,

-- , see also Jacutophyton cycles Stromatoporoids, 543, 544, 545, 559,

Stromatoid, 15 Subtidal environment, 153, 263, 265,

333

56 3

266, 312, 333, 381-388, 389, 393, 396, 397, 399, 401, 403, 404, 405, 406, 407, 422, 426, 443, 507; 510, 511, 51.2, 531-532, 543-563, 585-597, 605, 608, 632, 638, 639, 675, 677, 681, 685

639-643 -- , geochemical distinction, 188,

Sukhaya Tunguska Formation, 341

Sukhopit Group, 341 - Series, 366 Sukhotungusin suite, 366 Sulphides in stromatolites, 122, 126,

651-653, 669-671 Sulphur, 490 Sutton Lake Group, 372 Svetliella, 39, 347, 352, 361, 362, 368,

- svetlica, 340 - tottuica, 614 Svetlinian horizon, 350, 351 Svetlin suite, 366 Svetlyi Formation, 341 Svetly suite, 615, 627, 631, 632 Swaziland Sequence, 150, 181, 501 Sygynakhtakh Formation, 341 - suite, 366 Symploca 7aete-viridis, 120, 264 - thermalis, 143 Synechococcuq 88, 128, 130, 131, 132,

- lividus, 143, 299, 300 - minervae, 143, 299, 300

river, 323 --

694

493

Synechocystis, 128, 130, 131, 132 - aquatilis, 143 Synoptic profile, 7, 691 -- , see also Lamina shape .- relief, 305, 324, 325, 326, 382, 385,

393, 401, 405, 505, 506, 513, 527, 529, 531, 538, 557, 636, 639, 691

Taemas, 191 Takiyuak Formation, 601 Taltheilei Formation, 599, 600, 605,

Talyn Formation, 341 - suite, 366, 615 Tamala Eolianite, 395 Tamjout series, 650, 656 Taoudenni Basin, 252, 255, 257, 363,

Tapley Hill Formation, 191 Tarioufetia, 38, 363, 694 - hemispherica, 253, 518, 520, 521 Taseyeve Group, 341 Taseyev Series, 366 Taxis, 591 Taxonomy of stromatolites, see

Techniques, living stromatolites, 274-

-, microfossiliferous stromatolites, 317 Tennessee, 651 Ten upa luse 114 3 9 Tepee structures, 64, 70 Terminal Riphean, 337, 347, 348, 621 Texas, 64, 114, 171, 172 Thalassia, 382, 383 Thalkedar Limestone, 364 Thermal environments, 132, 136, 139,

-- , see also Hot springs Thermocline, 691 Thin-section, 10-11 Thiobacillus, 668 Thrombolites, 227-235, 236, 237, 238,

Thunder Bay, 313 Tidal amplitude, 673, 679, 680, 681 - cycles, 22 Tides, 107 Tien-Shan, 360, 361, 362 Tifounke Group, 367, 517, 518 Tifounkeia, 40, 363, 694 - ramificata, 256, 257, 518, 521

606, 608

369, 517-522

Classification of stromatolites

278, 480-481

141, 142, 143, 146, 180

240, 241, 403, 603, 691

INDEX 789

Tilemsina, 40, 335, 363, 369, 627, 694 Tillite, see Glacial sediments Tinnia, 39, 347, 694 - patomica, 340 Tochatwi Formation, 602 Tolypothrix, 128, 137, 138 Tooganinie Formation, 175, 178, 191 Toolonga Calcilutite, 393 Top Crossing, 191 Totta Formation, 341 - suite, 615 Transvaal Dolomite, 152, 156, 160, 167,

187, 188, 311, 502, 656 - Supergroup, 635-643, 655 Trapstones, 237, 239 Travertine, 70, 115, 117, 122, 123 Triassic, 559, 657 Trichomes, 129 Trichodesmium, 128, 135, 136, 151 Tropism, 691 Trucial Coast, see Persian Gulf True branching (cyanophytes), 138 Tsipanda Formation, 341, 615, 627,

Tsipandian horizon, 350, 351, 352 Tsipandin suite, 366 Tsumeb, 364 Tufted mat, 264, 265, 267 Tully Green Lake, 486 Tungusik Group, 341 - Series, 366 Tunguska River, 36 Tungussia, 34, 35, 40, 42, 49, 335, 347,

