Dissertation Tobias Weisenberger

ZEOLITES IN FISSURES OF

CRYSTALLINE BASEMENT ROCKS

INAUGURALDISSERTATION

zur

Erlangung des Doktorgrades

der Fakultät für Chemie, Pharmazie und Geowissenschaften

der Albert-Ludwigs-Universität Freiburg im Breisgau

vorgelegt von

TOBIAS WEISENBERGER aus Emmendingen

2009

Vorsitzender des Promotionsausschusses: Prof. Dr. Rolf Schubert

Referent: Prof. Dr. Kurt Bucher

Korreferent: Prof. Dr. Reto Gieré

Tag des Promotionsbeschlusses: 9.. Juli 2009

Saint Barbara statue in the Arvigo quarry–

Patron saint of geologists and firemen Saints day: 4th December

Der Tag der heiligen Barbara! - Feierlich stehen sie alle da, die Männer, die aus des Berges Nacht - das schwarze Gestein zu Tage gebracht, das dort gelegen seit Urweltzeit; - bald wird es vom roten Feuer gefreit. Feierlich stehen sie alle da.- Es ist 4. Dezember: St. Barbara! Du Schutzpatronin, St. Barbara, - Im Schmucke treten sie alle dir nah'; An dem Tschako wiegt sich die schwarze Feder, - Schwarz ist ja alles, Anzug und Leder. Dort sind die weißen, Musik trägt rot. - In Ordnung und Würde, wie nach Gebot beginnt der Zug, und wer ihn sah',- weiß, es ist heute St. Barbara! Zurück von der Kirche St. Barbara. - Und es geschieht, was immer geschah, Musik spielt lustig, die Federn winken, - in Oberschlesien will man auch trinken, sorglos sich freuen, den Tag genießen, - wen sollte das heitere Volk verdrießen? Und es geschieht, was immer geschah! - Nur einmal im Jahr ist St. Barbara!

St. Barbara; Poem after Käthe Gutwein

ABSTRACT I

ABSTRACT

The goal of the thesis is to study the occurrences and formation of zeolites hosted in

crystalline basement rocks. The low-grade fissure mineral assemblages including

zeolites are the key to the appreciation of water-rock interaction in hydrothermal and

geothermal systems at relatively low temperatures (< 250 °C) located in granites and

gneisses of the crystalline basement. Extensive work is done on zeolite occurrences in

sedimentary rocks often pyroclastic origin and volcanic rocks, whereas elements

necessary for zeolite formation derive from primary glass of from feldspar. In contrast

the formation of zeolites in granites and gneisses is poorly studied and no systematic

of evaluation and spatial distribution are carried out either chemical studies on

zeolites or formation consideration are done.

Therefore a systematic evaluation of zeolites in the Central Swiss Alps is

presented. Ca-zeolites occur in various assemblages in late fissures and fractures in

granites and gneisses. The systematic study of zeolite samples showed that the

majority of finds originate from three regions particularity rich in zeolite-bearing

fissures: (1) in the central and eastern part of the Aar- and Gotthard Massif, including

the Gotthard road tunnel and the Gotthard-NEAT tunnel, (2) Gibelsbach/Fiesch, in a

fissure breccia between Aar Massif and Permian sediments, and (3) in Penninic

gneisses of the Simano nappe at Arvigo (Val Calanca). The excavation of tunnels in

the Aar- and Gotthard massif give an excellent overview of zeolite frequency in

Alpine fissures, whereas 32 % (Gotthard NEAT) and 18 % (Gotthard road tunnel) of

all fissures are filled with zeolites. The number of different zeolites is limited to 6

species: laumontite, stilbite and scolecite are abundant and common, whereas

heulandite, chabazite and epistilbite occur occasionally. Ca is the dominant extra-

framework cations, with minor K and Na. Heulandite and chabazite additionally

contain Sr up to 29 and 10 mole%, respectively. Na and K content of zeolites tends to

increase during growth as a result of systematic changes in fluid composition and/or

temperature. The K enrichment of stilbite found in surface outcrops compare to

stilbite in the subsurface may indicate late cation exchange during interaction with

surface water. Texture data, relative age sequences derived from fissure assemblages

and equilibrium calculations shows that the Ca-dominated zeolites precipitated from

fluid with decreasing temperature in the order (old to young = hot to cold): scolecite,

ABSTRACT II

laumontite, heulandite, chabazite and stilbite. The components necessary for zeolite

formation are derived from dissolving primary granite and gneiss minerals. The

nature of these minerals depends on the metamorphic history of the host rock.

Zeolites in the Aar Massif derived from the dissolution of epidote or calcite and albite

that were originally formed during Alpine greenschist metamorphism. Whereas

albitization of plagioclase in higher grade rocks releases the necessary components for

zeolite formation, a process that is accompanied by a distinct porosity increase.

Zeolite fissures occur in the zone where fluid inclusions in earlier formed quartz

contain H2O dominated fluids. This is consistent with equilibrium calculations that

predict a low CO2 tolerance of zeolite assemblages particularly at low temperature.

Pressure decrease along the uplift and exhumation can increase zeolite stability. The

major zeolite forming reaction consumes calcite and albite; it increases pH and the

total of dissolved solids. The produced Na2CO3 waters are in accord with reported

deep groundwater (thermal water) in the continental crust, which are typically

oversaturated with respect to Ca-zeolites.

A detailed local study of the mineralogical, chemical and petrological evolution

of crystalline basement rocks in Arvigo was performed to assess information about

the evolution of fluid-rock interaction during uplift of the Alpine orogen. The Arvigo

fissures contain the assemblage epidote, prehnite, chlorite and various species of

zeolites. Fluid rock interaction takes place along a retrograde exhumation path which

is characterized with decreasing temperature by: (1) coexisting prehnite/epidote, that

reveals temperature conditions of 330 – 380 °C, (2) chlorite formation at temperature

of 333 ± 32 °C and (3) formation of zeolites <250 °C. The formation of secondary

minerals is related to the hydrothermal replacement reaction during albitization and

chloritization that releases components for the formation of Ca-Al silicates and form a

distinct reaction front. The fluid-rock interaction is associated with a depletion of

Al2O3, SiO2, CaO, Fe2O3 and K2O in the altered wall rock. The reaction is associated

with an increase in porosity up to 14.2 ± 2.2 %, caused by the volume decrease during

albitization and the removal of chlorite. The propagation of the sharp reaction front

through the gneiss matrix occurred via a dissolution-reprecipitation mechanism.

Zeolite formation is tied to the plagioclase alteration reaction in the rock matrix,

which releases components for zeolite formation to a CO2-poor, alkaline aqueous

fluid.

ABSTRACT III

A combined study of 40Ar/39Ar age dating, apatite fission track (FT) and chemical

characterization of tunnel and surface samples are present to carry out the position of

low-temperature water-rock interaction in respect to the Alpine history. Apatite FT

analysis yields an exhumation rate of 0.45 mm a-1, a cooling rate of 13 °C Ma-1 and a

geothermal gradient of 28 °C km-1. Combining these with the 40Ar/39Ar plateau age

for apophyllite of ∼2 Ma, a minimum formation temperature and depth of 70 °C and

2800 m, respectively can be assumed. Temperature-time evolution of fissures in the

Aar Massif and thermodynamic mineral evolution indicate that laumontite were

formed between 7 and 2 Ma before present at temperatures between 150 and 70 °C.

ACKNOWLEDGMENTS IV

ACKNOWLEDGMENTS

It is a great pleasure for me to thank the many people who made this thesis possible.

I am most thankful to my advisor Prof. Dr. Kurt Bucher. I thank him for

awarding the topic of this thesis and an outstanding supervision. I am glad that he

gave me the opportunity to continue my research interests on zeolites that I

experience during my diploma thesis. I always appreciated the discussion and

constructive criticism with him. A special appreciation has to be mentioned that he

gave me the freedom to develop and follow my own ideas. He enhanced me to

educate myself during numerous DMG workshops and teached me to deal with

thermodynamic calculations, which was not even easy with me. During theses years

Kurt must have used up a lifetime supply of red ink pens to teach me geological

common sense. This thesis would not have been possible without the great advise and

trust in me. Thank you!

Also, I want to thank Prof. Dr. Reto Gieré for the additional supervision,

numerous discussions and the takeover of the co-referee.

I want to thank my parents, who always gave me the liberty to follow my interest

and supported me during my education in a loving environment.

A special thank to all the people who supported me during my search for Alpine

zeolites and supplied samples: Peter Amacher, mineral representative of the NEAT

Amsteg-Sedrun section who provided high-quality minerals specimen and who was

always easy to contact for discussion. Beda Hoffmann and Peter Vollenweider from

the Swiss Natural History Museum in Bern, giving me the possibility to study their

mineral collection and their encourage during my work in the “dungeon”. Giovanni

Polti and Alfredo Polti SA for permission to do field work in the active quarry in

Arvigo, especially Luigi who took care about me during blasting, even we conduct

our conversion by signs, due to my lack in Italian language.

I appreciate all the help and support that I get by using technical equipment in

external research institutes: Prof. Dr. Stefan Graeser form the Mineralogical Institute

Basel, who provided me the possibility to use the FTIR instrument; Dr. Egbert Keller

from the Crystallographic Institute Freiburg, who guided me through DSC-TGA

measurements and Andreas Leemann from the Swiss Federal Laboratories for

ACKNOWLEDGMENTS V

Materials Testing and Research for impregnation of rock samples. Roelant van der

Lelij from the Department of Mineralogy in Geneva for the apophyllite dating and

helpful discussion and PD Dr. Meinert Rahn for helping with the apatite fission track

analysis and the always profitable conversations.

I wish to thank Zeng Lu, Fleurice, Siggi, Zhou Wei, Hiltrud and Duy Anh Dao

for their always open doors, where I find a sympathetic ear for discussion.

The Friedrich Rinne Stiftung of the Albert-Ludwigs-University, Freiburg for the

financial support.

Last but not least I want to thank Simon and Rune. I am very happy to call them

my friends. We had a wonderful time during our diploma thesis and they encouraged

me during my PhD study whenever I needed them. Thanks Rune for the long

insightful phone calls, the high email exchange rate and help when I need a cheer-up.

Thanks Simon for the friendship during the last years and all the help, which

contribute to the succeed of this thesis – feel free to ask me if you need a red ink pen

again!

TABLE OF CONTENTS VI

TABLE OF CONTENTS

ABSTRACT I

ACKNOWLEDGMENTS IV

TABLE OF CONTENTS VI

1. INTRODUCTION 1

1.1. LAYOUT OF THE THESIS 1

1.2. MOTIVATION OF THIS STUDY 3

1.3. ZEOLITES 4

1.4. ZEOLITE STRUCTURE AND CHEMISTRY 6

1.5. ZEOLITE OCCURRENCES, ZEOLITE FACIES AND ZEOLITE

ZONES 7

1.6. ZEOLITE STABILITY 11

1.7. GEOLOGICAL SETTING 12

1.8. REFERENCES 15

2. ZEOLITES IN FISSURES OF GRANITES AND GNEISSES OF

THE CENTRAL ALPS 20

2.1. ABSTRACT 21

2.2. INTRODUCTION 22


2.3.1. Metamorphic conditions during Alpine orogenesis 26

2.4. SAMPLING AND ANALYTIC METHODS 29

2.5. ZEOLITES IN THE CENTRAL ALPS 30

2.5.1. Spatial distribution 31

2.5.2. Field occurrences 35

2.5.2.1. General features 35

2.5.2.2. Aar Massif/Gotthard NEAT tunnel 43

2.5.2.3. Gotthard massif/ Gotthard road tunnel 45

2.5.2.4. Gibelsbach/Fiesch 46

2.5.2.5. Arvigo/Val Calanca 47

2.5.3. Mineralogy and crystal chemistry of zeolites and associated minerals 48

2.5.3.1. Chabazite-Ca 48

2.5.3.2. Heulandite-Ca 51

2.5.3.3. Laumontite 52

2.5.3.4. Scolecite 53

TABLE OF CONTENTS VII

2.5.3.5. Stilbite/Stellerite 54

2.6. DISCUSSION 55

2.6.1. Reactions and processes of zeolite formation 58

2.6.2. Assemblage stability and phase relationships involving zeolites 61

2.6.3. Fluid composition 64

2.7. CONCLUSIONS 67

2.8. ACKNOWLEDGMENTS 68

2.9. REFERENCES 68

3. POROSITY EVOLUTION, MASS TRANSFER AND

PETROLOGICAL EVOLUTION DURING LOW

TEMPERATURE WATER-ROCK INTERACTION IN

GNEISSES OF THE SIMANO NAPPE - ARVIGO, VAL

CALANCA, GRISONS, SWITZERLAND 76 3.1. ABSTRACT 77



3.4. PREVIOUS WORK 82

3.5. SAMPLING AND ANALYTIC METHODS 83

3.6. RESULTS 84

3.6.1. Petrography 84

3.6.1.1. Unaltered rock 86

3.6.1.2. Altered rock 86

3.6.1.3. Fissure minerals 87

3.6.1.4. Changes in modal mineralogy 87

3.6.2. Mineralogy and mineral chemistry 87

3.6.2.1. Plagioclase and its alteration products 92

3.6.2.2. Biotite-chlorite 95

3.6.2.3. Muscovite 95

3.6.2.4. K-feldspar 95

3.6.2.5. Quartz 96

3.6.2.6. Epdiote 96

3.6.2.7. Prehnite 97

3.6.2.8. Zeolites 99

3.6.3. Porosity 102

3.6.4. Whole rock geochemistry and mass changes 103

3.7. DISCUSSION 105

3.7.1. Mineral reactions 105

TABLE OF CONTENTS VIII

3.7.2. Mass changes and element mobility 107

3.7.3. Mineral stability and mineral equilibria 110

3.7.3.1. Prehnite and epidote 110

3.7.3.2. Chlorite 112

3.7.3.3. Zeolites 113

3.7.4. Mineral evolution 117

3.7.5. Fluid accessibility and composition 120

3.8. CONCLUSION 121


3.10. REFERENCES 124

4. TIMING AND MINERAL EVOLUTION DURING LOW-

TEMPERATURE FLUID-ROCK INTERACTION ON UPPER

CRUSTAL LEVEL: 40Ar/39Ar APOPHYLLITE-(KF) DATING

AND APATITE FISSION TRACK ANALYSIS ON ALPINE

FISSURES (CENTRAL ALPS/SWITZERLAND) 132 4.1. ABSTRACT 133



4.4. MATERIAL AND METHODS 137

4.4.1. Analytic 137

4.4.2. Samples 140

4.5. RESULT 140

4.5.1. Petrography and geochemistry 140

4.5.2. Mineralogy and geochemistry 143

4.5.2.1. Laumontite 143

4.5.2.2. Apophyllite-(KF) 144

4.5.3. Ar/Ar age 146

4.5.4. Apatite FT analysis 147

4.6. DISCUSSION 149

4.6.1. Mineral reaction 149

4.6.2. Depth and temperature estimation 151

4.6.3. Thermodynamic approach 152

4.6.4. Alpine history 154

4.7. CONCLUSION 155


4.9. REFERENCES 157

TABLE OF CONTENTS IX

APPENDIX

I OWN CONTRIBUTION i

II PUBLICATIONS ii

III CURRICULUM VITAE v

INTRODUCTION 1

1. INTRODUCTION

1.1. LAYOUT OF THE THESIS

The thesis is divided into 4 chapters. The 1st Chapter gives an overview and structure

of the thesis and the research interest that is associated with this thesis. Additionally

general information about zeolites, zeolite occurrences and the geological setting of

the working area are given to introduce into the topic of natural zeolites.

Chapter 2, 3 and 4 are separated chapters covering each a manuscript and

therefore analytic methods are described in each chapter, including measuring

methods, measuring conditions and standards that were used. The chapters are

arranged to start with a general overview and evaluation of zeolite occurrence and

formation in crystalline rocks and continue with two detailed local studies to

understand the process of zeolite formation in crystalline basement rocks and the

relation of zeolite formation in respect to the Alpine history (Table 1.1). At the

beginning of each chapter an overview of my contribution to finalize each manuscript

is given.

Chapter 2 is concerned with the documentation and compilation of all known

zeolite occurrences in the Central Swiss Alps to get information about the spatial

distribution. Zeolites hosted in granites and gneisses are chemically characterized to

get information about chemical changes during growth as well as information about

chemical variations in different settings within the Central Alps. Petrographic

observation yields information about zeolite-forming reactions that are formulated in

this chapter. Thermodynamic phase-diagram-modeling to approach the conditions at

which zeolites are formed is used to discuss the zeolite appearance in respect to fluid

composition variation in the Swiss Alps.

This work was presented at the “Zeolite06 Conference” in Socorro, USA in July

2006 and at the annual meeting of the “Deutsche Mineralogische Gesellschaft” in

Hannover, Germany in September 2006. The manuscript was submitted to Journal of

Metamorphic Geology (March 2009) as:

Weisenberger T. and Bucher K.: ZEOLITES IN FISSURES OF GRANITES AND GNEISSES

OF THE CENTRAL ALPS. Journal of Metamorphic Geology

INTRODUCTION 2

Chapter 3 presents a detailed petrographic, mineralogical and geochemical study

of the zeolite locality Arvigo. The active quarry gives good access to many factures

filled with zeolites, which are characterized by a leaching zone from which elements

for secondary mineral formation are derived. Mass balance calculation, element

mobility and porosity measurements were done in combination with thermodynamic

phase diagram modeling to show the approach and applicability of thermodynamic

modeling at such problems, including PT-path modeling, mineral and porosity

evolution and relation of fluid composition with respect to the zeolite formation.

This work was presented at the “17th Goldschmidt Conference” in Cologne,

Germany in August 2007 and at the “Swiss Geoscience Meeting” in Geneva,

Switzerland in November 2007. The manuscript will be shortly submitted to

Contributions to Mineralogy and Petrology as:

Weisenberger T. and Bucher K.: POROSITY EVOLUTION, MASS TRANSFER AND

PETROLOGICAL EVOLUTION DURING LOW TEMPERATURE WATER-ROCK INTERACTION IN

GNEISSES OF THE SIMANO NAPPE - ARVIGO, VAL CALANCA, GRISONS, SWITZERLAND.

Contributions to Mineralogy and Petrology

Chapter 4 is a combined study of apophyllite Ar/Ar age dating, apatite fission

track analysis, chemical phase analysis and thermodynamic modeling in order to

constrain the evolution of zeolites with respect to the Alpine history. This yields

information about minimum temperature conditions of laumontite formation.

Apophyllite age dating was performed by R. van der Lelij at the Mineralogical

department of the University Geneva, Switzerland. Apatite fission track analyses were

carried out by PD Dr. M. Rahn at the Mineralogical department of the University

Basel, Switzerland.

Parts of this work were presented at the “33rd International Geological Congress”

in Oslo, Norway in August 2008 and at the annual meeting of the “Deutsche

Mineralogische Gesellschaft” in Berlin, Germany in September 2008. The manuscript

will be shortly submitted to Mineralogical Magazine as:

INTRODUCTION 3

Weisenberger T., Rahn M., van der Lelij R., R. Spikings and Bucher K.: TIMING

AND MINERAL EVOLUTION DURING LOW-TEMPERATURE FLUID-ROCK INTERACTION ON

UPPER CRUSTAL LEVEL: 40AR/39AR APOPHYLLITE-(KF) DATING AND APATITE FISSION TRACK

ANALYSIS ON ALPINE FISSURES (CENTRAL ALPS/SWITZERLAND). Mineralogical Magazine

Table 1.1: Overview of thesis and major topics and major questions addressed to each chapter.

Zeolites in fissures of crystalline basement rocks Chapter 2: Zeolites in fissures of granites and gneisses of the Central Alps

Chapter 3: Porosity evolution, mass transfer and petrological evolution during low temperature water-rock interaction in gneisses of the Simano Nappe - Arvigo, Val Calanca, Grisons, Switzerland

Chapter 4: Timing and mineral evolution during low-temperature fluid-rock interaction on upper crustal level: 40Ar/39Ar apophyllite-(KF) dating and apatite fission track analysis on Alpine fissures (Central Alps/Switzerland)

• general overview and compilation of all zeolite localities in the Central Swiss Alps • what zeolite species occur in crystalline basement rocks? • spatial distribution of zeolites • chemical characterization of zeolites • zeolite forming reactions • what factors control the zeolite formation? • thermodynamic phase-diagram calculation approaching physical and chemical conditions

• detailed study of a zeolite locality to assess the formation of Ca-Al silicates during uplift • temporal evolution of secondary minerals • mineral reactions during water-rock interaction • mass transfer during water-rock interaction • porosity evolution during water-rock interaction • fluid-composition in relation to the formation of zeolites • thermodynamic phase-diagram calculation to approach physical and chemical conditions

• detailed study of a zeolite locality to assess the formation of Ca-Al silicates during exhumation • timing of zeolite formation • depth of zeolite formation • chemical change during formation of secondary minerals • relation of zeolite formation with respect to the Alpine history • estimation of laumontite forming temperature

1.2. MOTIVATION OF THIS STUDY

Already during my diploma thesis in 2004 and 2005 on zeolites in Iceland, one of the

worldwide famous zeolite localities (e.g. Walker 1959, 1960) I got attracted to the

zeolite mineral group. Simultaneous ongoing excavation of the New Gotthard

Railway Base tunnel (NEAT) supplied a large amount of zeolites specimen during the

drilling period. This finds, the possibility to get access to these specimen and the

current research interest on deep continental fluids triggered Kurt Bucher, my adviser

and me to found this PhD project about the zeolite formation in crystalline basement

rocks that is driven by fluid-rock interaction.

INTRODUCTION 4

But what makes zeolites in crystalline basement rocks so interesting?

Although zeolites are known in fissures and gashes of the crystalline basement

from mineral collectors since about 150 years (e.g. Kenngott 1866; Parker 1922;

Niggli et al. 1940; Huber 1943; Sigrist 1947; Stalder et al. 1998), they did not affect

the interest of the scientific research community. So far, previous publications about

zeolite occurrences and their genesis in the central Swiss Alps and other areas of

crystalline basement rocks are limited, which easily can be count on one hand

(Armbruster et al. 1996; Freiberger et al. 2001; Fujimoto et al. 2001; Ciesielczuk and

Janeczek 2004). Considering the latest special edition on natural zeolites (Bish and

Ming 2001), only two short notes were made on zeolites hosted in granites and

gneisses and so far no research was done on the chemical characterization and spatial

distribution of those zeolites, like in other environments. This lack of research interest

may be related to the economically non-profitable occurrences of zeolites hosted in

basement rocks compared to the well-known and widely used deposits of natural

zeolites from zeolitized volcanic tuffs and sediments.

Nevertheless, geochemical studies of deep continental fluids suggest that many

crystalline basement aquifers are oversaturated with respect to zeolites (e.g. Stober

and Bucher 1999; Bucher and Stober 2000). Therefore zeolites play a role when

considering mass transfer, porosity and permeability of these aquifers, which are

important research areas in relation to geothermal energy production as well as the

problematic storage of nuclear waste in crystalline basement rocks.

1.3. ZEOLITES

Zeolites are among the most common products of chemical interaction between

groundwaters and the Earth’s crust during diagenesis and low-grade metamorphism

(e.g. Bish and Ming 2001). Zeolite minerals occur in low temperature (<250 °C), low

pressure (<200 MPa), water saturated environments. The required amount of silica,

alumina, and alkali and alkali-earth cations necessary for the formation of zeolites is

commonly derived from dissolution of volcanic glass and primary phases.

INTRODUCTION 5

Zeolites are tectosilicates characterized by an open framework structure of Si and

Al surrounding channels of ~2-10 Å in size which contain molecular water and

charge-balancing cations of alkali and alkali-earth metals (e.g. Neuhoff et al. 2000;

Armbruster and Gunter 2001). Their unique and distinct crystal structures result in a

large molar volume, high cation-exchange capacities, and molecular sieve capabilities

(e.g. Gottardi and Galli 1985; Bish and Ming 2001). These properties lead to

widespread industrial application in water softening, catalysis, water and wastewater

treatment, agriculture, nuclear waste storage, heating and refrigeration and

construction industry (e.g. Murphy et al. 1978; Kalló 2001; Ming and Allen 2001;

Tchernev 2001; Hauri 2006).

During the past decades with the onset of analyses by electron microprobe,

thousands of zeolite crystals have been analyzed showing a wide compositional range,

and several new minerals with framework structures were discovered. This was

achieved by the advent of automated single crystal X-ray diffractometers, resulting in

much more detail concerning zeolite framework structures. With impeding

nomenclature problems, the International Mineralogical Association’s Commission

on New Minerals and Mineral Names assigned a subcommittee to review all minerals,

and proposes a new definition and a system of nomenclature of zeolites.

The report of the International Mineralogical Association, Commission on New

Minerals and Mineral Names, contains the following definition:

“A zeolite mineral is a crystalline substance with a structure characterized by

a framework of linked tetrahedra, each consisting of four O atoms

surrounding a cation. This framework contains open cavities in the form of

channels and cages. These are usually occupied by H2O molecules and extra-

framework cations that are commonly exchangeable. The channels are large

enough to allow the passage of guest species. In the hydrated phases,

dehydration occurs at temperatures mostly below about 400°C and is largely

reversible. The framework may be interrupted by (OH,F) groups; these

occupy a tetrahedron apex, which is not shared with adjacent tetrahedra”

(COOMBS et al. 1998).

INTRODUCTION 6

Dehydration properties from Cronstedt’s original definition (1756) and a framework

structure from Hey’s 1930 definition are retained. However, following the new

definition (COOMBS et al. 1998), the framework needs not to be only aluminosilicate.

Beryllosilicate, aluminophosphate, and a few similar compositions are allowed by

definition.

1.4. ZEOLITE STRUCTURE AND CHEMISTRY

The basic feature of all zeolite structures is an aluminosilicate framework

(tectosilicate) composed of (Si,Al)O4 tetrahedra, each oxygen of which is shared

between two tetrahedrons (Armbruster and Gunter 2001). The net negative charge on

the tectosilicate framework is balanced by the incorporation of cations (extra-

framework cations) in cages or channels. In most cases Ca2+, Na+ or K+ and less

frequently Li+, Mg2+, Sr2+ and Ba2+ are situated in cavities within the framework

structures. This feature can also be observed in feldspar and feldspathoid minerals.

But in contrast to feldspar and feldspathoid minerals the zeolite aluminosilicate

framework contains open cavities and open channels (i.e. they have lower densities)

through which ions can be either extracted or introduced in the structure (Armbruster

and Gunter 2001).

Their compositions are represented by the structural formula (1):

(A+z)y/z(B+3)y(Si)xO2(x+y) · nH2O (1)

Where A represents extra-framework cations (such as Na+, K+, Ca2+, Ba2+, Sr2+, Mg2+

and Fe2+), B are tetrahedral coordinated trivalent cations in the zeolite framework

(Al3+ and Fe3+), z is the charge of the extra-framework cations, n is the number of

moles of extra-framework molecular water, and x and y are the stoichiometric

coefficients for trivalent cations and Si4+ in tetrahedral sites, respectively. The

quantities y/z and 2(x+y) represent the stoichiometries of the extra-framework cations

and framework oxygens, respectively, necessary for maintaining charge balance in the

tectosilicate lattices of zeolites (Armbruster and Gunter 2001).

INTRODUCTION 7

An additional feature, which separates zeolites still further from the feldspar and

feldspathoid minerals, is the presence of water molecules within the structural

channels. These are relatively loosely bound to the framework and cations, and like

the cations they can be removed and replaced without disrupting framework bonds

(Deer et al. 2004).

Three types of solid solutions in zeolites are consistent with the stoichiometry of

equation (1). These solutions are not strictly coupled and can occur independently

from other substitutions as long as charge balance is maintained (Neuhoff et al. 2000).

The first of these is the solid solution within the tetrahedral sites. Tetrahedral

substitution of Si4+ and Al3+ observed in zeolites is highly variable, whereas the

substitution Fe3+ for Si4+ or Al3+ is limited. Secondly, solid solutions among extra-

framework cations are often quite extensive, as evidenced by the large ion-exchange

capacities of some zeolites (e.g. Colella 1996). Total extra-framework ion charge is

necessarily a function of Al3+ and Fe3+ content. Zeolites with high Si/Al ratios

commonly are richer in monovalent extra-framework cations than are more aluminous

samples of the same species. Twice as many monovalent ions as divalent ions are

necessary to compensate for charge imbalances caused by Al3+ in the framework, and

the additional monovalent ions often occlude H2O molecules present in isostructural

zeolites with divalent extra-framework cations (e.g. natrolite and scolecite, Ca-

heulandite and Na-heulandite). The third type of solid solution in zeolites is the

variation in water content, which a consequence of the loose bounding nature of

molecular water in zeolites, whereas the total water content is a sensitive function of

temperature, total pressure and the partial hydrostatic pressure (Neuhoff et al. 2000).

1.5. ZEOLITE OCCURRENCES, ZEOLITE FACIES AND

ZEOLITE ZONES

Zeolites are formed during reaction of aqueous fluids and rocks in several different

geological environments: Most zeolite occurrences formed during diagenetic

processes in sedimentary rocks (including volcanoclastic deposits) which can be

grouped into several types of geological environments or hydrological systems (Hay

1966, 1977; Hay and Sheppard 1977; Surdam 1977; Gottardi 1989; Hay and Sheppard

INTRODUCTION 8

2001), like (1) hydrologically open systems (e.g. Hay and Sheppard 2001), (2)

hydrologically closed systems (e.g. Langella 2001), (3) soil and surficial deposits (e.g.

Ming and Mumpton 1989), (4) deep marine sediments (e.g. Boles and Coombs 1977)

and (5) marine sediments from arc-source terrains (e.g. Boles and Coombs 1977) (Fig.

1.1).

Fig. 1.1: Schematic diagrams showing patterns of zeolite zoning in silicic tephra deposits in various genetic environments (modified after Hay 1977 and Neuhoff et al. 2000). (a) Plan view and cross-section view of zeolites formed in closed hydrological systems (e.g. Playa lakes). The mineral distribution in these systems reflects an increase in salinity during fluid evaporation. (b) Cross-section view of zeolites formed in an open hydrological system. Mineral distribution is taken to reflect pH changes during progressive interaction with the host rock. (c) Cross-section view of mineral distribution in hydrothermal systems. Mineral distribution reflects a temperature gradient during alteration. (d) Cross-section view of mineral distribution under ongoing burial of a stratigraphic sequence. Mineral distribution reflects a temperature gradient during burial.

Almost every known zeolite occurs in cavities of volcanic lava flows (e.g. Tertiary

lavas of Iceland, Deccan Plateau, India). These zeolites are formed either during

burial metamorphism of the lava pile (e.g. Walker 1960; Neuhoff 1999), during

INTRODUCTION 9

hydrothermal alteration of continental basalts (e.g. Walker 1963), or during diagenesis

in areas of high heat flow caused by active geothermal systems (e.g. Kristmannsdóttir

and Tómasson 1978; Weisenberger and Selbekk 2008).

Zeolites as products of hydrothermal crystallization are generally known from

active geothermal systems associated with volcanic rocks. Very little work has been

published on zeolite occurrences related to late stage crystallization of pegmatitic

bodies (e.g. Orlandit and Scortecci 1985), in hydrothermal ore veins (Deer et al.,

2004), as alteration along fault plains (e.g. Vincent and Ehlig 1988), and in

hydrothermal fractures and veins in granites and gneisses (e.g. Borchardt et al. 1990;

Borchardt and Emmermann 1993; Armbruster et al. 1996; Bish and Ming 2001;

Freiberger et al. 2001; Fujimoto et al. 2001).

Fig. 1.2: Temperature-pressure diagram showing the metamorphic facies, including the field of zeolite facies which represents the lowest metamorphic facies (modified after Winter 2001).

The importance of zeolites as low temperature alteration phases in the Earth’s crust

was noted early in the history of metamorphic petrology. However, Eskola (1939)

rejected the concept of a metamorphic zeolite facies. The issue was revisited in the

1950s when Rengarten (1950) proposed a “geochemical zeolite facies” in which

zeolite assemblages represent alteration of sediments in contact with aqueous

solutions of unusual composition. The benchmark papers of Fyfe et al. (1958) and

INTRODUCTION 10

Coombs et al. (1959) arise the present concept of zeolite facies as a necessary

intermediate metamorphic grade between diagenesis and greenschist facies. Since

then several zeolite facies distributions in various genetic environments were

described. Nowadays the zeolite facies is accepted as an intermediate facies between

the prehnite-pumpellyite facies and diagenesis (Fig. 1.2, 1.3)

The distribution of individual zeolite species within diagenetically altered or

metamorphosed sediments and volcanic rocks is commonly characterized by isograds

of first appearance of one or more zeolites bounding spatially restricted zones (e.g.

Coombs et al. 1959; Hay 1977; Kristmannsdóttir and Tómasson 1978).

Fig. 1.3: Temperature-pressure ranges of zeolite-forming environments. (adapted from Deer et al. 2004). Solid curves are experimentally determined stability limits of selected zeolites: (1) epidote + quartz + H2O = laumontite + prehnite, at low-pressure end and epidote + chlorite + quartz + H2O = laumontite + pumpellyite at high-pressure end (2) laumontite + quartz + H2O = heulandite (Cho et al. 1987). Dashed line represents a retrograde PT-path in Alpine fissures, determined by fluid inclusions in fissure quartz (Mullis et al. 1994).

The golden spike for zeolite facies mineralization was done by the British geologist

George Walker (1959, 1960, 1963). In his work, Walker made a careful study of

Iceland´s amygdules and mapped the zeolite distribution in Tertiary lavas of Eastern

Iceland. He recognized a systematic depth variation in the zeolite distribution of the

lava sequences. Similar observations were done during the same period by Coombs et

al. (1959) on low-grade meta-sediments in New Zealand.

INTRODUCTION 11

With their pioneering work Walker (1959, 1960) and Coombs et al. (1959)

identified regionally extensive mineral assemblages that define “depth” controlled

“zeolite zones” formed during burial metamorphism that have been found in different

environments during the last centuries (e.g. Bish and Ming 2001).

1.6. ZEOLITE STABILITY

The particular zeolite to form depends on any of five factors: (1) temperature, (2)

pressure, (3) primary rock composition, (4) fluid composition, and (5) the water to

rock ratio. The most important factor of all theses for a particular paragenesis is the

composition of the material to be altered and the composition of the altering solution

(e.g. Deer et al. 2004).

The effect of pressure on zeolite isograds usually cannot be assessed

independently from that of temperature; however, Iijima (1988) has demonstrated that

the temperatures corresponding to the isograds are essentially independent of

pressure.

Several reactions involving Ca-zeolites have been studied experimentally (e.g.

Liou 1971; Thompson 1971; Cho et al. 1987; Frey et al. 1991). In general the

maximum temperature and pressure limits of zeolite stability are in agreement with

observations on geothermal systems (Kristmannsdóttir and Tómasson 1978; Frey et

al. 1991). Above 400°C anorthite is stable relative to wairakite, the Ca-zeolite stable

at highest temperature (Frey et al. 1991).

There are large discrepancies between directly measured temperatures at the

position of some zeolite boundaries in boreholes and temperature calculated from

experimentally based thermodynamic data (Neuhoff et al. 2000). For example, in the

North Tejon oil field in California, heulandite and laumontite coexist at around 90 °C

(Noh and Boles 1993). This is a much lower temperature than the 240 ˚C derived

from equilibrium phase calculation. Neuhoff (1999) has demonstrated that these

discrepancies can be attributed for example to variations in structural order-disorder

and in different chemical composition between natural and synthetic minerals.

INTRODUCTION 12

1.7. GEOLOGICAL SETTING

The Alps form a part of a Tertiary orogenic belt that stretches from southern Europe

to Asia. The Alps formed as a result of the closure of Jurassic to Cretaceous Tethys

ocean basins during convergence of the Apulian and European plates (e.g. Trümpy

1960; Frisch 1979; Schmid et al. 2004). An orogenic belt characterized by stacked

nappes formed in the Tertiary when the Apulian and European plates collided. The

collision caused a complicated tectonic structure and a regional metamorphic

overprint. Deeply buried parts of the orogen were later exhumed and uplifted in the

late Tertiary (Trümpy 1980) and finally reached the erosion surface.

Fig. 1.4: Simplified geological map of Switzerland (modified after Spicher 1980). The dashed line mark the trail of the Gotthard NEAT tunnel.

INTRODUCTION 13

Zeolites in fissures occur predominantly in rocks that belong to large basement

windows exposed in the northern part of the Alps (Fig. 1.4). These so-called “external

massifs” of the Alps belongs to the European plate (e.g. Trümpy 1980). The massifs

represent parautochthonous units (Pfiffner 1986). Two major basement units are

distinguished in the central Swiss Alps: the Aar Massif and the Gotthard Massif. They

constitute of pre-Variscan basement, which is partly reworked by the Variscan and

Alpine orogenesis. The massifs form a 115 km long and 23 to 40 km wide SW-NE

trending outcrop. The large Aar Massif consists of pre-Variscan gneisses, pre-

Variscan granitoids, migmatitic granites and gneisses, lower and upper Carboniferous

intrusives and Carboniferous volcanics (Abrecht 1994). Many of the prominent high

Alpine peaks and the largest glaciers of the Alps are located in the Aar massif.

The Gotthard Massif is located to the south of the Aar Massif and is followed

further south by the north Penninic continental nappe stack. It consists of a poly-

metamorphic continental basement with Variscan granites. It is separated from the

Aar Massif in the north by the narrow Tavetsch Massif and the Mesozoic

metasediments of the Urseren zone.

The rocks have been overprinted by Tertiary Alpine metamorphism. The

metamorphic grade and Alpine peak metamorphism increases from nearly non-

metamorphosed rock units in the north, over greenschist facies rocks in the Aar- and

Gotthard Massif region up to amphibolite facies conditions in the Penninic nappes to

the south (Labhart 1977; Frey et al. 1980; Frey and Mählmann 1999). In general,

metamorphic grade increases within a specific tectonic unit from north (external

position) to south (internal position), as well as from the top to the bottom (Frey and

Mählmann 1999) (Fig. 1.5).

Deutsch and Steiger (1985) determined an age of 37 Ma for peak metamorphism

in the Lepontine Alps, whereas peak metamorphism in the Aar Massif is dated around

25 Ma (Grimsel area, Dempster 1986).

During uplift and erosion brittle deformation structures became dominant once

the rocks crossed the ductile-brittle transition zone. Semi-brittle shear zones related to

backthrusting, normal faults and the opening of fissures and gashes are typical

structures related to the late orogenic deformation phases including extension

structures related to uplift.

INTRODUCTION 14

The fissures are conductive to hot aqueous fluids, whereas the percolating hot

fluids could react with the surrounding rocks along the fissure walls to form

secondary fissure minerals.

Fig. 1.5: Schematic sketch of the T-t evolution of tectonic units in the Central Alps in relation to fissure formation and the timing of zeolite growth. (a) The T-t paths of individual tectonic units reflecting an increase of the Alpine peak metamorphism from north to south. The southern units have reached the ductile regime, northern units were deformed brittle. All units reached temperatures above the zeolite window (except for the parautochtonous cover rocks of the Aar massif). During uplift the units returned to the brittle deformation regime and extension fissures formed (b), subsequently zeolite-absent fissure assemblages developed (c) and finally the units entered the zeolite window (d).

Metamorphic conditions derived from fluid inclusion studies on Alpine fissure

material represent conditions during various stages of exhumation, decompression

and cooling. Mullis et al. (1994) determined minimum conditions for fissure mineral

formation ranging from 400-430°C in temperature and fluid pressures from 240 to

380 MPa for the southern Aar Massif (Zinggenstock), for the Gotthard Massif and the

northern Lepontine Alps.

An exhumations rate of 0.5 mm a-1 for the Reuss valley (northern Aar Massif) has

been proposed by Michalski and Soom (1990) for the past 27 Ma from apatite and

zircon fission track data. The corresponding cooling rate of 13°C Ma-1 agrees well

with other apatite fission track data that suggest uplift rates of 0.3-0.6 mm a-1 during

the last 6-10 Ma (Schaer et al. 1975). Using exhumation rates and trapping

temperatures of early fluid inclusions (Mullis 1996) determined the time of the first

INTRODUCTION 15

opening of fissures and precipitation of fissure minerals in the Aar Massif

(Zinggenstock) and Gotthard Massif (La Fibbia) to 20 to 15 Ma b.p..

The deposition of fissure minerals occured along the cooling and decompression

P-T-path after peak metamorphism (Fig. 1.5).

1.8. REFERENCES

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Armbruster, T., Kohler, T., Meisel, T., Nägler, T.F., Götzinger, M.A., and Stalder, H.A. (1996) The zeolite, fluorite, quartz assemblage of the fissure at Gibelsbach, Fiesch (Valais, Switzerland): crystal chemistry, REE patterns, and genetic speculations. Schweizerische Mineralogische und Petrographische Mitteilungen, 76, 131-146.

Bish, D.L., and Ming, D.W. (2001) Natural Zeolites: occurrence, properties, applications, 45, 654 p. Mineralogical Society of America, Geochemical Society.

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Cho, M., Maruyama, S., and Liou, J.G. (1987) An experimental investigation of heulandite-laumontite equilibrium at 1000 to 2000 bar Pfluid. Contributions to Mineralogy and Petrology, 97, 43-50.

INTRODUCTION 16

Ciesielczuk, J., and Janeczek, J. (2004) Hydrothermal alteration of the Strzelin granite, SW Poland. Neues Jahrbuch für Mineralogie, Abhandlungen, 179, 239-264.

Colella, C. (1996) Ion exchange equilibria in zeolite minerals. Mineralium Deposita, 31, 554-562.

Coombs, D.S., Alberti, A., Artioli, A., Armbruster, T., Colella, C., Galli, E., Grice, J.D., Liebau, F., Mandarino, J.A., Minato, H., Nickel, E.H., Passaglia, E., Peacor, D.R., Quartieri, S., Rinaldi, R., Ross, M., Sheppard, R.A., Tillmanns, E., and Vezzalini, G. (1998) Recommended nomenclature for zeolite minerals: report of the subcommittee on zeolites of the international mineralogical association, commission on new minerals and mineral names. Mineralogical Magazine, 62, 533-571.

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Cronstedt, A.F. (1756) Om en obekant barg art, som kallas Zeolites. Akad. Handl. Stockholm, 17, 120-123.

Deer, W.A., Howie, R.A., Wise, W.S., and Zussman, J. (2004) Framework silicates: silica minerals, feldspathoids and the zeolites. 982 p. The Geological Society, London.

Dempster, T.J. (1986) Isotope systematics in minerals: biotite rejuventation and exchange during Alpine metamorphism. Earth and Planetary Science Letters, 78, 355-367.

Deutsch, A., and Steiger, R. (1985) Hornblende K-Ar ages and the climax of Tertiary metamorphism in the Lepontine Alps (South-Central Switzerland); an old problem reassessed. Earth and Planetary Science Letters, 72, 175-186.

Eskola, P. (1939) Die Metamorphen Gesteine. In T. Barth, F.W. Correns, and P. Eskola, Eds. Die Entstehung der Gesteine, p. 263-407. Springer, Berlin.

Freiberger, R., Hecht, L., Cuney, M., and Morteani, G. (2001) Secondary Ca-Al silicates in plutonic rocks: implications for their cooling history. Contributions to Mineralogy and Petrology, 141, 415-429.

Frey, M., Bucher, K., Frank, E., and Mullis, J. (1980) Alpine metamorphism along the Geotraverse Basel-Chiasso - a review. Eclogae Geologicae Helvetiae, 73, 527-546.

Frey, M., de Capitani, C., and Liou, J.G. (1991) A new petrogenetic grid for low-grade metabasites. Journal of Metamorphic Geology, 9, 497-509.

Frey, M., and Mählmann, R.F. (1999) Alpine metamorphism of the Central Alps. Schweizerische Mineralogische und Petrographische Mitteilungen, 79, 135-154.

Frisch, W. (1979) Tectonic progradation and plate tectonic evolution of the Alps. Tectonophysics, 60, 121-139.

Fujimoto, K., Tanaka, H., Higuchi, T., Tomida, N., Othani, T., and Ito, H. (2001) Alteration and mass transfer inferred from the Hirabayshi GSJ drill penetrating the Nojima Fault, Japan. Island Arc, 10, 401-410.