351, 352, 361, 362, 363, 364, 368, 369, 373, 374, 375, 376, 379, 505, 614, 627, 694

629, 632

- bassa, 42 - colcimi, 340 - confusa, 340 - cumata, 518, 519, 520, 521 -globulosa, 253, 254, 255, 518, 520,

- indica, 253 - inna, 191, 362 -- , organic chemistry, 175, 178 - nodosa, 256, 257, 340 - wilkatanna, 191 -- , organic chemistry, 175, 178 Tunicata, 347, 694 Turkut suite, 366 Turuchania, 39, 361 Turukhan Formation, 341

521

- region, 323, 329, 330, 332, 333, 341,

- suite, 366 Tussock microstructure, 253-255, 259 Tynagh, 657

350, 352, 360, 361, 366

Uchur depression, 615, 621, 622, 624 - Group, 341 - Series, 360, 366 Uchur-Maya platform, 613, 614, 615 -- region, 326, 330, 331, 341, 350,

Udaipur, 653, 659 Uderey suite, 366 Uii group, 341, 615, 622 Uk Formation, 341 - suite, 361, 366 Ulkan series, 366 Ultrastructure, 194 Ultraviolet light, 152, 153, 502, 503 Umberatana Group, 191, 362, 363,'

Uniformitarianism, 407-409 Union Island Group, .602 Ural Mountains, 360, 361, 366 Urals, south, 341 Uranium in stromatolites, 655, 656, 659 Urquhart Shale, 649 Uricatella, 347 - urica, 255 Ust-Ilyin suite, 366 Ustkirba suite, 615 Ust-Kirbin suite, 366 Ust-Kotykan suite, 366 Utsingi Formation, 606, 608 Uyan Series, 366 Uy Series, 366

360, 361, 366, 613-633

367

Vacerrilla walcotti, 273-310, 494 Vaucheria, 116, 151 - geminata, 122 Vendian, 154, 158, 255, 256, 258, 259,

338, 348, 360, 361, 362, 369, 374, 512, 513, 694

Venerupis, 417, 419 Ventersdorp System, 364 Vermiculus, 518, 520, 522 Vermiform microstructure, 159, 255,

Vertical differentiation in algal mats,

Vetella, 39, 347, 694

259, 691

see Zonation

790 INDEX

Veteranen Group, 341 Victoria Island, 523 Vindhyachal Hills, 364 Vindhyan, 364, 367 Virgin Hills Formation, 544, 547, 551,

553, 559, 557, 558

Wall, 12, 582, 691 Water potential, ,144-145 West Congo sequence, 367 Wildbread Formation, 606 Wilkawillina Gorge, 191 - Limestone, 167, 191, 565 Williams Bore, 191 Wilpena Group, 191, 367 Wilson, 191 Windjana Limestone, 543, 547 Windward Lagoon, 197 Winnipegosis Formation, 182 Winston Point, 313 Wirrealpa, 565-584 Witwatersrand, 656 Wonoka Formation, 175, 178, 191 Woodcutters, 670 Wopmay orogen, 523, 524, 599, 600,

Wyloo Group, 367 Wyoming, 87 -, see also Yellowstone National Park

601

Xenococcus, 128, 133, 134

Yellowstone National Park, 25, 87-112, 114, 146, 180, 181, 187, 197, 253, 273-310, 489-498, 532, 669, 676, 677

Yenisey Mountains, 360, 361, 366 - Ridge, 341, 352 Yu . . ., see also Ju . . . Yudoma-Maya Trough, 614, 615, 621,

622, 625, 627, 629, 631 Yudoma suite, 366, 621, 622, 629-631 Yudomian, 338, 348, 349, 361 Yurmata Group, 341 Yurmatian Series, 352 Yurmatin Series, 366 Yusmastakh Formation, 341 - suite, 328, 366

Zaire, 645, 649, 654 Zambia, 645, 648, 663 Zambian Copperbelt, 645-650, 660 Zigazino-Komarov suite, 366 Zilgalgin suite, 366 Zilmerdak Formation, 341 - suite, 351, 366 Zonation, vertical, in microbial mats,

Zosterosphaera tripunctata, 164 25, 120, 125, 126

Stromatolites (791 p, 1976)

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