INTRODUCTION 17

Fyfe, W.S., Turner, F.J., and J., V. (1958) Metamorphic reactions and metamorphic facies. 259 p.

Gottardi, G. (1989) The genesis of zeolites. European Journal of Mineralogy, 1, 479-487.

Gottardi, G., and Galli, E. (1985) Natural Zeolites. 409 p. Springer, Berlin, Heidelberg, New York, Tokyo.

Hauri, F. (2006) Natural zeolites from southern Germany: Applications in concrete. In R.S. Bowmann, and S.E. Delap, Eds. Zeolite'06 - 7th International Conference on the Occurrence, Properties, and Utilization of Natrural Zeolites, p. 130-131, Socorro, New Mexico, USA.

Hay, R.L. (1966) Zeolites and zeolitic reactions in sedimentary rocks. Geological Society of America, Special Papers 85.

Hay, R.L., and Sheppard, R.A. (1977) Zeolites in open hydrologic systems. In F.A. Mumpton, Ed. Mineralogy and geology of Natural Zeolites, 4, p. 93-102. Mineralogical Society of Amercia, Short Course Notes.

Hay, R.L., and Sheppard, R.A. (2001) Occurrences of zeolites in sedimentary rocks. In D.L. Bish, and D.W. Ming, Eds. Natural Zeolites: occurrence, properties, applications, 45, p. 217-234. Mineralogical Society of America, Geochemical Society.

Huber, H.M. (1943) Die Kluftminerallagerstätten im südöstlichen Gotthardmassiv. Schweizerische Mineralogische und Petrographische Mitteilungen, 23, 475-537.

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Kalló, D. (2001) Applications of natural zeolites in water and wastwater treatment. In D.L. Bish, and D.W. Ming, Eds. Natural Zeolites: occurrence, properties, applications, 45, p. 519-550. Mineralogical Society of America, Geochemical Society.

Kenngott, A. (1866) Die Mineralien der Schweiz nach ihren Eigenschaften und Fundorten ausführlich beschrieben. 460 p. Verlag W. Engelmann, Leipzig.

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Langella, A., Cappelletti, P., and de'Gennaro, M. (2001) Zeolites in closed hydrologic systems. In D.L. Bish, and D.W. Ming, Eds. Natural Zeolites: occurrence, properties, applications, 45, p. 235-260. Mineralogical Society of America, Geochemical Society.

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INTRODUCTION 18

apatite and K-Ar mica ages. Schweizerische Mineralogische und Petrographische Mitteilungen, 70, 373-387.

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INTRODUCTION 19

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ZEOLITES IN BASEMENT ROCKS 20

2. ZEOLITES IN FISSURES OF GRANITES AND

GNEISSES OF THE CENTRAL ALPS


2.1. ABSTRACT

Six different Ca-zeolites occur widespread in various assemblages in late fissures and

fractures in granites and gneisses of the Swiss Alps. The zeolites form as a result of

water-rock interaction at relatively low temperatures (<250 °C) in the upper

continental crust. The low-grade fissure mineral assemblages are the key to the

appreciation of water-rock interaction in hydrothermal and geothermal systems

located in granites and gneisses of the crystalline basement. The zeolites typically

overgrow earlier minerals of the fissure assemblages, but zeolites also occur as single

stage fissure deposits in granite and gneiss. They represent the most recent fissure

minerals formed during uplift and exhumation of the Alpine orogen. A systematic

study of zeolite samples showed that the majority of finds originate from three regions

particularity rich in zeolite-bearing fissures: (1) in the central and eastern part of the

Aar- and Gotthard Massif, including the Gotthard road tunnel and the Gotthard-

NEAT tunnel, (2) Gibelsbach/Fiesch, in a fissure breccia between Aar Massif and

Permian sediments, and (3) in Penninic gneisses of the Simano nappe at Arvigo (Val

Calanca).

The excavation of tunnels in the Aar- and Gotthard massif give an excellent

overview of zeolite frequency in Alpine fissures, whereas 32 % (Gotthard NEAT

tunnel, 12000-18555) and 18 % (Gotthard road tunnel) of all fissures are filled with

zeolites. The number of different zeolites is limited to 6 species: laumontite, stilbite

and scolecite are abundant and common, whereas heulandite, chabazite and epistilbite

occur occasionally. Ca is the dominant extra-framework cations, with minor K and

Na. Heulandite and chabazite additionally contain Sr up to 29 and 10 mole%,

respectively. Na and K content of zeolites tends to increase during growth as a result

of systematic changes in fluid composition and/or temperature. The K enrichment of

stilbite found in surface outcrops compare to stilbite in the subsurface may indicate

late cation exchange during interaction with surface water. Texture data, relative age

sequences derived from fissure assemblages and equilibrium calculations shows that

the Ca-dominated zeolites precipitated from fluid with decreasing temperature in the

order (old to young = hot to cold): scolecite, laumontite, heulandite, chabazite and

stilbite.


The components necessary for zeolite formation are derived from dissolving

primary granite and gneiss minerals. The nature of these minerals depends on the

metamorphic history of the host rock. Zeolites in the Aar Massif derived from the

dissolution of epidote or calcite and albite that were originally formed during Alpine

greenschist metamorphism. Whereas albitization of plagioclase in higher grade rocks

releases the necessary components for zeolite formation, a process that is

accompanied by a distinct porosity increase. Zeolite fissures occur in the zone where

fluid inclusions in earlier formed quartz contain H2O dominated fluids. This is

consistent with equilibrium calculations that predict a low CO2 tolerance of zeolite

assemblages particularly at low temperature. Pressure decrease along the uplift and

exhumation can increase zeolite stability. The major zeolite forming reaction

consumes calcite and albite; it increases pH and the total of dissolved solids. The

produced Na2CO3 waters are in accord with reported deep groundwater (thermal

water) in the continental crust, which are typically oversaturated with respect to Ca-

zeolites.

Keywords: zeolite, granite, water-rock interaction, laumontite, Swiss Alps

2.2. INTRODUCTION

The interaction of rocks with hot water circulating on the fractures of the continental

crust produces fissure minerals in the open porosity of the fissures and leaches the

original rock matrix. The chemical composition of the fluid thereby monitors the

water-rock interaction process that controls the dissolution of primary minerals, as

well as the precipitation of secondary minerals in the open spaces (e.g. Nordstrom et

al., 1989; Stober & Bucher, 1999; Bucher et al., 2009). Veins and mineralized

fractures are ubiquitous in regional metamorphic terrains. They bear considerable

information on fluid movement, fluid-rock interaction and fluid sources (e.g. McCaig

et al., 1990). Detailed mineralogical and petrological study of the low-grade fissure

mineral assemblage provides quantitative access to fluid-rock interaction. From such

data the evolution of porosity and permeability of the total system and the leached

rock matrix can be deduced. These flow properties of fractured rocks are required for

the understanding of geothermal systems and fluid migration in the upper continental


crust (e.g. Gianelli et al., 1998; Neuhoff et al., 1999; Weisenberger & Selbekk, 2008).

The crystalline basement of the upper continent crust consists predominantly of

granites and gneisses. A major mineral of both rock types is plagioclase.

Hydrothermal alteration of basement rocks along fractures attacks predominantly

plagioclase and replaces the mineral with secondary Ca-Al silicates such as epidote,

prehnite and zeolites. Zeolites are the predominant secondary Ca-Al mineral at low

temperature (< 250 °C) (e.g. Gottardi, 1989; Bish & Ming, 2001; Fig. 2.1).

Although zeolites are known from fissures and gashes of the crystalline basement

from reports by mineral collectors for more than 150 years (e.g. Kenngott, 1866;

Parker, 1922; Niggli et al., 1940; Huber, 1943; Sigrist, 1947; Stalder et al., 1998)

(Table 2.1), they did not create much interest in the scientific research community.

Previous publications on zeolite occurrences and their origin in the Central Swiss

Alps and other areas of crystalline basement rocks are limited (Armbruster et al.,

1996; Freiberger et al., 2001; Fujimoto et al., 2001; Ciesielezuk & Janeczek, 2004).

Fig. 2.1: Pressure-temperature ranges of environments of zeolite formation. (adapted from Deer et al., 2004). Solid curves are experimentally determined stability limits of selected zeolites: (1) epidote + quartz + H2O = laumontite + prehnite, at low-pressure end and epidote + chlorite + quartz + H2O = laumontite + pumpellyite at high-pressure end (2) laumontite + quartz + H2O = heulandite (Cho et al., 1987). Dashed line represents a retrograde PT-path in Alpine fissures, determined by fluid inclusions in fissure quartz (Mullis et al., 1994).


In general, zeolites are among the most common products of chemical interaction

between fluids and the crustal rocks during diagenesis and low-grade metamorphism

(e.g. Bish & Ming, 2001). Zeolite minerals occur in low temperature (<250 °C), low

pressure (<200 MPa), water saturated environments. The required constituents for the

formation of zeolites are commonly derived from dissolution of volcanic glass (e.g.

Sheppard & Hay, 2001) and zeolites are common in altered volcanic rocks. In granites

and gneisses, the necessary components can be derived from the alteration of

feldspars, particularly plagioclase, and other aluminous silicates (Engvik et al., 2008).

Temperature and pressure largely control the kind of zeolite that will be formed. P

and T is usually a function of burial depth or temperature changes during

hydrothermal overprint. The composition of the altered material and the composition

of the hydrothermal fluid are further controls on the product zeolite mineralogy (e.g.

Bucher & Stober, 2001; Deer et al., 2004).

Zeolites as products of hydrothermal crystallization are generally know from

active geothermal systems associated with volcanic rocks. Very little work has been

published on zeolite occurrences related to late stage crystallization of pegmatitic

bodies (e.g. Orlandit & Scortecci, 1985), in hydrothermal ore veins (Deer et al.,

2004), as alteration along fault plains (e.g. Vincent & Ehlig, 1988), and in

hydrothermal fractures and veins in granites and gneisses (e.g. Borchardt et al., 1990;

Borchardt & Emmermann, 1993; Armbruster et al., 1996; Bish & Ming, 2001;

Freiberger et al., 2001; Fujimoto et al., 2001).

The scarcity of reports on zeolite in fissures of granites and gneisses is

astonishing because of the obvious very widespread occurrence of zeolites in granites

and gneisses. For example in the latest special edition on natural zeolites (Bish &

Ming, 2001), only two short notes were made on zeolites hosted in granites and

gneisses. Nevertheless zeolite formation in fractures and cavities in granitic gneisses

is a frequent feature in the continental crust.

In this paper, we present the “uncommon” zeolite occurrences in fractures and

cavities in granites and granitic gneisses in the Central Swiss Alps. The presented data

allow for a consistent model of zeolite formation in basement rocks (Table 2.2; Fig.

2.2).



The Alps forms a part of a Tertiary orogenic belt that stretches from southern Europe

to Asia. The Alps formed as a result of the closure of Jurassic to Cretaceous Tethys

ocean basins during convergence of the Apulian and European plates (e.g. Trümpy,

1960; Frisch, 1979; Schmid et al., 2004). An orogenic belt characterized by stacked

nappes formed in the Tertiary when the Apulian and European plates collided. The

collision caused a complicated tectonic structure and a regional metamorphic

overprint. Deeply buried parts of the orogen were later exhumed and uplifted in the

late Tertiary (Trümpy, 1980) and finally reached the erosion surface.

Fig. 2.2: Map of Switzerland and the Central Swiss Alps. (a) Outline of Switzerland and the position of the external massifs (modified after Labhart, 1977). Numbered points marks zeolites localities and the numbers corresponds to Table 2.1. The positions of the Gotthard road tunnel and the Gotthard NEAT tunnel are marked central external massifs in gray. Black star marks the position of the zeolite locality Arvigo, Val Calanca/GR. (b) Simplified geological map with dashed line. (c) Spatial distribution of each zeolite, showing no preferred distribution.

Zeolites in fissures occur predominantly in rocks that belong to large basement

windows exposed in the northern part of the Alps. These so-called “external massifs”

of the Alps thus belong to the European plate (e.g. Trümpy, 1980). The massifs

represent parautochthonous units (Pfiffner, 1986). Two major basement units are

distinguished in the Central Swiss Alps: the Aar Massif and the Gotthard Massif.

They constitute to the pre-Variscan basement, which is partly reworked by the


Variscan and Alpine orogenesis. The massifs form a 115 km long and 23 to 40 km

wide SW-NE trending outcrop. The large Aar Massif consists of pre-Variscan

gneisses, pre-Variscan granitoids, migmatitic granites and gneisses, lower and upper

Carboniferous intrusives and Carboniferous volcanics (Abrecht, 1994). Many of the

prominent high Alpine peaks and the largest glaciers of the Alps are located in the

Aar massif.

The Gotthard Massif is located to the south of the Aar Massif and is followed

further south by the north Penninic continental nappe stack. It consists of a poly-

metamorphic continental basement with Variscan granites. It is separated from the

Aar Massif in the north by the narrow Tavetsch Massif and the Mesozoic

metasediments of the Urseren zone (Fig. 2.2).

2.3.1. Metamorphic conditions during Alpine orogenesis

All rocks have been overprinted by the Tertiary Alpine metamorphism. The

metamorphic grade and Alpine peak metamorphism increases from nearly non-

metamorphosed rock units in the north, over greenschist facies rocks in the Aar- and

Gotthard Massif region up to amphibolite facies conditions in the Penninic nappes to

the south (Labhart, 1977; Frey et al., 1980; Frey & Mählmann, 1999). In general,

metamorphic grade increases within a specific tectonic unit from north (external

position) to south (internal position), as well as from the top to the bottom (Frey &

Mählmann, 1999).

Metamorphic conditions of the Central Alps exceed the zeolite facies with the

exception of the parautochthonous units in the north. A continuous zone of very low-

grade metamorphism (= anchizone) of up to 15 km width can be delineated along the

southern Helvetic nappes and in the parautochthonous cover of the Aar Massif

basement (Frey & Mählmann, 1999). Peak metamorphism determined from the

Taveyanne greywacke (Glarus Alps) in parautochthonous units north of the Aar

Massif range from zeolite (Lmt + Prh + Pmp + corrensite), to prehnite-pumpellyite

(Prh + Pmp + Ep), pumpellyite-actinolite (Pmp + Act + Ep) and lower greenschist

facies (Act + Ep) (Rahn et al., 1994). Metamorphic conditions in parautochthonous

units north of the Aar Massif corresponds to a temperature range of 240-300 °C and

200-300 MPa (Frey & Mählmann, 1999; Rahn et al., 1994).


Fig. 2.3: Detailed geological map of the eastern Aar Massif (modified after Labhart, 1977). Numbered points marks zeolites localities and the numbers corresponds to Table 2.1. G = glacier

Tertiary greenschist facies metamorphism (= epizone) overprinted the old basement

rocks of the Aar- and Gotthard Massif. The following isograds have been located

from north to south: (1) first appearance of green biotite (Steck & Burri, 1971), (2)

disappearance of stilpnomelane (Jäger et al., 1967), (3) the transformation isograd of

microcline/sanidine (Bambauer & Bernotat, 1982; Bernotat & Bambauer, 1982; Frey

& Mählmann, 1999). This corresponds to maximum temperatures in the Northern Aar

Massif of about 270 °C. This temperature is above the zeolite window defined by data

on illite crystallinity, vitrinite reflection and fluid inclusion measurements by

Breitschmid (1982). The typical Alpine mineral assemblages of the Central Aar

granite (Fig. 2) is Qtz + Ab + Kfs + Chl + Ms + Cal. Alpine green biotite occurs along

a mappable isograd (Steck & Burri, 1971) in the Aar granite suggesting a minimum

temperature at about 420 °C. The microcline/sanidine transformation isograd is

located further south in the Aar granite, suggesting a further increase of metamorphic

grade with minimum temperatures of 450 °C (Bambauer & Bernotat, 1982). The first

appearance of oligoclase in granitic gneisses in the Gotthard Massif (Steck, 1976) still

further south marks the beginning of amphibolite facies conditions (about 500 ˚C).

Metamorphic peak conditions regularly increase southward in the Lepontine Alps

(the Penninic nappe stack). The stacked nappes involve staurolite-bearing micaschists,

tremolite marbles and amphibolites all characteristic of amphibolite facies conditions


(Frey & Mählmann, 1999). Peak conditions during Eocene to Miocene metamorphism

at the Arvigo zeolite locality (Fig. 2.2) in the Simano nappe (a member of the

Penninic nappe stack range from 600-680 °C and 550 to 600 MPa (Engi et al.; 1995;

Todd & Engi, 1997; Nagel et al., 2002). Deutsch & Steiger (1985) determined an age

of 37 Ma for peak metamorphism in the Lepontine Alps, whereas peak metamorphism

in the Aar Massif is dated around 25 Ma (Grimsel area, Dempster, 1986).

During uplift and erosion brittle deformation structures became dominant once

the rocks crossed the ductile-brittle transition zone. Semi-brittle shear zones related to

backthrusting, normal faults and the opening of fissures and gashes are typical

structures related to the late orogenic deformation phases including extension

structures related to uplift. The fissures can be divided into two characteristic groups

based on geometry and morphology (Mullis et al., 1994; Mullis, 1995): (1) tension

gashes and (2) interboudin gaps. Tension gashes generally develops parallel to the

maximum stress (σ1) and perpendicular to the maximum elongation, at an angle of

around 45° to the shear plane (Huber, 1948; Mullis et al., 1994). They are mostly

arranged in en echelon fashion (Ramsay, 1967). The variation in length reach from

<10 cm up to more than 10 meters. Interboudin gaps usually develop parallel to the

direction of maximum extension in rocks of different viscosity. In contrast to tension

gashes the interboudin gaps rarely exceed the size of 1 meter. The orientation of both

types of fissures is normal to foliation or schistosity (Mullis et al., 1994).

Both structures are conductive to hot aqueous fluids. The percolating hot fluids

could react with the surrounding rocks along the fissure walls to form secondary

fissure minerals.

Metamorphic conditions derived from fluid inclusion studies on Alpine fissure

material represent conditions during various stages of exhumation, decompression

and cooling. Mullis et al. (1994) determined minimum conditions for fissure mineral

formation ranging from 400-430 °C in temperature and fluid pressures from 240 to

380 MPa for the southern Aar Massif (Zinggenstock), for the Gotthard Massif and the

northern Lepontine Alps.

An exhumations rate of 0.5 mm a-1 for the Reuss valley (northern Aar Massif) has

been proposed by Michalski & Soom (1990) for the past 27 Ma from apatite and

zircon fission track data. The corresponding cooling rate of 13 °C Ma-1 agrees well


with other apatite fission track data that suggest uplift rates of 0.3-0.6 mm a-1 during

the last 6-10 Ma (Schaer et al., 1975). Using exhumation rates and trapping

temperatures of early fluid inclusions (Mullis, 1996) determined the time of the first

opening of fissures and precipitation of fissure minerals in the Aar Massif

(Zinggenstock) and Gotthard Massif (La Fibbia) to 20 to 15 Ma b.p..

The deposition of fissure minerals, hence also the zeolite minerals described in

this paper, occured along the cooling and decompression P-T-path after peak

metamorphism.

2.4. SAMPLING AND ANALYTIC METHODS

Samples of fissure minerals from surface outcrops and from road and rail tunnels in

the Central Swiss Alps provided assemblage data, relative age relationships and

chemical composition data from zeolites and associated minerals. Samples were

supplied from four sources (Table 2.2): (1) The Swiss Natural History Museums Bern

(SNHMB) made the vast collection of Alpine minerals available for our systematic

study of the spatial distribution of zeolites, the museum keeps samples of important

zeolite localities, including the Gotthard road tunnel. (2) Excellent mineral samples

from the Amsteg-Sedrun section of the new Gotthard rail base tunnel currently under

construction were made available by the mineral representative Peter Amacher. He

saved the specimens during excavation work during the last few years. (3) The

Mineralogical Museum University Freiburg provided samples from classic localities.

(4) During several field trips in the summer seasons 2006 and 2007 mineral samples

were collected from the Aar- and Gotthard Massif, Gibelsbach and Arvigo/Val

Calanca.

Quantitative zeolite analysis were performed at the Institute of Mineralogy and

Geochemistry, University of Freiburg, using a CAMECA SX 100 electron

microprobe equipped with five WD spectrometers and one ED detector with an

internal PAP-correction program (Pouchou & Pichior, 1991). Major and minor

elements for zeolites were determined at 15 kV accelerating voltage and 8 nA beam

current with a defocused electron beam of 20 µm in diameter with counting time up to


20 s. Na and K were counted first to minimize the Na and K loosed during

determination. Since zeolites lose water when heated, the crystals were mounted in

epoxy resin to minimize loss of water due to the electron bombardment. Natural and

synthetic standards were used for calibration. The standards employed were: albite

(Na), periclase (Mg), wollastonite (Si), barite (Ba), hematite (Fe), celestine (Sr),

orthoclase (K), anorthite (Ca), rhodonite (Mn) and rutile (Ti). Identification of various

minerals was obtained by a BRUKER AXS D8 Advance X-ray powder diffractometer

(XRD) and the DIFFRACplus v5.0 software for evaluation.

The content of zeolite water was determined by heating the samples that had been

equilibrated with air of 50 % relative humidity to 873 K for 24 h and measuring the

weight loss. For some samples zeolite water was determined by measuring mass loss

between 298 and 1273 K by scanning-heating TGA (thermogravimetry analysis) at a

heating rate of 10 K min-1 on a Netzsch STA 449C Jupiter simultaneous DSC-TGA

(differential scanning calorimetry - thermogravimetry analysis) apparatus. The charge

balance of zeolites formulas is a reliable measure for the quality of the analysis and

which correlates with the difficulties related to the thermal instability of zeolites in

microprobe analysis. A usefull error test investigates the charge balance between the

non-framework cations and the amount of tetrahedral Al (Passaglia, 1970). Analyses

are considered acceptable if the sum of the charge of the extra-framework cations

(Ca2+, Sr2+, Na+, and K+) is within 10% of the framework charge (Al3+).

2.5. ZEOLITES IN THE CENTRAL ALPS

In contrast to other zeolite environments (e.g. basalts, alkaline lakes) only few

different zeolite species occur in granites and gneisses. Laumontite, scolecite,

heulandite, stilbite and chabazite mark the dominant species whereas epistilbite

occurs only at three sites (Table 2.1; no 4, 25 & 57). Additionally thomsonite,

phillipsite, natrolite, mesolite and analcime were found in the Central Alps associated

with rocks of basic to ultramafic composition (e.g. amphibolites, metabasalts,

serpentinites), e.g. within the Zermatt-Saas ophiolite (Table 2.1; no 67) or the

Geisspfad (Binntal) serpentinite (Table 2.1; no 54) and the zeolite-bearing Taveyanne

greywacke (Glarus Alps, Rahn et al., 1994). These zeolite occurrences will not be


discussed here. Faujasite was mentioned by Parker (1922) in the eastern part of the

Central Aar granite body (Fig. 2.2) but the mineral has never been confirmed.

2.5.1. Spatial distribution

The spatial distribution of zeolites in Alpine fissures in crystalline basement rocks of

the Central Alps is summarized in Figs 2.2 & 2.3 and Table 2.1. It follows from the

data compilation that zeolites are not evenly distributed but rather have preferentially

been reported from a relatively small number of localities. In a broad zone

surrounding these localities zeolites are very common and widespread. These focus

areas typically cover some km2. It is important to note that the patchy zeolite

distribution does not result from sampling bias. Sampling bias can be excluded

because the main target of the mineral hunters is rock crystal and smoky quartz,

which is found in fissures over the entire area. However, quartz-bearing fissures are

commonly devoid of zeolite minerals.

Fig. 2.4: Photograph and schematic illustration showing typical vein characteristics of Alpine fissures. (a) Photograph of a vein hosted in a biotite-rich gneiss (Arvigo/Val Calanca); hammer for scale. The vein is characterized by a 1 cm leaching zone trending in vertical direction, which appears to be lighter colored, due to the removal of mafic minerals. The open space of the fissure is filled with secondary minerals, mainly chlorite. (b) Schematic sketch of an Alpine fissure. The open fissures provide pathways for hot fluids. Leaching caused by fluid-rock interaction change primary mineralogy and composition of the host rock, visible as alteration zone along to the fissure wall. Secondary minerals precipitate in the open space. (c) Schematic sketch of a zeolite bearing Alpine fissure, exhibit euhedral mineral assemblages. Zeolites overgrow earlier formed minerals in the following order, as it observed in nature: Qtz → Ep → Prh → Sco. The alterations zones seem to be proportional to the aperture of fissures.

One of the prime sources of Alpine fissure quartzes is the Central Aar granite (Figs

2.2 & 2.3), which forms an over 100 km long and 8 to 10 km wide intrusive body


striking in SW-NE direction. It contains abundant fissure zeolites only in an eastern

area centering around Piz Giuv (Figs 2.2 & 2.3). In the Giuv area zeolites are

abundant in surface outcrops but also in fissures opened during major tunnel

construction (e.g. Gotthard road tunnel, Stalder et al., 1980; new base rail tunnel

Gotthard NEAT). In the tunnels zeolite fissures occur up to 1500 to 2000 m below the

surface. Fissures in other parts of the Central Aar granite do not contain zeolites (Fig.

2.2). The Aar granite is rather uniform and homogeneous in mineralogical and

chemical composition (Labhart, 1977).

Fissure zeolites are not restricted to a specific lithological unit (Figs 2.2 & 2.3).

In the Giuv area, for example, zeolite-bearing fissures occur in most lithologies of the

crystalline basement (Fig. 2.3). Zeolite occurrence is not controlled by changes in

bulk composition of basement rocks. However, zeolites disappear south of the Aar

Massif basement abruptly and are not present in the meta-sediments of the Tavetsch

Massif (Fig. 2.3). In the Gotthard massif zeolites also occur in basement granites and

gneisses but are absent in fissures in the meta-sediments of the cover units (e.g.

Stalder et al., 1980).

The distribution of zeolites in vertical direction is accessible thanks to large

numbers of zeolite samples recovered during tunnel construction operations (e.g.

Gotthard road and rail tunnel, NEAT Gotthard base railway tunnel, NEAT Lötschberg

base rail tunnel). These outstanding data show that laumontite is the dominant zeolite

mineral in Alpine fissures. Scolecite, heulandite, chabazite and stilbite are present in

distinctly smaller numbers of fissure assemblages (Stalder et al., 1980; Amacher,

pers. com.). Laumontite is also abundant at the Arvigo locality (Fig. 2.2), where

various zeolites are found in an active quarry with large blocks of basement gneisses.

However, laumontite has been reported only sporadically, although from many

localities, from surface outcrops where the fissures have been exposed to weathering

conditions (Stalder et al., 1998) (Fig. 2.2). In contrast, scolecite, heulandite, chabazite

and stilbite are the dominant zeolites species in fissures of surface outcrops (e.g.

Niggli et al., 1940; Huber, 1943; Sigrist, 1947; Huber, 1948; Stalder et al., 1998).

The data show that no zonal distribution pattern can be recognized for any of the

zeolite minerals. Both, a zonal regional distribution as well as a vertical zonal

distribution is absent. This lack of mineral zone patterns in the distribution of Alpine

fissure zeolites is in contrast to other environments (Langella et al., 2001; Sheppard &


Hay, 2001; Utada, 2001a, b) where zeolites are formed (e.g. deep marine sediments,

hydrothermal alteration and burial metamorphism in volcanic rocks, saline high

alkaline lakes).

Characteristic of Alpine zeolite fissures is that in many of the fissures a

succession of zeolites minerals can be found as a paragenesis in a single fissure.

Table 2.1. Zeolite localities in the Central Swiss Alps (numbers corresponds to numbers shown in Figs 2.2 & 2.3). No Locality Ca Sco Lmt Stb Heu Cha AFMb Fc Referencesd 1 Bäregg, Oberaar,

Grimsel BE x x x Qtz, Kfs,

Hem, Cal, Chl, Py

x Stalder, 1964; Stalder et al., 1998; SNHMB; °

2 Lötschberg tunnel BE x Qtz, Cal x SNHMB 3 Bächistock GL x x x Niggli et al., 1940; SNHMB 4 Arvigo/Val

CalancaEpi GR x x x x Qtz, Kfs,

Ttn, Act, Ep, Prh, Chl, Ap

xxx Stalder et al., 1998; Wagner, 2000a, b; SNHMB; °

5 BergellPhil * GR x x x x x Qtz, Kfs, Grs, Prh, Ttn, Cal

x Hirschi, 1925; Stalder et al., 1998; SNHMB

6 Cuolm da Vi GR x x Kenngott, 1866; Huber, 1948; SNHMB

7 Davos GR x x x Stalder et al., 1998; SNHMB 8 Drumtobel/Sedrun GR x x x x x Qtz, Kfs,

Cal, Act, Ttn

xxx Stalder et al., 1998; SNHMB; °

9 Misox GR x Chl x SNHMB 10 OberalpstockTho GR Qtz x Niggli et al., 1940; Stalder et

al., 1998 11 Sedrun GR x x x Kenngott, 1866; SNHMB 12 Sella/Gotthard GR x Qtz, Hem,

Chl x Kenngott, 1866; SNHMB; °

13 Stgegia, Medel GR x Qtz x SNHMB 14 Val Casatscha GR x x Cal x Stalder et al., 1998; Huber,

1943; ° 15 Val Cristallina GR x x Cal x SNHMB 16 Val Giuv GR x x x x Qtz, Kfs,

Ap xxx Kenngott, 1866; Sigrist, 1947;

Stalder et al., 1998; SNHMB; °

17 Val Maighels* GR x Chl x Huber, 1943; Stalder et al., 1998; SNHMB

18 Val Medel GR x x Huber, 1943; SNHMB 19 Val Mila GR x Qtz x Niggli et al., 1940; Sigrist,

1947; SNHMB 20 Val Muretto/Bergell GR x x Stalder et al., 1998 21 Val Nalps GR x x Prh, Tur x SNHMB 22 Val Punteglias GR x Qtz, Chl x Huber, 1943; Sigrist, 1947;

Stalder et al., 1998; SNHMB 23 Val Russein GR x x Qtz, Ep, Chl xx Huber, 1943; Stalder et al.,

1998; SNHMB 24 Val Strem GR x x x x Qtz, Chl xxx Kenngott, 1866; Niggli et al.,

1940; Huber, 1948; Stalder et al., 1998; SNHMB; °


continue Table 2.1 25 BiascaEpi TI x Kfs, Ttn,

Ep. Chl x Stalder et al., 1998; SNHMB

26 Camperio/Passo del Lucomango

TI x Qtz, Hem, Cal, Chl,

Prh

x Wagner et al., 1972; SNHMB

27 Domodossola TI x Chl x SNHMB 28 La Fibbia, Gotthard TI x Qtz Chl,

Hem xx Kenngott, 1866; SNHMB

29 Lodrino TI x Qtz, Chl x Stalder et al., 1998; SNHMB 30 Mt Ceneri TI x x Py x Toroni, 1984; Stalder et al.,

1998 31 Pizzo Lucendro TI x x x Qtz Kfs,

Hem, Ms x Stalder et al., 1998; SNHMB

32 Val Baveno TI x x x Stalder et al., 1998 33 Val Canaria* TI x x SNHMB 34 Val Maggia* TI x x x Kenngott, 1866; Stalder et al.,

1998; SNHMB 35 Val Vergeletto, road

tunnel TI x x x Simonetti, 1971; Stalder et al.,

1998; SNHMB 36 Andermatt UR x x SNHMB 37 Brunnital UR x Cal x SNHMB 38 Chrützlistock UR x x Qtz xxx Kenngott, 1866; Niggli et al.,

1940; Sigrist, 1947; Weibel, 1963; Stalder et al., 1998; SNHMB; °

39 Etzlital UR x x x x xxx Kenngott, 1866; Niggli et al., 1940; Sigrist, 1947; Stalder et al., 1998; SNHMB; °

40 Fedenstock UR x Qtz, Chl x SNHMB 41 Fellilücke UR x x Qtz, Flt xxx Niggli et al., 1940; Sigrist,

1947; SNHMB; ° 42 Fellital UR x x x Qtz, Chl xxx Niggli et al., 1940; Sigrist,

1947; SNHMB; ° 43 Göscheneralp UR x x x Qtz, Chl x Kenngott, 1866; Stalder et al.,

1998; SNHMB 44 Gotthard road

tunnel UR x x x x Qtz, Kfs,

Cal, Ep, Prh, Chl, Py

xxx Stalder et al., 1980, 1998; SNHMB; °

45 Griesserental UR x Kfs x SNHMB 46 Maderanertal UR x Qtz, Chl, Py x SNHMB 47 NEAT, Amsteg -

Sedrun UR x x x x Qtz, Kfs,

Cal, Hem, Act, Ep,

Chl, Apo, Py, Anh

xxx °

48 Piz Giuv UR x xxx Kenngott, 1866; Niggli et al., 1940; Sigrist, 1947; Huber, 1948; SNHMB; °

49 Riental UR x x Qtz xxx Kenngott, 1866; Niggli et al., 1940; SNHMB; °

50 Schattig Wichel & Piz Giuv

UR x x x x x Qtz, Kfs, Ap, Ttn, Act, Chl, Ep, Prh

xxx Kenngott, 1866; Niggli et al., 1940; Sigrist, 1947; Huber, 1948; Stalder et al., 1998; SNHMB; °

51 Schijenstock UR x x Qtz, Flt, Hem

xx Niggli et al., 1940; Stalder et al., 1998; °

52 Tiefengletscher UR x Qtz, Gn x SNHMB


continue Table 2.1 53 Arolla VS x x Qtz x Stalder et al., 1998; SNHMB 54 Binntal Meso, Nat, Phil * VS x x x x x Qtz, Kfs,

Hem, Chl, Ttn,

x Kenngott, 1866; Keusen & Bürki, 1969; Stalder et al., 1998; SNHMB

55 Fieschergletscher VS x x Qtz, Kfs, Ep, Ap

x Kenngott, 1866; Niggli et al., 1940; Stalder et al., 1998; SNHMB

56 Furka tunnel VS x Qtz, Kfs, Chl, Ttn

x Wälti, 1984; SNHMB

57 Gibelsbach/FieschEpi VS x x x x Qtz, Kfs, Flt xxx Kenngott, 1866; Koenigsberger, 1917; Armbruster et al., 1996; Stalder et al., 1998; SNHMB

58 Gornergletscher VS x x Stalder et al., 1998 59 Gredetschtal/Brig VS x Qtz, Kfs,

Chl, Ttn xx Niggli et al., 1940; Stalder et

al., 1998; SNHMB 60 Grosses Sidelhorn VS x Qtz, Chl x SNHMB 61 Lax VS x x x Qtz, Chl x Kenngott, 1866; Stalder et al.,

1998; SNHMB 62 Lötschental* VS x x x x Qtz, Kfs,

Cal, Act xx Fellenberg von, 1893; Stalder

et al., 1998; SNHMB 63 Martiny VS x Qtz x SNHMB 64 Massaschlucht VS x Kfs, Cal x SNHMB 65 Mättital VS x x Qtz, Cal x SNHMB 66 Nuffenenpass VS x Qtz x SNHMB 67 Pollux/ZermattNat VS xx Stalder et al., 1998 68 Simplontunnel VS x Chl x Stalder et al., 1998; SNHMB

Mineral abbreviations used after Bucher & Frey (2002). Following abbreviations are used for zeolites: Sco = scolecite, Lmt = laumontite, Stb = stilbite, Heu = heulandite, Cha = chabazite. a Canton (Swiss districts, BE = Bern, GL = Glarus, GR = Grisons TI = Ticino, UR = Uri, VS = Valais). b major associated fissure minerals. c frequency of zeolites (x = sporadic zeolite occurrence, xx = cumulative zeolites occurrence, xxx = zeolites occur frequently). d SNHMB = collection of the Swiss Natural History Museum Bern.. Epi epistilbite. Meso mesolite. Nat natrolite. Phil phillipsite. Tho thomsonite. * = zeolites hosted in basic rocks (e.g. amphibolite). ° = localities from which samples were analyzed in this study

2.5.2. Field occurrences

2.5.2.1. General features

The zeolites reported in this paper occur exclusively in fissures, veins and gashes.

Rock forming zeolite minerals such as in the Taveyanne greywacke (Rahn et al.,

1994) are not considered here. The fissure zeolites cover and coat the walls of

fractures in granites and gneisses but also occur as pore and cavity filling in leached

host rocks. Zeolites typically overgrow earlier formed minerals of the fissure

assemblage, but they also occur as single stage fissure deposits in granites and

gneisses (e.g. Kenngott, 1866; Parker, 1922; Stalder et al., 1980; Armbruster et al.,

1996).


Assemblage and distribution data of zeolite minerals are compiled in Table 2.1.

Zeolites are present as transparent, white or brownish crystals. The crystal size is

relatively small and usually does not exceed 1 cm. This is in contrast to other fissure

minerals, which commonly reach crystal sizes of several cm to dm. Typically zeolites

occur as very small (< 1 mm) inconspicuous whitish coatings on earlier formed

minerals or direct on the fracture wall, which makes them difficult to recognize. Some

zeolite crystals have a green color (Table 2.2; A8332), due to small inclusions of

chlorite.

Fig. 2.5: Fissures in the Gotthard NEAT tunnel. (a) Tunnel head wall showing fissure, that strikes in S-W direction. (b) Sigmoidal fissure. (c) Photomicrograph of laumontite needles. (d) Laumontite covering earlier formed quartz as dense mats.

The fissures tend to be lens-shaped with large long- to short-axis ratios. Open fissure

cavities range from cm to several tens of meters in length. The aperture of the

fractures varies from mm to tens of cm (some open crystal caves of > 1m have been

found). The fractured host rock is commonly chemically leached on both sides of the

fissure (Fig. 2.4). The leached zone is usually lighter colored than the unaltered host

rock, due to the lack of primary dark minerals such as biotite in the former (Fig. 2.4).

The leached zone shows a higher porosity compared with the host rock (e.g. Parker,

1922; Sigrist, 1947; Mercolli et al., 1984; Ciesielezuk & Janeczek, 2004). The

secondary porosity of the leached rock is locally filled with secondary minerals,

forming a zone of impregnation (e.g. Huber, 1943; Mercolli et al., 1984). Leaching

zones are not always present in the fissures. Leaching zones range from mm to m

scale, but it seems that the width of leaching zone is related to the aperture of the

fracture. Leaching always increases towards the central open space. Frequently much

of the primary rock material has been removed and the remaining alteration products

form a very porous crumbly disjointed mass (Table 2.2; DT, TW34).


The dominant minerals in Alpine fissures are quartz, adularia and chlorite. There

is a remarkable absence of clay minerals (except chlorite) in all Alpine fissures. In the

most common multi-mineral veins and fissures the volume fraction of zeolites is very

low (Huber, 1943; Sigrist, 1947; Stalder et al., 1980). However, single-mineral zeolite

fissures are common and widespread. In tunnel fissures where laumontite is the

dominant zeolite covering fissure walls as dense mats (Figs 2.5 & 2.6), the zeolite

may be modally abundant also in multi-mineral veins (Stalder et al., 1980). Because

zeolites tend to be well preserved in fissures opened in the progress of tunnel

constructions, tunnel data give a clue on the frequency and the abundance of zeolites

in fractured granite and gneiss. During construction of the 16 km long Gotthard road

tunnel 225 fissures were recorded during a systematic evaluation of fissures in the

main- and security tunnel in the northern section (from north portal to 7 km to the

south). 41 of the fissures contained zeolites (18 % of all fissures; Stalder et al., 1980).

In some lithologies, for instance in the Southern gneisses of the Aar Massif, zeolites

were found in 50 % of the fissures (Stalder et al., 1980). Similar data are available

from the new Gotthard NEAT tunnel. In the section 12000 to 18555 meters (Fig. 2.7),

26 of 83 fissures (32 %) yield zeolite species (P. Amacher pers. com.).

Zeolites normally occur together with other mineral species in a vein. A general

chronology of zeolite-bearing Alpine fissures is compiled in Table 2.3. The schematic

sequence of successive minerals and mineral assemblages in Alpine fissures is based

on observed textural (overgrowth) relationships providing relative age and sequence

from a large number of fissures (Table 2.2).

Zeolites generally overgrow all earlier formed minerals and usually represent the

latest mineral formed in Alpine fissure. In some veins euhedral crystals of apophyllite

overgrow locally laumontite as (e.g. Gotthard NEAT, Table 2.2; KB868; Arvigo,

Wagner et al., 2000a, b; Gotthard road tunnel, Stalder et al., 1980). All six Ca-zeolites

(Table 2.3) never occur together, a maximum of three different successive zeolites

may be found in a single fissure. Single zeolite veins are common but many veins

contain at least two different zeolites. The structures of multi-zeolite veins show that

zeolites never co-precipitate but rather are always diachronous.

Laumontite is the most common zeolite. It occurs in mono-zeolite veins and in

fissures with stilbite or scolecite or both. Samples from the Gotthard NEAT tunnel

(Table 2.2; TW01.2, TW01.3, TW02) consistently show that early laumontite is


overgrown by late stilbite. This observation is confirmed by data from the Gotthard

road tunnel (Stalder et al., 1980). Samples from Arvigo (Table 2.2; A8) demonstrate

that early scolecite is overgrown by laumontite (Fig. 2.6). However, one sample

suggests an inverse growth relationship in which laumontite formed before scolecite

(Table 2.2; Arvigo1).

Table 2.2. Zeolite assemblages in Alpine fissures analyzed in this study No- Mineral assemblages Sourcea, b, c Nod Description

A3* Stb Fr 24 light brownish Stb crystals, Val Strem A4* Qtz-Kfs-Stb P.A 47 Stb covers Qtz and Kfs as dense mats, up to 3 mm long white euhedral

needles, Gotthard NEAT A4190 Hm-Cc-Lmt SNHMB 26 Lmt on Cc scalenoeder, which overgrows Hm, Camperio A5203* Qtz-Heu SNHMB 1 euhedral Heu crystals up to 1cm in size growing on top of Qtz crystals

in a fissure of altered sericite-gneiss, Oberaar/Grimsel A6* Qtz-Stb P.A. 47 Stb forming flat-topped crystals and fan-like crystal aggregates, grown

on Qtz, Gotthard NEAT A8* Kfs-Prh-Ep-Chl-Sco-Lmt Fr 4 Sco needles and Lmt on top of Kfs, Ep, Prh and Chl in a fissure hosted

in an orthogneiss, Arvigo/Val Calanca A8.1* Qtz-Kfs-Ep-Chl-Sco-Heu P.A. 4 Sco needles and Heu on top of Kfs, Ep and Chl in a fissure hosted in

an orthogneiss, Arvigo/Val Calanca A8.2* Heu-Stb Fr 24 coffin shaped Heu crystals with a blocky habit, followed by Stb, as

dense mats of crystals, up to 4 mm in size, Val Strem A8332* Kfs-Ap-Chl-Cha SNHMB 55 euhedral Cha crystals up to 5 mm, associated with Kfs and Ap; Cha

appear to be green, because of Chl inclusions, Fieschergletscher Arvigo1* Ep-Prh-Sco-Lmt Fr 4 Sco tufts and Lmt cover Ep and Prh in a fissure hosted in an

orthogneiss Arvigo/Val Calanca Arvigo12I*

Lmt Fr 4 anhedral Lmt crystals, fills up a highly porous zone in a leaching zone hosted in an orthogneiss, Arvigo/Val Calanca

Arvigo13*

Sco Fr 4 Sco needles up to 4 cm in length located in fissure, Arvigo/Val Calanca

Arvigo2* Lmt Fr 4 Lmt, which totally fill up a 4 cm wide and 10 cm long boudinage gash; crystal sizes up to 2 cm. Arvigo/Val Calanca

B12* Sco-Heu-Stb Fr 24 euhedral crystals of Sco, Heu and Stb, Schattig Wichel, B981* Chl-Lmt SNHMB 44 green Lmt crystals up to 1 cm in length, green colour of Lmt appear

because of Chl inclusions, Gotthard road tunnel B3140 Qtz-Kfs-Cc-Chl-Cc-Sco SNHMB 4 fissure mineralization in chloritized gneiss, Arvigo, Chaba1* Qtz-Cha Fr 51 rhombohedral, transparent Cha crystals, up to 2 mm in size, grown

after Qtz in altered granite, Schijenstock Chaba2* Cha Fr 28 rhombohedral Cha crystals in a fissure of granite Stella/Gotthard DT* Heu-Stb Fr 11 light brownish Stb crystals, forming 5 cm long fanlike bow ties, which

associated with early formed Heu crystals on top of a highly porous matrix, Drumtobel/Sedrun

Fi1* Qtz-Heu-Stb Fr 57 euhedral Stb crystals, up to 1 cm in size associated with Heu and Qtz, Gibelsbach/Fiesch

Fi2* Flt-Stb Fr 57 Stb on green fluorite, Gibelsbach KB868* Qtz-Kfs-Lmt-Apo P.A. 47 Apophyllite-(F) overgrows Lmt, Kfs and Qtz in a fissure of Gotthard

NEAT K614* Qtz-Cha-Stb Fr 50 smoky Qtz, covered by Cha rhomboeder and fan-shape Stb aggregates,

Val Val NL4b* Lmt P.S. 47 Lmt which fills up a 0.2 mm wide vein in a highly porous and altered

gneiss, Gotthard NEAT R1* Qtz-Stb Fr 49 light brounish Stb, forming radial groups of 1 cm in diameter, hosted

in quartz lenses in paragneiesses, Riental TW01* Lmt P.A. 47 Lmt on Qtz, Gotthard NEAT


continue Table 2.2 TW01.2* Qtz-Chl-Lmt-Stb P.A. 47 Stb associated with Lmt and Chl as early phases in a fissure of the

Gotthard NEAT TW01.3* Lmt-Stb P.A. 47 Lmt associated with Stb as early phase in a fissure of the Gotthard

NEAT TW02* Qtz-Lmt-Stb P.A. 47 Lmt on Qtz associated with Stb as early phase in a fissure of the

Gotthard NEAT TW03* Cc-Lmt P.A. 47 Lmt associated with early formed Cc in a fissure of Gotthard NEAT TW11.1* Qtz-Chl-Lmt P.A. 47 Lmt associated with early formed Chl and Qtz in a fissure of Gotthard

NEAT TW20* Sco-Heu-Stb Fr 50 euhedral crystals of Sco, Heu and Stb, Schattig Wichel TW34* Heu-Stb Fr 24 light brownish Stb crystals, forming 5 mm long bow ties, which

associated with early formed Heu crystals, on a highly porous matrix, Val Strem

4172 Cc-Stb SNHMB 37 Stb grow on Cc, Brunnital 7844 Qtz-Chl-Cc-Sco SNHMB 50 Sco and Cc growing on Chl and Qtz, Schattig Wichel 7890 Cc-Heu SNHMB 46 Heu on Cc scalenoeder, Maderaneetal 30992 Qtz-Cc-Lmt SNHMB 2 Lötschberg exploring tunnel 35370* Stb SNHMB 44 light brownish Stb crystal, up to 3 mm, grown on top of fine Qtz

crystals, fissure breccia 2830 meter after north portal, Gotthard road tunnel

35508* Kfs-Chl-Stb SNHMB 44 light pinkish Stb, southern granite gneisses, (3240m) Gotthard road tunnel

35795* Lmt SNHMB 44 Lmt from the Gotthard road tunnel 35843* Sco-Stb SNHMB 44 Stb on top of Sco from a fissure of the Gotthard road tunnel 36728* Qtz-Cc-Ms-Ttn-Stb SNHMB 44 Stb crystals up to 5 mm growing on Cc, Fibbia granite gneisses,

Gotthard road tunnel s SNHMB = collection of the Natural History Museum Bern. b P.A. = sample provided by Peter Amacher. c Fr = sample collected during field work or owned by the Mineralogical Museum University Freiburg. d number corresponds to Table 2.1 and Figs 2.2 & 2.3. * = sample analysed during this study

Frequently heulandite succeeds scolecite (Table 2.2; A8.1, TW20; Fig. 2.6). Single

heulandite crystals are penetrated by older scolecite tufts (Fig. 2.6e). In some veins

additional stilbite forms the latest zeolite in this succession (Sco → Heu → Stb).

Heulandite occurs either in the assemblage scolecite-heulandite or heulandite-stilbite

or scolecite-heulandite-stilbite, whereas the assemblage laumontite-heulandite has not

been found. Heulandite in mono-zeolite veins is rare and occurs occasionally grown

on calcite as substrate mineral (Table 2.2; 7890).

Stilbite appears with all other zeolites except epistilbite, but it also appears

frequently in mono-zeolite fissures (Table 2.2; A4, A6, R1, 35370). In samples from

the NEAT tunnel stilbite associated with earlier formed laumontite. Overgrowth of

stilbite on heulandite can be found at all zeolite localities (Table 2.2; DT, Fi1, TW34).

Stilbite is always the last zeolite in the succession (Table 2.2). Chabazite occurs

widespread but in small amounts and irregularly. It is subordinate to laumontite,

stilbite, scolecite and heulandite (Fig. 2.2). It is associated with and overgrown by


later formed stilbite (Fig. 2.6; Table 2.2; K614). At some localities it also occurs in

single zeolite veins (Table 2.2; A8332, Chaba1, Chaba2). Chabazite has never been

observed together with heulandite, laumontite and scolecite. The rarest zeolite is

epistilbite, which is only known from 3 localities (Table 2.1; no 4 25, 57). It is not

associated with other zeolite species.

From these observations the following growth chronology can be deduced (from

old to young): Sco → Lmt → Heu → Cha → Stb. Note that wairakite, a common Ca-

zeolite elsewhere, has not been found in the Central Alps.

Most zeolites occur associated with earlier formed minerals in the veins. These

minerals include silicates, oxides, sulfides, sulfates, carbonates, phosphates and

halides (Table 2.1). Zeolites always overgrow all these minerals (except apophyllite).

Although a large number of different minerals occur together with zeolites, in a single

fissure not more than six different minerals are normally present. A full list of

minerals occurring with zeolites in fissures from different localities is given in Table

2.1 and the data from the analysis of mineral assemblages in this study are given in

Table 2.2. It is evident from the data compilation that zeolites do not occur

preferentially in veins with a preferred mineral assemblage. The data presented in

Table 2.2 demonstrate that zeolites occur unrelated to the type of minerals deposited

in the vein prior to zeolite formation.

Table 2.3. Crystallization chronology of zeolite bearing Alpine fissures determined in this study. quartz adularia

± titanite ± actinolite ± ilmenite

± anhydrite

epidote calcite fluorite

± hematite

prehnite chlorite calcite

zeolites - scolecite - laumontite - heulandite - chabazite - stilbite

apophyllite

early

late

Zeolites grow on a substrate, which is either the host rock surface or an earlier formed

mineral. Preferential substrate minerals are quartz, adularia and calcite. Locally

special and unique assemblages in fissures associated with zeolites occur. For

instance, the assemblage quartz - adularia - green fluorite - zeolite occurs in a fissure

breccia at Gibelsbach near Fiesch (Figs 2.2 & 2.6d; Table 2.1; no 57). A fissure at

Tiefengletscher (Table 2.1; no 55) exhibits a unique paragenetic relation of laumontite

overgrowth on galena and smoky quartz. Often titanium minerals, like titanite,


brookite, rutile and ilmenite appear in fissures as accessory phase in minor amounts.

Remarkable is the sporadic presence of sulfide and sulfate minerals in veins from the

same rock type. For instance in the Central Aar granite at tunnel level (Gotthard

NEAT) the assemblages anhydrite-chlorite-laumontite-stilbite and quartz-chlorite-

calcite-pyrite-laumontite are present (P. Amacher, pers. com.). Museum quality

hematite roses from La Fibbia and Pizzo Lucendro (Table 2.1; no 28, 31), are

associated with stilbite in the complete paragenesis of quartz-adularia-muscovite-

hematite-rutile-stilbite (Stalder et al., 1998). Similar assemblages are known from the

same Fibbia gneiss, approximately 1500 m beneath La Fibbia at the level of the

Gotthard road tunnel (Stalder et al., 1980, 1998).

The most common mineral that occurs together with zeolites is quartz (Figs 2.5 &

2.6; Table 2.1 & 2.2). If quartz and zeolites are the only minerals in a fissure, the

zeolite grows directly on the crystal surface of euhedral quartz (Table 2.2; A5203, A6,

A8.1, Fi1). Substrate quartz does not show any sign of leaching and dissolution.

Alpine fissure quartz crystals overgrow and include a wide range of different minerals

such as chlorite, rutile, anhydrite and many others (Stalder et al., 1998). However,

zeolite inclusions in quartz are unknown. This implies that quartz growth always

precedes the onset of zeolite growth. Additional phases in fractures of gneiss and

granite are often adularia and chlorite, which are together with quartz the most

common minerals in Alpine fissures (Stalder et al., 1998). Sample KB868 (Table 2.2)

provides a full growth chronology of these phases with the sequence quartz →

adularia → laumontite → apophyllite. Similar to quartz-substrate, the contact surface

between adularia and laumontite is planar and even without leaching or dissolution

textures (Table 2.2; A8332, A8). Chlorite occurs in large quantities in Alpine fissures

and in association with zeolites and marks the most frequent Mg-Fe silicate in Alpine

fissures (Stalder et al., 1998). Chlorite very often occurs as unconsolidated chlorite

sand in fissures forming vermicular grains, with grain sizes less than 2 mm. Because

of the loose nature, chlorite is reworked and often fills up open spaces in the porous

rock matrix and in open spaces between older minerals as well as between younger

minerals. Inclusions of chlorite in later formed species including zeolites give them a

green color typical of many localities (Table 2.2; A8332). The observation suggests

that chlorite formed prior to zeolites. In addition to the frequent substrate minerals

quartz and adularia, also calcite often serves as substrate for zeolite growth (Table 2.1


& 2.2; TW03, 7844, A4190, 30992, 7890, 4172). Calcite may occur in different

generations during vein mineralization (Table 2.3). Calcite formed before and after

chlorite growth and it occurs with all zeolites except chabazite and epistilbite.

Fig. 2.6: Representative zeolite assemblages from the Central Swiss Alps. (a) Radial stilbite aggregates overgrows chabazite rhombohedra on quartz; Val Val/GR. (b) Fan-shape stilbite overgrows coffin shaped heulandite crystals; Val Strem/GR. (c) Laumontite fibrous; Gotthard NEAT/UR. (d) Stilbite associated with green fluorite; Gibelsbach/VS. (e) Coffin shaped heulandite grows on scolecite needles, which grow on quartz covering highly porous rock matrix; Schattig Wichel/UR. (f) Scolecite grows on Alpine fissure quartz; Gotthard NEAT tunnel/UR. (g) Scolecite on epidote, adularia and quartz; Arvigo/GR. (h) Laumontite and scolecite associated with calcite that shows dissolution Arvigo/GR.


Calcite often occurs as scalenoeders on which zeolites developed (Table 2.2; 30992,

7890). It also occurs as paper spar that often shows evidence of dissolution and

resorption (Table 2.2; TW03; Fig. 2.6).

Prehnite and epidote are Ca-Al silicate minerals that formed prior to the zeolites

at some localities (Table 2.1 & 2.2). Actinolite that often forms asbestos fibers also

occurs locally. Pumpellyite, a diagnostic mineral of the prehnite-pumpellyite and

pumpellyite-actinolite metamorphic facies, has never been found in Alpine gneiss and

granite fractures. However prehnite and epidote do occur in granite and gneisses

(Table 2.1 & 2.2), but zeolites are often lacking (Stalder et al., 1998) in the

succession. The Arvigo quarry, fissures around Piz Giuv and the Gotthard NEAT

tunnel (Table 2.1 & 2.2) represent localities were prehnite and epidote are frequently

found associated with zeolites. The mineral sequence deduced from these localities is:

quartz-adularia-actinolite-epidote-prehnite-zeolite (Table 2; A8).

2.5.2.2. Aar Massif/Gotthard NEAT tunnel

During the ongoing excavation of the Gotthard NEAT tunnel (NEw Alp Transit, 57

km long rail base tunnel through the Central Alps, Figs 2.2 & 2.3) a large number of

mineralized fissures and veins were opened, many of which contained a museum

quality mineral specimens (Figs 2.5 & 2.6). Of particular interest for the present

research was tunnel section under Chrüzlistock (Figs 2.3 & 2.7) and the surface

outcrops above to the tunnel section, where zeolites have been frequently found in

fissures (e.g. Kenngott, 1866; Parker, 1922; Niggli et al., 1940; Huber, 1943; Sigrist,

1947; Stalder et al., 1998). In other sections of the NEAT tunnel no zeolites have

been found so far or the tunnel drilling is still under progress.

Figure 2.7 gives an overview of the main fissure mineral assemblages in relation

to the lithologies in the tunnel section, where zeolites have been found. Zeolite

occurrences in surface outcrops vertically above the tunnel correspond to those in the

tunnel section. In surface outcrops zeolites are concentrated around the Giuv syenite

(Fig. 2.3) a W-E trending igneous body of syenitic composition (Labhart, 1977). But

the syenite-unit pinches out with depth and is not present on tunnel level (500 m)

(Fig. 2.7). In total more than 25 different mineral species have been found in veins

and fractures in the studied section of the NEAT tunnel (7950-8850 meter in distance


to the north portal). Adularia, albite, amianthus, calcite, chlorite, pyrite and quartz

occur evenly distributed over the whole section. Anatase, anhydrite, ankerite, apatite,

apophyllite, chalcopyrite, epidote, fluorite, galena, graphite, hematite, milarite,

spahlerite, synchisite and titanite occur only sporadically or in traces (P. Amacher

pers. com.). In surface outcrops the following additional minerals were found

associated with zeolites (Huber, 1948): datholite, limonite, molybdenite and prehnite.

To be mentioned is the well-known “Skolezitkehle” (Huber, 1948), located at Schattig

Wichel, the north wall of Piz Giuv, where a numerous zeolite-bearing fissures were

found.

Fig. 2.7: Main fissure mineral assemblages in the Gotthard NEAT tunnel. Profile shows the drilling section Amsteg-Sedrun, from tunnel meter 12000 to 18550 (distance from north portal at Erstfeld), where zeolite minerals were recovered. In the section between 7950 and 12000 meter no zeolites were found in fissures. Drilling in the three other tunnel sections is still in progress. Tunnel level is 500 meter above sea level and overburden ranges between 1290 and 2130 meter. The position of Chrüzlistock (Fig. 2.3) is marked with a black star; (data derived from P. Amacher, pers. com.).

In the tunnel section laumontite is the most abundant zeolite mineral, which is in

sharp contrast to observations from surface outcrops (Table 2.1), where laumontite

occurs only sporadically in small amounts (Huber, 1943; Sigrist, 1947; Stalder et al.,

1998). Stilbite (stellerite) follows as the second most common zeolite in the tunnel,

whereas scolecite, heulandite and chabazite occur only sporadically in the tunnel,


again in contrast to observations from surface outcrops where stilbite, scolecite and

heulandite occur in similar proportions and chabazite is relatively rare (e.g. Huber,

1948).

Fissures that are filled with zeolites only indicate a late formation of the fracture.

A systematic study of fissure chronology by Heijboer (2006) focused on quartz

formation. It showed that zeolite species formed during the last mineral precipitation

phase, during which the orientations of fissure and veins are generally SE-NW or NE-

SW, respectively (Huber, 1946; Heijboer, 2006).

2.5.2.3. Gotthard massif/ Gotthard road tunnel

Excavation of the 16 km long Gotthard road tunnel in the seventies of the last century

(Fig. 2.2) supplied considerable amounts of fissure minerals including zeolites from

the Gotthard massif (Stalder et al., 1980). This amount of material is in distinct

contrast to the small number of finds of zeolite veins from surface outcrops in the

Gotthard massif (Fig. 2.2; Table 2.1). Thus zeolites are very frequent in both the

Gotthard road tunnel and in the Gotthard NEAT tunnel.

The spatial distribution of zeolites in the Gotthard road tunnel can be related to

the type of host rock of the fissure. The host rock controls the proportion of fissures

filled with zeolites and the nature of the dominant zeolite present (Fig. 2.8). A

detailed study of minerals in fissures of the Gotthard road tunnel during excavation by

Stalder et al. (1980), has shown, that zeolites occur in different modal proportions in

different lithologies but that they are not present in all rock types. Figure 2.8 gives an

overview of fissure bearing lithologies and the relative amount of fissure that contain

zeolites. The units along the tunnel sections are (from N to S): Central Aar granite

(ZArg), Southern gneisses (SGn), Permian-Carboniferous (PC), Gurschen gneiss

(GGn), Fibbia granite gneiss (FGgn) and Tremola Series (TrS). Where CO2 fluids

were found in fluid inclusions of quartz, e.g. in GGn fissures are zeolites devoid and

in the PC unit only a small number of fissures (<10 % of all fissures) contain

laumontite (Stalder et al., 1980). Laumontite with minor amounts of stilbite,

heulandite scolecite and chabazite are typically found in the ZArg (17 % of fissure

hold laumontite), SGn (45 %) and TrS (25 %) units. A distinctly different zeolite

population pattern is observed in the FGgn rock unit. Stilbite (>50 %) represents the


major zeolite in the Fibbia gneisses, whereas laumontite, scolecite and chabazite were

only found in less than 10 % of all fissures (Stalder et al., 1980). Similar to the

Gotthard NEAT tunnel apophyllite is associated with laumontite and marks a younger

fissure generation.

Fig. 2.8: Simplified geological cross section through the Gotthard road tunnel and incidence of zeolites (% of fissures with zeolites, n = number of fissures) in fissures for different lithological units (Stalder et al., 1980). (ZAgr: Central Aar granite, SGn: Southern gneisses, Me: Mesozoic units, PC: Permian-Carboniferous, GGn: Gurschen gneiss, GGr: Gamsboden granite gneisses, GGn: Guspis gneiss, FGgn: Fibbia granite gneiss, SGn: Sorescia gneiss, TrS: Tremola Series.

2.5.2.4. Gibelsbach/Fiesch

The mineral fissures from Gibelsbach (Fig. 2.2) have been described already by

Kenngott (1866). The mineral assemblage of green octahedral fluorite, quartz,

adularia, albite and various zeolite species (Fig. 2.6d) is unique to this locality. In

addition to the five major zeolites (scolecite, laumontite, heulandite, stilbite/stellerite,

and chabazite) Gibelsbach is one of three localities (Table 2.1; no 4, 25), where

epistilbite is found. Single crystal studies have shown that the mineral previously

identified as stilbite is in fact stellerite, the orthorhombic Ca-endmember equivalent

of monoclinic stilbite (Armbruster et al., 1996).

The Gibelsbach mineral veins are hosted in a brecciated, highly porous and

strongly foliated granite (Armbruster et al., 1996). The zeolite-bearing zone is

bordered to the south by Permian sedimentary rocks occurring between Aar- and


Gotthard Massif and to the north by a coarse grained granite of the southern part of

the Aar Massif (Zbinden, 1949) (Fig. 2.2). Important at this locality is that in contrast

to most other zeolite localities chlorite or other iron-magnesium silicates have not

been observed in the outcrop.

2.5.2.5. Arvigo/Val Calanca

Banded biotite gneiss and coarse-grained light colored two-mica gneiss are mined as

building stones in Arvigo (Fig. 2.4a). The rocks belong to the upper Simano nappe of

the crystalline Penninic basement. The Arvigo quarry became famous for a large

number of Alpine fissure minerals, which occur in extensional fractures and cavities

of the granitic gneisses. Arvigo is the prime zeolite locality within the Penninic

nappes (Table 2.1; no 25, 30, 32, 35). The Arvigo fissures contain the assemblage

quartz, adularia, epidote, prehnite, chlorite and all zeolite species known from Alpine

fissures. Apophyllite is present as the latest mineral. In general, pale green sheaves of

epidote are overgrown by prehnite, chlorite and zeolites marking the most common

mineral assemblages in Arvigo (Fig. 2.6g). Scolecite and laumontite are the dominant

zeolite species, whereas heulandite, chabazite, stilbite and epistilbite can only found

sporadically. The occurrences of mesolite (Weiß & Forster, 1997) as overgrowth on

scolecite could not be confirmed during this study. Comparing to other surface

outcrops, the Arvigo quarry shows the highest abundance of laumontite in the Alps.

This is related to the active quarrying, which reveals a steady supply of fresh

unaltered material. More than 40 different minerals species are found in fissures of

the Arvigo quarry, most of the minerals are rare (full list of minerals see Wagner et

al., 2000a and b). Remarkable is the appearance of babingtonite, which usually is

associated with zeolites in basic igneous rocks (Armbruster, 2000; Armbruster et al.,

2000). The mineral assemblages vary with host rock composition and mineralogy, but

zeolite species are always present as latest mineral species in the fractures. Fluid-rock

interaction with the host rock, forms noticeable leaching zones (Fig. 2.4). These 6-7

cm wide leaching zones are usually lighter colored than the dark country rock, due to

the removal of mafic minerals. The high abundance of zeolites in Arvigo is

remarkable and unique for the crystalline Penninic nappes. There are a large number


of quarries in the area quarrying similar rocks from the same or the neighboring

nappes but not in any of them comparable zeolites finds have been made.

2.5.3. Mineralogy and crystal chemistry of zeolites and associated

minerals

2.5.3.1. Chabazite-Ca

Chabazite ((Ca0.5,Na,K)4(Al4Si8O24) •12 H2O, Passaglia & Sheppard, 2001) forms

granular pseudorhombohedral, 1 to 4 mm large, transparent to translucent colorless or

white crystals (average 2 mm; Fig. 2.6a). The triclinic crystals (Armbruster & Gunter,

2001) often form penetrating twins with corners projecting from the faces. It grows on

fissure quartz crystals (rock crystals) and is often associated and overgrown by

stilbite.

Chabazite in Alpine fissures can be classified as chabazite-Ca (Coombs et al.,

1998), because the main extra-framework cation is Ca. Ca occupies on an average 60

mole% of the extra-framework cations (range from 56 to 66 mole%). Representative

analyses are given in Table 2.4. K and Sr are present in high proportions in the

channels structure of chabazite (Fig. 2.9), the chabazite-K content ranges from 24 to

35 mole%, with an average value of 30 mole% of all extra-framework cations. Sr

occupies up to 10 mole% of the extra-framework sites, making Sr to an important

zeolite cation similar to heulandite. Other extra-framework cations such as Na, Mg,

Ba and Fe occur in traces only and can therefore be neglected. Fissure chabazite

Si/(Si+Al) ratio of 0.70 to 0.71 (Fig. 2.10), which is higher than the mean value of

amygdaloidal chabazite crystals (0.67; Passaglia & Sheppard, 2001) that can be

related to the coupled substitution K+ + Si4+ = Ca2+ + Al3+. Zoned chabazite shows

that the K content increases and the Ca decreases from core to the rim.


Table 2.4. Representative zeolite analysis Sample no. A8332 Chaba1 Chaba2 Chaba2 A5203 TW20 TW20 A8.1 Fi1 TW34 Analysis no. 2 2 3 4 1 6 7 12 4 6 wt.% Cha Cha Cha Cha Heu Heu Heu Heu Heu Heu SiO2 52.76 53.59 54.37 53.80 54.64 60.74 59.10 56.30 56.67 57.76 Al2O3 19.47 19.24 19.51 19.53 16.51 15.99 15.08 16.56 15.60 15.61 CaO 8.35 8.52 8.00 8.48 5.47 6.77 6.36 6.87 6.07 5.99 SrO 1.88 1.58 2.18 2.29 4.35 2.40 2.35 0.92 2.31 3.08 BaO -- 0.08 0.01 0.00 - - - - 0.00 0.28 Na2O 0.17 0.08 0.20 0.15 0.14 0.03 0.03 0.38 0.11 0.07 K2O 3.75 3.72 3.70 2.73 1.53 1.36 1.36 2.47 2.28 1.40 H2O 13.58* 13.14* 12.02* 13.00* 17.30* 12.67* 15.68* 16.47* 16.93* 15.75* Totala 100.00 100.00 100.00 100.00 100.00 100.00 100.00 100.00 100.00 100.00 Si 8.300 8.364 8.385 8.360 26.573 27.446 27.636 26.713 27.139 27.276 Al 3.610 3.539 3.546 3.577 9.463 8.517 8.312 9.259 8.805 8.688 Ca 1.407 1.425 1.322 1.412 2.850 3.276 3.184 3.491 3.115 3.031 Sr 0.171 0.143 0.195 0.206 1.227 0.629 0.638 0.253 0.641 0.843 Ba - 0.005 0.001 0.000 - - - - 0.000 0.052 Na 0.052 0.024 0.060 0.045 0.132 0.027 0.024 0.350 0.102 0.064 K 0.753 0.741 0.728 0.541 0.949 0.786 0.811 1.496 1.393 0.843 O 24 24 24 24 72 72 72 72 72 72 H2O 7.125 6.840 6.183 6.737 28.074 19.088 24.452 26.073 27.041 24.807 E%b -8.89 -9.92 -7.23 -6.55 2.47 -1.50 -2.17 -0.81 -2.40 -1.60 Si/(Si+Al) 0.70 0.70 0.70 0.70 0.74 0.76 0.77 0.74 0.76 0.76 Ca/(Ca+Na+K+Sr) 0.59 0.61 0.57 0.64 0.55 0.69 0.68 0.62 0.59 0.63 Na/(Ca+Na+K+Sr) 0.02 0.01 0.03 0.02 0.03 0.01 0.01 0.06 0.02 0.01 K/(Ca+Na+K+Sr) 0.32 0.32 0.32 0.25 0.18 0.17 0.17 0.27 0.27 0.18 Sr/(Ca+Na+K+Sr) 0.07 0.06 0.08 0.09 0.24 0.13 0.14 0.05 0.12 0.18 Sample no. DT A8.2 A8 TW11.1 TW03 TW01.2 35795 NL4b Arviog12I Arvigo2 Analysis no. 5 3 2 6 1 8 3 9 10 7 wt.% Heu Heu Lmt Lmt Lmt Lmt Lmt Lmt Lmt Lmt SiO2 57.53 57.66 52.23 53.00 53.18 52.33 51.49 51.95 52.06 52.43 Al2O3 15.08 15.96 21.89 20.87 20.64 21.46 21.63 20.64 21.48 21.15 CaO 5.96 5.98 12.20 11.35 11.28 11.97 12.02 11.06 11.65 11.82 SrO 1.95 3.72 0.03 0.12 0.03 0.05 0.24 0.27 0.17 0.00 BaO 0.41 0.50 0.00 - - - - 0.00 0.00 0.00 Na2O 0.26 0.08 0.04 0.12 0.08 0.04 0.08 0.16 0.00 0.09 K2O 1.97 1.51 0.05 0.33 0.33 0.05 0.05 0.46 0.32 0.12 H2O 16.81* 14.532* 15.24 14.17* 14.457* 14.08* 14.49* 15.44* 14.19* 14.37* Totala 100.00 100.00 101.69 100.00 100.00 100.00 100.00 100.00 100.00 100.00 Si 27.442 27.062 16.036 16.362 16.443 16.149 16.013 16.313 16.124 16.228 Al 8.478 8.827 7.923 7.594 7.521 7.806 7.928 7.640 7.841 7.713 Ca 3.046 3.005 4.014 3.754 3.735 3.957 4.005 3.721 3.866 3.919 Sr 0.539 1.013 0.005 0.021 0.005 0.008 0.043 0.049 0.031 0.000 Ba 0.077 0.092 0.000 - - - - 0.000 0.000 0.000 Na 0.240 0.068 0.023 0.072 0.050 0.023 0.048 0.099 0.000 0.053 K 1.199 0.905 0.019 0.130 0.129 0.021 0.020 0.185 0.126 0.047 O 72 72 48 48 48 48 48 48 48 48 H2O 26.743 22.747 13.884 14.591 14.909 14.486 15.029 16.170 14.658 14.837 E%b -3.57 -4.80 -1.95 -2.06 -1.81 -2.11 -2.90 -2.36 -0.99 -2.85 Si/(Si+Al) 0.76 0.75 0.67 0.68 0.69 0.67 0.67 0.68 0.67 0.68 Ca/(Ca+Na+K+Sr) 0.61 0.60 0.99 0.94 0.95 0.99 0.97 0.92 0.96 0.98 Na/(Ca+Na+K+Sr) 0.05 0.01 0.01 0.02 0.01 0.01 0.01 0.02 0.00 0.01 K/(Ca+Na+K+Sr) 0.24 0.18 0.00 0.03 0.03 0.01 0.00 0.05 0.03 0.01 Sr/(Ca+Na+K+Sr) 0.11 0.20 0.00 0.01 0.00 0.00 0.01 0.01 0.01 0.00


continue Table 2.4

Sample no. TW20 35843 A8 A4 Arvigo1 Arvigo13 Arvigo13 TW02 TW01.2 35370 Analysis no. 2 3 3 1 6 3 10 2 4 1 wt.% Sco Sco Sco Sco Sco Sco Sco Stb Stb Stell SiO2 45.72 44.64 45.67 45.39 45.64 46.02 46.20 58.12 63.07 59.13 Al2O3 24.86 24.44 24.70 25.14 24.65 24.96 25.14 15.21 17.09 14.12 CaO 14.05 13.63 13.90 13.69 14.14 14.07 13.99 7.75 8.72 7.87 SrO 0.02 0.04 0.00 0.10 0.00 0.04 0.00 0.06 0.00 0.00 BaO - - - 0.00 0.00 0.00 0.01 - - - Na2O 0.04 0.09 0.22 0.09 0.10 0.14 0.10 0.41 0.91 0.02 K2O 0.01 0.02 0.00 0.00 0.00 0.02 0.00 0.07 0.02 0.04 H2O 15.25* 17.09* 15.45* 14.01 15.38* 14.65* 14.55* 18.37* 10.18* 18.80* Totala 100.00 100.00 100.00 98.43 100.00 100.00 100.00 100.00 100.00 100.00 Si 24.298 24.249 24.330 24.200 24.312 24.296 24.325 27.546 27.255 28.062 Al 15.569 15.647 15.507 15.797 15.476 15.532 15.599 8.495 8.705 7.898 Ca 8.000 7.933 7.932 7.819 8.070 7.961 7.890 3.938 4.035 4.002 Sr 0.006 0.013 0.000 0.032 0.000 0.013 0.000 0.018 0.000 0.000 Ba - - - 0.000 0.000 0.000 0.001 - - - Na 0.044 0.095 0.228 0.091 0.103 0.139 0.104 0.378 0.763 0.018 K 0.005 0.014 0.000 0.000 0.000 0.012 0.003 0.041 0.013 0.024 O 80 80 80 80 80 80 80 72 72 72 H2O 27.035 30.966 27.451 27.700 27.325 25.798 25.550 29.035 14.671 29.760 E%b -3.08 -2.30 -3.64 0.01 -4.73 -3.66 -1.82 2.00 -1.81 -1.84 Si/(Si+Al) 0.61 0.61 0.61 0.61 0.61 0.61 0.61 0.76 0.76 0.78 Ca/(Ca+Na+K+Sr) 0.99 0.98 0.97 0.98 0.99 0.98 0.99 0.90 0.84 0.99 Na/(Ca+Na+K+Sr) 0.01 0.01 0.03 0.01 0.01 0.02 0.01 0.09 0.16 0.00 K/(Ca+Na+K+Sr) 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.01 0.00 0.01 Sr/(Ca+Na+K+Sr) 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 Sample no. 35370 35843 35843 R1 R1 Fi1 DT A4 A6 A6 Analysis no. 5 7 10 1 2 7 9 5 4 7 wt.% Stb Stb Stb Stb Stb Stb Stb Stb Stb Stb SiO2 58.11 61.46 61.62 55.94 56.91 57.40 61.36 58.15 56.17 61.80 Al2O3 14.14 15.64 15.54 15.37 15.36 17.21 15.99 15.14 15.04 17.11 CaO 7.74 8.52 8.52 7.23 7.62 8.65 7.41 7.93 7.72 8.55 SrO 0.00 0.00 0.00 0.04 0.00 0.17 0.00 0.03 0.09 0.06 BaO - - - 0.00 0.02 0.00 0.03 0.06 0.00 0.08 Na2O 0.27 0.35 0.23 0.36 0.11 0.30 0.76 0.56 0.52 0.81 K2O 0.03 0.06 0.07 1.50 0.95 0.87 1.26 0.00 0.04 0.05 H2O 19.71* 13.90* 13.91* 20.64 20.64 15.34* 13.17* 18.06* 20.36* 11.46* Totala 100.00 100.00 100.00 100.00 100.00 100.00 100.00 100.00 100.00 100.00 Si 27.932 27.627 27.691 27.185 27.342 26.579 27.527 27.509 27.353 27.133 Al 8.010 8.286 8.230 8.803 8.698 9.392 8.454 8.443 8.630 8.854 Ca 3.986 4.103 4.102 3.765 3.923 4.291 3.562 4.017 4.030 4.022 Sr 0.000 0.000 0.000 0.011 0.000 0.046 0.000 0.008 0.025 0.014 Ba - - - 0.000 0.004 0.000 0.005 0.011 0.001 0.013 Na 0.252 0.305 0.200 0.339 0.102 0.269 0.661 0.509 0.487 0.692 K 0.018 0.034 0.040 0.930 0.582 0.514 0.721 0.001 0.023 0.029 O 72 72 72 72 72 72 72 72 72 72 H2O 31.598 20.848 20.854 33.454 33.074 23.691 19.706 28.497 33.080 16.789 E%b -2.81 -3.05 -2.54 -0.52 1.88 -0.98 -0.88 -1.61 -0.14 0.39 Si/(Si+Al) 0.78 0.77 0.77 0.76 0.76 0.74 0.77 0.77 0.76 0.75 Ca/(Ca+Na+K+Sr) 0.94 0.92 0.94 0.75 0.85 0.84 0.72 0.89 0.88 0.85 Na/(Ca+Na+K+Sr) 0.06 0.07 0.05 0.07 0.02 0.05 0.13 0.11 0.11 0.15 K/(Ca+Na+K+Sr) 0.00 0.01 0.01 0.18 0.13 0.10 0.15 0.00 0.01 0.01 Sr/(Ca+Na+K+Sr) 0.00 0.00 0.00 0.00 0.00 0.01 0.00 0.00 0.01 0.00 *H2O calculated by difference. aTotals include traces of Mg, Ti, Mn and Fe. b E % = 100*((Al)-(Na+K)+2(Mg+Ca+Sr+Ba)/(Na+K)+2(Mg+Ca+Sr+Ba)), measure of charge balance


2.5.3.2. Heulandite-Ca

The heulandite group of minerals is represented by heulandite (Na,K)Ca4(Al9Si27O72)

•24 H2O and clinoptilolite (Na,K)6(Al6Si30O72) •24 H2O (Armbruster & Gunter,

2001). In Alpine fissures only heulandite occurs as member of the heulandite group.

Heulandite forms crystals up to 12 mm in size, but the average size of the crystals is 1

to 4 mm. The monoclinic crystals of space group 2/m occur in tabular habit parallel

{010} and with its typical coffin-shaped appearance (Fig. 2.6b, e). Heulandite crystals

are transparent to translucent and colorless or white in color and have a subconchoidal

to uneven cleavage.

Fig. 2.9: Extra-framework cations (Ca-Na-K-Sr) in Alpine zeolites. All zeolites are dominated by Ca. Heulandite and chabazite incorporate a significant amount of Sr of up to 30 mole%.

Representative analyses of heulandite are given in Table 2.4. The average

composition of heulandite, determined by 46 microprobe analysis from 9 different

samples (Table 2.2), shows the composition of Ca3.09Na0.12K1.01Sr0.80(Al8.81Si27.13O72)

•24 H2O, which is very similar to the heulandite composition

(Ca3.37Na0.07K0.88Sr0.55(Al8.42Si27.49O72) •24 H2O) determined by Armbruster et al.

(1996) from Gibelsbach in the western Aar Massif (Fig. 2.2; Table 2.2). Heulandite

from Alpine fissures is heulandite-Ca (Coombs et al., 1998), with significant amounts


of Sr and K (Fig. 2.9; Table 2.4). Ca occupies on an average 62 mole% of all extra-

framework sites, 16 mole% are occupied by Sr (maximum 29 mole%) and K is found

on 20 mole% of the sites (maximum 31 mole%). Na is below 10 mole%, in most

samples Na occurs, like Ba, Mg and Fe in traces only (Fig. 2.9).

Heulandite is often zoned which Ca decreasing and K and Sr increasing from

core to rim. Samples (A8.1, A8.2, DT) from Arvigo and Drumtobel/Sedrun are low in

Sr and enriched in Na compared with the other samples (Fig. 2.9). The Si/(Si+Al)

content range between 0.74 and 0.77 (Fig. 2.10), which agrees with the definition of

heulandite (Coombs et al., 1998), that can be distinguished by clinoptilolite, 0.8 <

Si/(Si+Al).

2.5.3.3. Laumontite

Laumontite is a monoclinic (space group C2/m) zeolite. It forms thin, elongated fibers

or prisms elongated along the c-axis with a squared cross-section. The common

crystal form of laumontite is the {110} prism. Commonly twinning occurs on {100}

to form “swallow tail” or “V” twins. It is normally white with a common length

between <1 to 15 mm. The cleavage is perfect, with a slight pearly luster on the broad

cleavage surface. Laumontite is the most widespread zeolite in Alpine fissures,

exposed underground in tunnels sections or in active quarries. Because the mineral

decomposes by dehydration at room temperature and decays to a powdery mass,

laumontite occurs rarely in surface outcrops.

The composition of laumontite (79 analysis from 14 samples; Table 2.2 & 2.4)

formed in Alpine fissures is close to endmember composition Ca4(Al8Si16O48) •18

H2O (Armbruster & Kohler, 1992). Ca is the dominant extra-framework cation

(average value of 96 mole%), with Na and K typically below 5 mole% (Fig. 2.9;

Table 2.4). Maximum values for K and Na are 6 and 7 mole%, respectively. Other

elements occur only in traces. The extra-framework cations Na and K increase from

core to rim of zoned crystals. The Si/(Si+Al) content varies only slightly between

0.67 and 0.69 (Fig. 2.10), with an average ratio of 0.68. Alkalis increase with

increasing Si and decreasing Ca and Al during growth, related to the coupled

substitution of Si4+ + (Na+, K+) = Al3+ + Ca2+.


Fig. 2.10: R2+ - R+ - Si compositional diagram of Alpine zeolites. Si/Al ratio increases in chronological order.

2.5.3.4. Scolecite

Fibrous scolecite occur as white crystals with vitreous or slightly silky luster, forming

characteristic radiating sprays. The transparent to translucent crystals are thin

prismatic with squared cross sections. Scolecite needles range between 1 and 20 mm

in length, normally 3 to 6 mm. Scolecite (Ca8(Al16Si24O80) •24 H2O) has the same

structural framework as natrolite (Na16(Al16Si24O80) •16 H2O) and mesolite

(Na16Ca16(Al48Si72O240) •64 H2O) (Armbruster & Gunter, 2001). Minerals of this

group are distinguished by a Ca2+ + H2O ↔ 2Na+ substitution and by symmetry

(Gottardi & Galli, 1985). The variation in composition of natrolite, mesolite and

scolecite is very small. (e.g. Deer et al., 2004). 42-microprobe analyses from 7

different samples (Table 2.2) indicate no major chemical variation (Fig. 2.9). Ca

occupies an average of 99 mole% of the extra-framework cation sites (Fig. 2.9).

Minor amounts of Na up to 4 mole%, but in most analyses around 1 mole% can be

observed. In contrast K and Sr are not significantly incorporated in the framework

structure of scolecite and therefore occur only in traces, like Ba, Mg and Fe. The


small substitution range can also be recognized in the Si/(Si+Al) ratio, whereas the

ratio varies in a small range between 0.60 and 0.61 (Fig. 2.10).

2.5.3.5. Stilbite/Stellerite

A complete solid solution exists between stellerite (Ca4(Al8Si28O72) •28 H2O) and

stilbite (NaCa4(Al9Si27O72) •30 H2O). Although the tetrahedral framework of stellerite

and stilbite is identical, with the symmetry Fmmm (Gottardi & Galli, 1985), the

overall crystal symmetry is different due to different locations of the extra framework

cation sites. Stoichiometric stellerite is orthorhombic, with only one extra framework

cation site fully occupied by Ca. In contrast there is an additional extra framework

cation site in stilbite, occupied by Na. The additional Na site in stilbite leads to a

reduction of symmetry from orthorhombic Fmmm in stellerite to monoclinic C2/m in

stilbite (Gottardi & Galli, 1985; Quartieri & Vezzalini, 1987).

The habit of stilbite and its chemical composition is very variable. They occur as

thick tabular crystals flattened on {010} with pointed terminations, as sheaf like

aggregates or globular aggregates, as characteristic “bow-tie” crystals or as clusters of

elongated six-faced crystals. Stilbite is transparent to translucent, colorless to white in

color, whereas stilbite from surface outcrops in the Riental valley hosted in quartz

veins (Table 2.1; no 49) occurs as light reddish globular aggregates. Crystal sizes of

stilbite/stellerite range from 1 to 12 mm, with a frequent size between 3-5 mm.

Historically and in old museum collections stilbite and stellerite were summarized by

the synonym “desmine”, which is not longer an official term for zeolites (Coombs et

al., 1998). Single x-ray diffraction analysis of “stilbite” from Gibelsbach by

Armbruster et al. (1996), has shown that zeolites described stilbite from Alpine

fissure are stellerite. Between stilbite and stellerite exists a complete solid solution

(Fridriksson et al., 2001) and it is not always possible to determine, whether it

stellerite or stilbite is present. The chemical composition shows a large variation (Fig.

2.9).

112 microprobe analyses from 20 samples (Table 2.2 & 2.4) were measured to

evaluate stilbite composition. Different groups of stilbite can distinguished on the

basis of the K content (Fig. 2.9). One sample from the Gotthard road tunnel (35370) is

pure Ca-endmember thus it can be classified as stellerite (Coombs et al., 1998).


Chemically stellerite and stilbite-Ca can be rather similar, so the formal mineral name

stellerite is restricted to specimens of nearly stoichiometric Ca4(Al8Si28O72) •28 H2O.

It is distinguished from stilbite-Ca by containing low Na2O and high SiO2. The

maximum amount of Na, K, Mg and Fe in stellerite is below 0.2 atoms per formula

unit (apfu) (Coombs et al., 1998) and the sample 35370 has in average 0.18 apfu on

the basis of 72(O), well within the stellerite field (Coombs et al., 1998). The

Si/(Si+Al) content for stellerite is constant with a ration 0.78 (Fig. 2.10). All other

samples show Na-Ca zoning, with Ca high in the core and low at the rim. Sodium

concentration shows an inverse pattern. Using the K content, 2 different groups can be

distinguished. Samples from Gotthard road tunnel and Gotthard NEAT tunnel (Table

2.2) have a K content of less than 3 mole% of the extra-framework cation side,

whereas samples from surface outcrops (Table 2.2) can be distinguished by an

elevated K content between 8 and 20 mole% of the extra-framework site (Fig. 2.9). A

coupled substitution of Si4+ + K+ = Al3+ + Ca2+ can be seen in stilbite samples from

tunnel sections, with a maximum Na value of 0.84 (apfu), whereas the Si/(Si+Al)

ratio range between 0.75 and 0.78, with an average ratio of 0.77 (Fig. 2.10). Stilbite

samples from the surface indicate a coupled substitution including Na and K. These

samples are characterized by elevated K contents up to 0.93 apfu and slightly higher

Na values. Si/(Si+Al) ratio ranges between 0.74 and 0.77, with an average ratio of

0.76 (Fig. 2.10).

2.6. DISCUSSION

The distribution patterns of the Alpine zeolites (Fig. 2.2c) show that the vein zeolites

are not products of prograde zeolite facies metamorphism. They are rather related to

the cooling and uplift history of the Alpine orogen. The zeolites formed in late

fractures in areas that have experienced much higher metamorphic peak conditions

(Fig. 2.11) and also underwent pervasive ductile deformation. Zeolites formed when

the rocks entered the brittle deformation regime and cooled to temperature below

about 250˚C (see below). The ductile-brittle transition is located at 350-400˚C and can

be tied to the formation of fissure assemblages that indicate the highest temperature

(about 400˚C, Mullis et al., 1994) corresponding to a depth of about 12 km.

Consequently early fissures do not contain zeolites but rather a sequence of minerals


and assemblages that are summarized in Table 2.3. The onset of zeolite formation

corresponds to the point where the local cooling path of the different areas entered the

“zeolite window” (Fig. 2.11). This may have happened at different times in the

different areas (Fig. 2.11; all cooling path merge eventually, which may not

necessarily have been the case). Zeolites overgrowing earlier formed fissure minerals

(e.g. quartz; Figs 2.4 & 2.5) suggest that the fracture formed before the cooling path

entered the zeolite window. Monomineralic zeolite fissures indicate, in contrast, that

the fracturing occurred in the zeolite window (Fig. 2.11).

Fig. 2.11: Schematic sketch of the T-t evolution of tectonic units in the Central Alps in relation to fissure formation and the timing of zeolite growth. (a) The T-t paths of individual tectonic units reflecting an increase of the Alpine peak metamorphism from north to south. The southern units have reached the ductile regime, northern units deformed brittle. All units reached temperatures above the zeolite window (except for the parautochtonous cover rocks of the Aar massif). During uplift the units returned to the brittle deformation regime and extension fissures formed (b), subsequently zeolite-absent fissure assemblages developed (c) and finally the units entered the zeolite window (d).

The presented data show that first scolecite, then laumontite, heulandite, chabazite

and finally stilbite precipitated from the hot aqueous fluids with decreasing

temperature along the cooling path. This means that scolecite formed when the

cooling path first entered the zeolite window, stilbite is the low-temperature zeolite

that may still form in deep groundwater environments in the crystalline basement of

the Central Alps such as in 40˚C fissures of the Gotthard rail base tunnel (Seelig et

al., 2007). Zeolite formation thus can be related to the T-interval from 250˚C to 50˚C

with a characteristic mean temperature of about 200˚C where most of the dominant

zeolite laumontite formed.


Heulandite and chabazite require Sr in the fluid to become stable together with a

pure Ca zeolite. Sr-bearing heulandite and chabazite is associated with stilbite or

scolecite (Fig. 2.6). From this it also follows that heulandite associated with scolecite

(Fig. 2.6e) indicate a higher temperature regime compared to chabazite associated

with stilbite (Fig. 2.6a) that point to lower temperature conditions.

Chemical zoning and compositional variation of laumontite, heulandite, chabazite

and stilbite indicate changes in fluid composition or temperature during growth.

Similar pattern were also observed in Icelandic geothermal system (Fridriksson et al.,

2001), where a Na increase with decreasing temperature during formation of stilbite

was confirmed. Remarkable are two distinct geochemical patterns comparing stilbite

collected in tunnels and at surface outcrops. Because samples from surface outcrops

formed earlier than samples from tunnel sections, the elevated values of K in surface

samples suggests a late temperature dependent ion exchange: Na+ = K+ as a result of

interaction with surface water. The zoning pattern in laumontite, heulandite and

chabazite seems to be related to a decrease in temperature during formation. But late

ion exchange with surface water after the formation of zeolite could also explain the

observed compositional patterns in zeolite. Present day deep fluids in the Aar and

Gotthard basement are dominated by Na2SO4 and Na2CO3 (Seelig et al., 2007) and

would thus support the presence of an ion exchange component.

A remarkable result of this study is the observation of significant differences of

the observed rare occurrences of zeolites in surface outcrops and the very abundant

occurrence of zeolites in tunnel fissures. This is particularly the case for the

abundance of laumontite in fissures. The mineral is extremely rare at the surface and

dominant in subsurface samples. The absence of laumontite in surface outcrops is a

consequence of the instability of the zeolite in the presence of air. Laumontite

exposed to low humidity air at low temperature partially dehydrates and decomposes

(Blum, 1843; Armbruster & Kohler, 1992) and is easily eroded. We conclude from

our data and observations that Ca-zeolites are the prime alteration products together

with chlorite that form from the fundamental reaction (1):

granite + meteoric water = laumontite (Ca-zeolite) + chlorite + deep groundwater (1)


It is important to recognize that granite alteration by water-rock interaction along

fissures at temperatures of 150 ± 100˚C in a cooling orogen does not produce clay

minerals with the exception of chlorite.

2.6.1. Reactions and processes of zeolite formation

The geologic overall context of the reported zeolites suggests that they precipitated

from aqueous solutions circulating in open fissures. The zeolite crystals do not grow

directly on the expense of primary minerals of unaltered granite. The structures and

textures described imply that primary minerals of basement rocks dissolved along

fractures into hot water (aqueous fluid). The hot aqueous liquid reached high degrees

of super saturation with respect to a succession of Ca-zeolite minerals upon cooling,

as observed and documented from many basement fluids at temperatures comparable

to the imposed main temperature of zeolite formation of the Central Alps (Urach

geothermal site, German continental deep drilling site KTB; Stober & Bucher, 2004,

2005). The process is a classic dissolution-precipitation process observed in many

metamorphic environments (Carmichael, 1969).

The observed Ca-zeolites require the presence of Ca, Al and Si in the hot fluid to

form. Source minerals of the rock matrix for necessary constituents are plagioclase,

clinozoisite, quartz and earlier formed fissure calcite. The dissolution of plagioclase

with an appreciable component of anorthite such as oligoclase in the samples from

Arvigo (Fig. 2.12) provides a source for Ca2+.

plagioclase + H2O ⇒ albite + Ca2+ + 2 AlO2

- + 2 SiO2,aq + H2O (2)

Na4CaAl6Si14O40 + H2O ⇒ Na4Al4Si12O32 + Ca2+ + 2 AlO2- + 2 SiO2,aq + H2O

Reaction (2) describes the dissolution of An-component of oligoclase, release of its

component into the hot fluid and the production of a residual solid phase namely

albite. Continuous albitization is accompanied by a porosity increase (Fig. 2.12),

which provides the necessary permeability for water infiltration. The albitization

reaction proceeds, as long fluid pathways are available to transport fluid

undersaturated with plagioclase to the reaction interface of plagioclase and albite. In


the Arvigo samples plagioclase parent grains in the leaching zone are usually

completely replaced by albite. Preserved plagioclase exists in the unaltered host rock,

which could be an effect of the lack of porosity due to higher temperature

transformation reaction (e.g. biotite-chlorite), or because fluid pathways get clogged

by consumption of earlier formed porosity. The components released by plagioclase

dissolution are continuously precipitated as zeolite in the open fissure by the model

reaction (3):

Ca2+ + 2 AlO2

- + 4 SiO2,aq + 4 H2O ⇒ laumontite (3)

Ca2+ + 2 AlO2- + 4 SiO2,aq + 4 H2O ⇒ CaAl2Si4O12 •4 H2O

Fig. 2.12: EMPA images showing products of albitization process (Pl + Qtz + H2O = Ab + Lmt). (a) BSE image showing pores in albite formed by albitization of primary plagioclase. Laumontite is easily recognized by its cleavage. A porosity of ~ 14 % was measured by image analyzing methods using IMAGEJ. (b) Ca distribution image showing irregular grain boundary of laumontite.

The additional silica necessary for the formation of zeolite may either derive locally

from dissolution of primary quartz or from externally derived SiO2,aq and added by

the fluid. Reactions (4) and (5) represent generic reactions describing plagioclase

dissolution and albite and zeolite precipitation:

plagioclase + 4 H2O + 2 SiO2,aq ⇒ laumontite + albite (4)

Na4CaAl6Si14O40 + 4 H2O + 2 SiO2,aq ⇒ CaAl2Si4O12 •4 H2O + Na4Al4Si12O32


plagioclase + 7 H2O + 5 SiO2,aq ⇒ stilbite + albite (5)

Na4CaAl6Si14O40 + 7 H2O + 5 SiO2,aq ⇒ CaAl2Si7O18 • 7H2O + Na4Al4Si12O32

Keep in mind that the reactions involve a transport step between dissolution and

precipitation as the later two processes proceed at spatially different locations. The

two reactions represent net balances of reaction (2) and (3) and are thus independent

of pH.

During albitization of oligoclase in Arvigo Ca2+ and Al3+ were transported in

solution from the rock matrix to the open fissure resulting in a porosity increase of 14

% in the matrix rock. This agrees well with the measured porosity in thin section (Fig.

2.12).

Zeolite formation in the Aar Massif requires another source for Ca2+ because

plagioclase in the Aar granite and gneiss is pure albite. Alpine regional

metamorphism reached only greenschist facies metamorphism in the Aar Massif (Fig.

2.11), resulting in a complete transformation of prealpine plagioclase to albite and a

separate Ca-phase such as clinozoisite (epidote) or calcite. In the southern Gotthard

massif plagioclase contains up to 18% anorthite component (Steck, 1976) as a result

of increasing peak metamorphic grade toward the south (Fig. 2.11).

Thus Ca-zeolite may precipitate in fissures of the Aar massif from dissolved

components that have been provided by the dissolution of clinozoisite and albite.

Clinozoisite (epidote) is present in the host rock as result of Alpine greenschist facies

metamorphism.

2 clinozoisite + 15 SiO2,aq + CO2 + 20 H2O ⇒ 3 stilbite + calcite (6)

2 Ca2Al3Si3O12(OH) + 15 SiO2,aq + CO2 + 20 H2O ⇒ 3 CaAl2Si7O18 •7 H2O + CaCO3

This plausible reaction mechanism co-precipitates zeolite and calcite. The assemblage

is common in fissures (Table 2.2; TW03, 30992, 36728). If the fluid does not reach

calcite saturation, the zeolite is the single product phase of the reaction and models

the common zeolite (Lmt, Stb) fissures.


In some fissures, the dissolution of secondary calcite, originally formed from

plagioclase breakdown during Alpine greenschist facies metamorphism, may have

provided Ca2+ for zeolite formation during fluid-rock interaction:

2 albite + calcite + SiO2,aq + 7 H2O ⇒ stilbite + 2 Na+ + CO3

2- (7)

2 NaAlSi3O8 + CaCO3 + SiO2,aq + 7 H2O ⇒ CaAl2Si7O18 •7 H2O + 2 Na+ + CO32-

Reaction (7) consumes calcite and forms stilbite that is accompanied by an increase in

pH and the total of dissolved solids (TDS). The proposed reaction is supported by the

high Na+ and CO32- concentrations, high pH, and high degrees of oversaturation with

respect to stilbite in deep groundwater reported from the Gotthard rail base tunnel

(Seelig et al., 2007).

2.6.2. Assemblage stability and phase relationships involving zeolites

Assemblage stability and phase relationships involving zeolites have been computed

with the computer program Domino/Theriak (de Capitani & Brown, 1987) using the

thermodynamic data by Bermann (1988), Evans (1990), Frey et al. (1991) and

Maeder & Bermann (1991). For scolecite and chabazite we adopted the

thermodynamic data of Johnson et al. (1983) and Ogorodova et al. (2002) and the

heat capacity function was predicted from the equation [4Sco + 2Qtz = 3Lmt + An]

and [7 Cha + 16 Qtz = 6 Stb + An], respectively.

The P-T model for the system CaAl2Si2O8–SiO2–H2O (Fig. 2.13) agrees well

with field observations (Table 2.3). The predicted sequence of zeolites along a

cooling path Lmt - Heu -Stb corresponds to the observed sequence in Alpine fissures.

On Fig. 2.13 scolecite is restricted to low-pressures and would not form along the

deduced cooling path in clear contrast to the field evidence. We conclude that the

thermodynamic data for Sco needs improvement.

Fissure zeolites form at hydrostatic pressures below about 100 MPa

corresponding to 10 km depth and temperatures below 400˚C. The effect of pressure

on zeolite formation is thus assumed to be negligible. The absence of pumpellyite in

Alpine fissures is consistent with low-pressure zeolite formation. In basaltic systems


with excess Qtz and Chl the laumontite-pumpellyite association is not stable at

pressures below 100 MPa (Cho et al., 1986).

Fig. 2.13: Assemblage stability diagram in the Ca-Al-Si-O-H system. A retrograde cooling path adapted from Mullis et al. (1994) shows cooling during exhumation. Bulk: CaAl2Si2O8 + 5 SiO2 + 50 H2O; An + Qtz + water.

Zeolite forming reactions require dissolved silica in the hot water on the reactant side.

Consequently the reactions depend on the activity of SiO2 in the hot fluid (Fig. 2.14).

The computed phase fields show the same sequence as observed in nature where the

sequence also corresponds to a chronological order that can be related to a cooling

path. The phase topology has been computed at 100 MPa and it shows a common

boundary between Lmt and Stb. At lower pressure (10 MPa on Fig. 2.14) an inverted

topology with a common boundary between Heu and Cha is predicted to be more

stable than Lmt - Stb. The 100 MPa topology is consistent with the presented data,

since laumontite-stilbite is a frequently observed zeolite assemblage (Table 2.2) and

heulandite-chabazite has not been found.


During cooling and uplift (Fig. 2.11) zeolites start to form in fissures with the

first appearance of scolecite and laumontite at a temperature of 280-300°C. Assuming

a plausible crustal temperature gradient of 30˚C km-1 this would correspond to 10 km

depth and a hydrostatic pressure of about 100 MPa.

Fig. 2.14: Equilibrium T- aSiO2 diagram at P = 10 MPa and 100 MPa, for the Ca-Al-Si-O-H system. Standard state for aSiO2 is the pure stable SiO2 solid at P and T of interest. Thus quartz saturation is given at aSiO2 = 1. Note the topology inversion [Heu, Cha]⇔[Lmt, Stb] between 10 and 100 MPa pressure. Also note that pure chabazite-Ca is predicted to form from fluids with SIQtz < 0.

Temperature information can also be derived from chlorite geothermometry. As

described above, chlorite forms earlier or with the zeolites and therefore provides

maximum temperatures for zeolite formation. A temperature of 325°C ± 23°C (n =

170) has been calculated using the geothermometric calibration of Chatelineau

(1988). The empirical calibration is based on the Al content on the tetrahedral sites.

The temperature of 325˚C is interpreted as the maximum temperature for the zeolite

window (Fig. 2.11). It probably marks the point when the crust has cooled along the

path to start the first zeolites to form.


The low temperature limit of zeolite formation is not easy to establish. Stilbite is

predicted to be thermodynamically stable in the presence of water at ambient P-T

conditions. This is consistent with the observation that recent tunnel waters (40°C)

from the Gotthard NEAT tunnel is strongly oversaturated with respect to stilbite,

suggesting that stilbite formation is still under progress (Seelig et al., 2007).

Fig. 2.15: Equilibrium T-XCO2 diagram at P = 10 MPa, 50 MPa and 100 MPa, for the Ca-Al-Si-O-H-C system with excess quartz. (Bulk: CaAl2Si2O8 + 5 SiO2 + 50 H2O). Zeolites tolerate decreasing amounts of CO2 along cooling path.

2.6.3. Fluid composition

Zeolites in Alpine fissures are very irregularly distributed. For example zeolites in

fissures of the central Aar granite (ZAgr) are frequent in some and rare in other

regions. However, the bulk rock composition of ZAgr varies very little over the

outcrop area of nearly 100 km (Labhart, 1977). Consequently, the different frequency

of zeolite occurrence could be related to variations in externally derived fluids on the

fractures. The consequence is that in such a case the rocks would be unable to buffer

the fluid composition. This is in contrast to modern groundwaters in the basement,


which are controlled by the local lithology. It is obvious that H2O and CO2 will have a

prime control on zeolite reactions (Zen, 1961). At relatively high pCO2, Ca-zeolites are

replaced by clay minerals (Fig. 2.15).

It is evident that zeolites require low aCO2 conditions (Fig. 2.15). With increasing

CO2 in the fluid zeolites are replaced by kaolinite at low temperature and by margarite

and calcite at high temperature. Stilbite does not tolerate much CO2; its presence

indicates low CO2 fluids. CO2 tolerance increases with temperature in the order Stb <

Heu < Lmt < Sco. Zeolite CO2 tolerance decreases with increasing pressure and at

pressure condition at about 100 MPa, zeolites decompose (or will not form) in fluids

witch exceed 5 mole% CO2.

It can be concluded from the relationships that the CO2 content of the fluid

controls the presence or absence of zeolites in a particular local area. The composition

of fluid inclusions may thus give important information of fluid compositions during

zeolite growth. Unfortunately, no fluid inclusions are preserved in zeolite minerals.

Abundant fluid inclusions are present in Alpine fissure quartzes (Mullis, 1995).

However, as documented above fissure quartz is clearly older than zeolites. Fluids

trapped in earlier quartz are thus unrelated to zeolite forming fluids.

However, the fluids trapped in Alpine fissure quartzes show a very distinct N-S

compositional trend (Frey et al., 1980). The Central Alps can be divided into four

zones (from North to South: the higher hydrocarbon zone, 100-200 °C; the methane

zone, 210-270 °C; the H2O zone, 240-430 °C; and the CO2 zone, 300-450 °C),

whereas the Aar Massif and the northern part of the Gotthard Massif characterized by

fluids dominated by H2O (>80 mole% H2O, <10 mole% CO2). Zeolites reported and

discussed here formed in the CO2-poor H2O-dominated zone. The zeolite-rich zone of

the eastern part of the Aar Massif and the northern part of the Gotthard Massif

indicate H2O dominated fluids (Mullis et al., 1994). In the Penninic Alps to the South

fluids tend to be dominated by CO2, explaining the absence of zeolite species in this

region. The exceptional Arvigo zeolite locality in the Calanca valley within the CO2-

rich region can be explained by locally H2O-rich fluids. Detailed studies have shown

that the fluid composition changes from early CO2 dominated fluids (40 mole% CO2)

to late stage CO2-absent fluids (Wagner et al., 2000 a, b). Mullis et al. (1994)

observed that in general, every distinct fissure followed the same general retrograde


fluid evolution path, leading to a final water-rich fluid. Early H2O-poor fissure fluids

have been unable to form zeolites and are thus responsible for the lack of the high-

temperature Ca-zeolite wairakite.

Fig. 2.16: Activity diagram for the system Ca-Al-Si-O-H depicting mineral stability as a function of Log aSiO2,aq versus Log aCa2+/a2

H Standard state for aqueous species is a hypothetical one molal solution at infinite dilution. Quartz, apophyllite and wollastonite saturation and saturation of calcite as a function of pCO2marked with dashed lines. Deep waters from crystalline basement are shaded in gray: a = water samples from the Gotthard NEAT tunnel (Seelig et al., 2007), b = water samples from a granite located in Stripa/Sweden (Nordstrom et al., 1989), c = range of typical deep water from crystalline basement (Bucher & Stober, 2000), d = deep waters from the Urach drill site (Stober & Bucher, 2004).

Still, fluid inclusions in quartz crystals represent fluid compositions prior to zeolite

formation. Regarding present-day fluid compositions in crystalline basement rocks of

the continental crust, fluids are generally oversaturated in respect to zeolites. Deep

continental fluids have the potential to form zeolites (Stober & Bucher, 2004, 2005).

High-pH waters from the NEAT tunnel in the basement of the Aar Massif (Seelig et

al., 2007), water from the basement at Stripa, Sweden, (Nordstrom et al., 1989), Bad

Urach (Stober & Bucher, 2004) in the Black Forest basement (Bucher & Stober,

2000) are all oversaturated in respect to zeolites (Fig. 2.16). aSiO2,aq, pH and other

fluid composition parameters may control the zeolite that will form in addition to

temperature. Textural relationships presented above show that the fluids are not

quartz-saturated during zeolite formation. Variations in aSiO2,aq control the zeolite

assemblages along the cooling path. The following sequence of zeolites is present

with decreasing aSiO2,aq: Lmt - Heu – Stb, Lmt – Stb and Lmt – Cha – Stb (Fig. 2.14).

Thus the regional absence of zeolite and the observed different local sequences of


zeolites formed during cooling and uplift can be related to local variations in CO2 and

SiO2 content of the fissure fluid.

2.7. CONCLUSIONS

Zeolites are common minerals in late Alpine fissures of crystalline basement rocks of

the Central Alps. The fissure zeolites precipitated from hot aqueous fluids that

developed their load of dissolved solids from fluid-rock interaction. The zeolites form

dense mats and crusts on fissure walls in surface outcrops and in rail and road tunnel

up to 2000 meter below the surface. Three regions are particularly rich in zeolite-

bearing fissures in granites and gneisses: (1) in the central and eastern part of the Aar-

and Gotthard Massif, including the Gotthard road tunnel and the Gotthard-NEAT

tunnel, (2) Gibelsbach/Fiesch with a large supply of zeolites in a fissure breccia

between Aar Massif and Permian sediments, and (3) in Penninic gneisses of the

Simano nappe at Arvigo (Val Calanca). Ca-zeolites precipitated from the low-CO2

aqueous fluid with decreasing temperature in the following sequence: scolecite,

laumontite, heulandite, chabazite and stilbite. Within this sequence an increase of the

Si/Al ratio of the zeolites can be observed. Heulandite and chabazite incorporate

significant amounts of extra components with respect to the Ca-Al-Si-O-H system.

Stilbite samples from tunnel sections and from surface outcrops show distinctly

different compositions that indicate a post-growth K-Ca exchange.

The components needed for zeolite formation were derived from the reaction of

the hot fluid in the fissures with “primary” rock. The reaction dissolved plagioclase in

rocks of the amphibolite facies (e.g. Arvigo) and precipitated albite and zeolite. In

greenschist facies granites and gneisses of the Aar Massif, the reaction dissolved

clinozoisite-epidote and/or calcite and precipitated zeolites that form the observed

zeolite veins of this region. The dissolution process is accompanied with a porosity

increase in the leaching zone. The remarkable and astonishing lack of zeolites in late

fissures in many other regions with the potential for zeolite fissures is difficult to

explain but could be related to pCO2 above critical threshold value that makes zeolite

formation impossible.


2.8. ACKNOWLEDGMENTS

We are grateful to Peter Amacher who provided high-quality mineral specimens from

the Gotthard NEAT tunnel. Beda Hoffmann and Peter Vollenweider from the Swiss

Natural History Museum in Bern, giving us the possibility to study their mineral

collection. Special thanks go to the technicians and staff of the Mineralogical-

Geochemical Institute, University of Freiburg: Isolde Schmidt for her help during

sample preparation and for the XRD analyses; Melanie Katt for the careful

preparation of fragile thin sections; Hiltrud Müller-Sigmund for her useful advise

during EMP analyses and her patience with us at the electron microprobe. A special

thanks deserved to the Friedrich Rinne foundation for the financial support

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LOW TEMPERATURE WATER-ROCK INTERACTION -ARVIGO 76

3. POROSITY EVOLUTION, MASS TRANSFER AND

PETROLOGICAL EVOLUTION DURING LOW

TEMPERATURE WATER-ROCK INTERACTION IN

GNEISSES OF THE SIMANO NAPPE - ARVIGO, VAL

CALANCA, GRISONS, SWITZERLAND


3.1. ABSTRACT

Low-grade mineral assemblages are the key to the appreciation of water-rock

interaction in hydrothermal and geothermal systems located in granites and gneisses.

Zeolite formation is an important process in rocks of the continental crust. It takes

place at temperatures below 250°C under hydrothermal conditions. A detailed study

of the mineralogical, chemical and petrological evolution of crystalline basement

rocks in Arvigo was performed to assess information about the evolution of fluid-rock

interaction during uplift of the Alpine orogen. The Arvigo fissures contain the

assemblage epidote, prehnite, chlorite and various species of zeolites.

Fluid rock interaction takes place along a retrograde exhumation path which is

characterized with decreasing temperature by: (1) coexisting prehnite/epidote, that

reveals temperature conditions of 330 – 380 °C, (2) chlorite formation at temperature

of 333 ± 32 °C and (3) formation of zeolites <250 °C. The formation of secondary

minerals is related to the hydrothermal replacement reaction during albitization and

chloritization that releases components for the formation of Ca-Al silicates and form a

distinct reaction front. The fluid-rock interaction is associated with a depletion of

Al2O3, SiO2, CaO, Fe2O3 and K2O in the altered wall rock. The reaction is associated

with an increase in porosity up to 14.2 ± 2.2 %, caused by the volume decrease during

albitization and the removal of chlorite. The propagation of the sharp reaction front

through the gneiss matrix occurred via a dissolution-reprecipitation mechanism.

Zeolite formation is tied to the plagioclase alteration reaction in the rock matrix,

which releases components for zeolite formation to a CO2-poor, alkaline aqueous

fluid.

Keywords: water-rock interaction, laumontite, prehnite, epidote, albitization, Arvigo, Swiss Alps

3.2. INTRODUCTION

Fracture related fluid-rock interaction in gneisses and granitic rocks under

hydrothermal conditions often causes changes in mineralogy and geochemistry and

has been investigated by numerous authors (e.g. Ferry 1979; Mercolli et al. 1984;


Parneix and Petite 1991; Ciesielczuk and Janeczek 2004). Ca-Al-silicates, like

epidote, prehnite and zeolites thereby formed in the open space and therefore gives

information about temperature and pressure conditions as well as information about

fluid composition during fluid-rock interaction and precipitation of secondary phases

(e.g. Thompson 1971; Surdam 1973; Liou 1979; Liou 1985; Cho et al. 1986; Bevins

et al. 1991; Young et al. 1991; Rose et al. 1992; Diegel and Ghent 1994; Gianelli et

al. 1998; Faryad and Dianiska 2003). Detailed mineralogical and petrological study of

the low-grade mineral assemblage can give an appreciation of the fluid-rock

interaction and the porosity and permeability evolution, which is an important factor

of geothermal systems and fluid migration in the upper continental crust (e.g. Gianelli

et al. 1998; Neuhoff et al. 1999; Weisenberger and Selbekk 2008).

Ca-Al-silicate formation is a widespread and frequent feature in volcanic rocks of

basaltic to acidic composition and in sedimentary environments, where elements

necessary for the formation of secondary minerals mainly derive from dissolution of

primary glass (e.g. Walker 1960, 1963; Hay 1966, 1977; Hay and Sheppard 1977;

Surdam 1977; Gottardi 1989; Neuhoff et al. 1999; Hay and Sheppard 2001).

However, infrequently similar Ca-Al-silicate phases have been reported in gneisses

and granite, whereas elements most likely released during feldspar alteration (e.g.

Freiberger et al. 2001; Faryad and Dianiska 2003; Weisenberger and Bucher 2009). A

detailed study of the zeolite distribution in the Central Swiss Alps (Weisenberger and

Bucher 2009) has shown that zeolites hosted in gneisses and granites occur frequently

and has to be kept in mind during fluid-rock interaction in the upper continental crust.

The purpose of the paper is to report the results of an investigation of the

hydrothermal alteration and metasomatic albitization of granitic gneisses of the

Simano nappe. Samples were collected in an active quarry in Arvigo (Val Calanca,

Switzerland; Fig. 3.1). Epidote, prehnite and zeolites display the main fissure

minerals. In contrast only three other localities in the Lepontine Alps are known

where zeolites occur (Biasca, Val Baveno, Val Vergeletto road tunnel; Stalder et al.

1998). Hydrothermal alteration in gneisses and granitic rocks is macroscopically seen

as localized metasomatic bleached zone adjacent to fractures, which are filled by

hydrothermal minerals. These expose the interface between generally unaltered and

altered or albitized rock, respectively. This allows us to a detailed study of

petrological changes, mass transfer and porosity evolution during metasomatism of


the upper continental crust, which contains a significant amount of plagioclase. The

results provide an example of how observations of petrological changes, mass transfer

and porosity evolution can be integrated in the geochemical interpretation of fluids in

crystalline rocks of the upper continental crust.


The Arvigo locality is situated in the N-S striking Alpine valley Val Calanca, Grisons

(Fig. 3.1). The rocks exposed in Arvigo belong to the Simano nappe, which is

allocated to the lower Penninic basement nappes of the Central Alps and represents

paleogeographically the southern passive margin of the European plate (Wenk 1955).

The Simano nappe is a metamorphic complex including several metagranitic

bodies of Caledonian and Variscan age (Jenny et al. 1923; Keller 1968; Köppel and

Grünenfelder 1975; Schaltegger et al. 2002), whereas the upper parts of the Simano

nappe mainly consist of pre-Mesozoic gneisses and micaschists, intercalated with

numerous amphibolite and calcsilicate lenses (Schaltegger et al. 2002; Rütti et al.

2005). The large Simano basement nappe is located between the underlying Leventina

nappe and the Adula nappe as hanging wall (Berger et al. 2005), which is separated

by thin Mesozoic metasediments (Fig. 3.1).

Barrovian-type Alpine metamorphism gradually increases from north to south

from lower amphibolite facies condition in the southern Gotthard Massif to upper

amphibolite facies conditions southwards to the Insubric Line (Frey et al. 1980; Frey

and Mählmann 1999). Temperature determination by Engi et al. (1995) suggest a

temperature range in the Simano nappe between 550°C in the north and 700°C

conterminous to the Insubric Line. Corresponding pressure estimates in the Simano

nappe achieve maximum pressures of about 650-700 MPa in a region approximately

20 km north of the Insubric Line. Peak metamorphic conditions for the region around

Arvigo yield a temperature range from 600-680°C and pressure conditions of 550 to

600 MPa (Engi et al. 1995; Todd and Engi 1997; Nagel et al. 2002).


Fig. 3.1: Geological sketch maps: (a) Simplified tectonic map of the Simano nappe complex (modified after Spicher 1980); rectangle marks the section in part b. (b) Simplified geological map of the studied area along the NS trending Val Calanca valley (modified after Berger et al. 2005). (c) Outline of Switzerland, rectangle marks the location of the tectonic sketch in section a.

The rocks are mined as building stones in a quarry south of the town Arvigo (Fig.

3.1). The Arvigo gneisses general strike in NNW-SSE orientation with a dip of 20° to

30° in NE direction (146/30° NE to 158/31° E) that is parallel to the inclination of the

valley. The Arvigo quarry became famous for a large number of Alpine fissure

minerals, which occur in extensional fractures and cavities of the granitic gneiss

(Ruppe 1966; Simonetti 1971; Wagner 1968, 1980, 1981, 1983; Weiß and Forster

1997; Wagner et al. 2000a, b). Fissures and gashes formed by brittle deformation

were generated during exhumation and uplift of the Alpine orogen (Mullis et al.


1994). Fissures are generally perpendicular to the schistosity plane, whereas fissures

are often brecciated due to late deformation stage.

Fig. 3.2: Field relationships and schematic sketches of fracture related hydrothermal alteration. Mineral abbreviation used after Bucher and Frey (2002). (a) Fissure mineral assemblages: Qtz-Kfs-Ep-Sco. (b) Assemblage: Cal-Sco-Lmt, whereas calcite shows corrosion. Coin for scale.(c) Vein photograph from a vein hosted in biotite-rich gneisses (Arvigo/Val Calanca). Hammer for scale. The vein is characterized by a ∼1 cm leaching zone trending in vertical direction, which appears to be lighter, due to the removal of biotite. The open space of the fissure is filled with secondary phases, which consists mainly on chlorite. (d) Schematic sketch of an Arvigo fissure. Extension forces lead to the opening of fissure, which describe pathways for fluids that increase the permeability and drive on fluid flow through the rock fissure. Leaching during fluid-rock interaction change primary mineralogy and geochemistry of the host rock, which is marked by an alteration zone, which grow perpendicular to the fissure orientation. Alteration zone shows an increase in porosity and appears to be lighter due to removal of primary minerals. Fluid, which gets saturated in respect to secondary minerals precipitate this in the open space, which affected a decrease of permeability and can hence stop the fluid movement. (e) Schematic sketch of a zeolite bearing Alpine fissure, exhibit euhedral mineral assemblages. Zeolite species overgrow earlier formed minerals in the following order, as it observed in nature: Qtz-Ep-Prh-Sco.


The Arvigo fissures contain more than 40 different minerals (Weiß and Forster 1997;

Armbruster 2000; Armbruster et al. 2000, Wagner et al. 2000a, b) hosted in fissures,

whereas the main mineral assemblage is characterized by epidote, prehnite and

zeolites (Fig. 3.2). The Arvigo gneisses are penetrated by various small (<10 cm)

aplitic dykes as well as pegmatitic dykes, which consists of coarse-grained quartz,

biotite and sulfides, like pyrrhotite or molybdenite.

3.4. PREVIOUS WORK

Due to the ongoing active mining, the Arvigo quarry exhibits a large suite of minerals

and the quarry became famous for mineral collecting. Previous work is limited to

fissure mineral descriptions due to the extensive collection possibility (Ruppe 1966;

Simonetti 1971; Graeser and Stalder 1976; Brughera 1984; Wagner, 1968, 1980,

1981, 1983; Weiß and Forster 1997; Wagner et al. 2000a, b). Due to the presence of

epidote, which known to crystallize with low CO2 contents and the knowledge about

CO2 rich fluids in the Lepontine Alps (e.g. Poty et al. 1974; Mullis et al. 1994) fluid

inclusion studies on quartz crystals were made (Wagner et al. 2000a, b; Stalder 2007)

to get information about the evolution of the fluid inclusions.

Fluid inclusion analysis (Wagner et al. 2000a, b; Stalder 2007) shows a distinct

change of the fluid composition with time. Fluid inclusions in the core of quartz

crystals are rich in CO2 (up to 40 vol. % CO2, Stalder 2007), whereas the fluid

inclusions in the rim are CO2 free. Mineral inclusions in the core zone, which

corresponds to the CO2 solution, are hornblende, ilmenite, biotite and carbonate

(Wagner 2000a). Those minerals often found as precursor phases of Ca-Al-silicates of

the mineral. The rim zone contains also mineral inclusions, which change from the

inner part of the rim to the outer part, with the following sequence amianthus, epidote,

chlorite and calcite, which occur over the whole rim zone (Wagner 2000a). Zeolites

instead were not found as inclusions. Homogenization temperatures of fluid

inclusions in the core are up to 365°C (Wagner et al. 2000a, b; Stalder 2007). In

contrast fluid inclusions in the rim yield homogenization temperatures of 160 to


230°C. NaCl content increases from core to rim up to 5.9 wt. % NaCl (Wagner et al.

2000a, b; Stalder 2007).

3.5. SAMPLING AND ANALYTIC METHODS

A suite of representative samples of different veins and fissure was collected in the

field in the years 2006, 2007 and 2008. A subset of the samples was selected for

petrographic, bulk-rock and electron-microprobe studies.

Mineralogical analyses were carried out by point counting of more than 1600

evenly spaced points in each thin section using standard polarized microscope.

Quantitative zeolite analysis were performed at the Institute of Geosciences

(Mineralogy - Geochemistry), University of Freiburg, using a CAMECA SX 100

electron microprobe equipped with five WD spectrometers and one ED detector with

an internal PAP-correction program (Pouchou and Pichior 1991).

Major and minor elements for zeolites were determined at 15 kV accelerating

voltage and 10 nA beam current with a defocused electron beam of 20 µm in diameter

with counting time up to 20 s. Na and K were counted first to minimize the Na and K

loss during determination. Since the zeolite loses water when heated, the crystals were

mounted in epoxy resin to minimize loss of water due to the electron bombardment.

Natural and synthetic standards were used for calibration. The standards employed

were: albite (Na), periclase (Mg), wollastonite (Si), barite (Ba), hematite (Fe),

celestine (Sr), orthoclase (K), anorthite (Ca), rhodonite (Mn), fluorite (F) and rutile

(Ti). The charge balance of zeolites formulas is a reliable measure for the quality of

the analysis and which correlates with the difficulties related to the thermal instability

of zeolites in microprobe analysis. A useful error test investigates the charge balance

between the non-framework cations and the amount of tetrahedral Al (Passaglia

1970). Analyses are considered acceptable if the sum of the charge of the extra-

framework cations (Ca2+, Sr2+, Na+, and K+) is within 10% of the framework charge

(Al3+).


Identification of various minerals was obtained by a BRUKER AXS D8 Advance

X-ray powder diffractometer (XRD) and the DIFFRACplus v5.0 software for

evaluation. Whole rock analyses were performed by standard X-ray fluorescence

(XRF) techniques at the Institute for Geosciences (Mineralogy - Geochemistry) at the

University Freiburg/Germany, using a Philips PW 2404 spectrometer. Pressed powder

and Li-borate fused glass discs were prepared to measure contents of trace and major

elements, respectively. The raw data were processed with the standard XR-55

software of Philips. Relative standard deviations are < 1 % and < 4 % for major and

trace elements, respectively. Loss on ignition was determined gravimetrically by

calculated by heating at 1100 °C for 2 hours.

A slap, 7 mm in thickness, of sample Arvigo 12 was impregnated with

fluorescent epoxy under high pressure conditions at the EMPA (Swiss Federal

Institute for Materials Testing and Research) in Dübendorf, Switzerland to point out

the porosity. Information of the percentage of porosity was conducted by digital

analysis of photomicrographs. Images of the alteration profile Arvigo 12 were

digitized using the software package ImageJ 1.38x (Wayne Rasband, National

Institute of Health, USA) to calculate the area of porosity.

Assemblage stability and phase relationships involving zeolites have been

computed with the computer program Domino/Theriak (de Capitani and Brown 1987)

using the thermodynamic data by Bermann (1988), Evans (1990), Frey et al. (1991)

and Maeder and Bermann (1991). For scolecite and chabazite we adopted the

thermodynamic data of Johnson et al. (1983) and Ogorodova et al. (2002) and the heat

capacity function was predicted from the equation [4Sco + 2Qtz = 3Lmt + An] and [7

Cha + 16 Qtz = 6 Stb + An], respectively.

3.6. RESULTS

3.6.1. Petrography

Fracture related hydrothermal alteration and metasomatism extends a few centimeters

perpendicular out from discrete fractures (Fig. 3.2), Macroscopically the alteration

zone is characterized by a light and porous leaching zone. The fracture or altered wall


rock respectively is associated with precipitation of hydrothermal mineral

paragenesis.

Fig. 3.3: Sample Arvigo 12 that shows a complete section from unaltered host rock to the different alteration zone and fissure precipitation. (a) Hand specimen of the profile. Alteration grade increases in arrow direction. Schistosity is horizontal in respect to the sawed section. (b) Sawed slices showing the position of XRF analysis presented in Table 3.8. (c) Sketch of different alteration areas: unaltered gneiss (G) light altered zone (L), wherein biotite starts, the medium altered zone starts with the first occurrence of turbidity plagioclase, marked by the red dashed line. These zone graduals in the highly altered part (H), wherein chlorite is absent. Secondary phases are precipitated along the fissure wall (V). Additional a thin chlorite in vein in the light altered section is shown. (d) Thin section photographs along the profile. Red dashed line marks the sharp contact from the zone of albitization and the unaltered plagioclase. Albitization produce plagioclase turbidity due to the formation of porosity. Thin section numbers corresponds to Table 3.1.


Table 3.1: Modal mineral distribution through the alteration profile (see Fig. 3.3). mineral TS 12.1 TS 12.2 TS 12.3 plagioclase/albite 49.2 53.8 52.9 quartz 1.5 1.4 2.4 K-feldspar 17.3 13.5 14.0 biotite 0.0 8.7 10.3 muscovite 20.4 17.6 14.5 chlorite 10.4 3.6 4.5 others (apatite, calcite, epidote, titanite ilmenite, laumontite, prehnite)

1.3 1.5 1.5

counted points 2103 2179 1687

3.6.1.1. Unaltered rock

The unaltered gneisses (G, Fig. 3.2) in Arvigo are represented by a suite of biotite-

gneisses and two-mica gneisses, showing distinct differences in their mineralogy. The

dark to dark-colored rocks varying from anhedral/subhedral fine to medium grained

biotite gneisses with an alternation of mafic and felsic layers, representing the

schistosity in a scale of 2 to 4 mm to augen-gneisses with feldspar crystals up to 12

mm in size. The primary minerals are: plagioclase, quartz, K-feldspar, biotite,

muscovite and ±hornblende with the accessories apatite, zircon, rutile, ilmenite and

titanite. The modal composition varies in a wide range from biotite, biotite-

hornblende, biotite-poor, and biotite-muscovite to muscovite dominated gneisses.

3.6.1.2.. Altered rock

The extent of fracture-related hydrothermal alteration ranges in difference stages,

showing differences in altered mineralogy, fissure minerals and porosity. Generally

the altered rocks are marked by an increase in lightness and porosity.

Slightly altered gneisses (L, Fig. 3.3, 3.4) differ from unaltered in the prevalence

of biotite over chlorite. The alteration grade of biotite increases gradually in fissure

direction (Fig. 3.3). In medium altered gneisses (M, Fig. 3.3), almost all biotite is

replaced by chlorite. Microscopically observation shows the development of turbidity

in plagioclase (Figs 3.3, 3.4). In the outer zone (H, Fig. 3.3) all chlorite is dissolved.


3.6.1.3. Fissure minerals

As mentioned above, more than 40 minerals are known, but most of them occur only

infrequently. The Arvigo locality is characterized by the extensive occurrence of

epidote, prehnite, chlorite and zeolites, beside the prevalent fissure minerals quartz,

adularia and calcite (Fig. 3.2). Quartz, adularia represents earlier formed minerals in

the fissure overgrown by Ca-Al silicates, chlorite and calcite, whereas chlorite appear

in two temporal stages. The sequence of crystallization fissures can be determined by

field and microscopy observations as followed: quartz, adularia, chlorite I, epidote,

prehnite, chlorite II, calcite and zeolites. Scolecite and laumontite are by far the

dominant zeolite species, whereas heulandite, chabazite, stilbite and epistilbite can

only be found sporadically.

3.6.1.4. Changes in modal mineralogy

Changes in the modal mineral composition along the alteration profile (Fig. 3.3) are

given in Table 3.1. Plagioclase is the dominant mineral in the gneisses and due to

alteration the modal percentage decreases with alteration. Quartz decreases in the

highly altered zone, suggesting a quartz consuming reaction during alteration.

Considering the potassium bearing phases, an increase of the modal contents of K-

feldspar and muscovite is accompanied by the biotite decrease up to the absence of

biotite in the altered zone (M, Fig. 3.3). This decrease in biotite is again associated

with the increase in chlorite in the rock matrix. Calcite, epidote, prehnite and

laumontite occur as accessory minerals only in TS 12.1 (Fig. 3.3). Titanite and

ilmenite are inversely correlated to the biotite decrease.

3.6.2. Mineralogy and mineral chemistry

To study the change of petrography and chemical composition of primary and

secondary minerals to obtain changes in the mineralogy, mass change and porosity

during fluid-rock interaction sample Arvigo 12 (Fig. 3.3) were taken. This sample was

chosen, because the sample provides unaltered and altered zones in hand-specimen

scale.


Fig. 3.4: Representative microphotographs of typical mineral assemblages in the fresh and altered rocks (Sample Arvigo 12, Fig. 3.3). (a) Unaltered plagioclase crystal showing albite twins. (b) Albitized plagioclase grains. Albite crystals appear to be turbidity due to the porosity increase during albitization. Red dashed line marks the grain boundary between two albite grains. The orientation of pores seems to be depended on the crystallographic orientation. (c) Alteration front between the medium altered and highly altered zone marked by the red dashed line. Plagioclase to the right is unaltered, whereas plagioclase left of the red dashed line are albitized. (d) Typical mineral assemblages not affect by hydrothermal alteration. (e) Biotite alteration to chlorite and K-feldspar in the light altered zone. K-feldspar is etched by H2F and stained by Na3Co(NO2)6 to make it distinguishable from plagioclase. (f) Complete alteration of biotite to chlorite, K-feldspar and ilmenite in the light altered zone.


Table 3.2: Plagioclase composition along the profile (see Fig. 3.3, 3.5, 3.6). x-direction* 783 1482 5402 5876 5702 11756 11766 12977 16722

y-direction* 8989 19263 20361 14628 5452 2314 11678 21050 19899

SiO2 68.72 68.78 69.25 68.86 68.18 68.91 67.98 69.03 68.81

Al2O3 19.50 19.51 19.79 19.69 19.61 19.71 19.43 19.69 19.87

BaO 0.00 0.04 0.00 0.02 0.03 0.00 0.00 0.00 0.00

SrO 0.00 0.00 0.00 0.00 0.00 0.03 0.00 0.04 0.09

CaO 0.14 0.23 0.12 0.24 0.12 0.30 0.12 0.13 0.27

Na2O 11.96 11.92 11.83 11.86 12.02 11.99 11.99 11.99 12.07

K2O 0.03 0.04 0.03 0.05 0.05 0.02 0.04 0.04 0.04

Total+ 100.48 100.60 101.03 100.82 100.04 101.02 99.65 100.98 101.18

Si 2.991 2.991 2.993 2.987 2.982 2.985 2.986 2.990 2.978

Al 1.000 1.000 1.008 1.007 1.011 1.006 1.005 1.005 1.013

Ba 0.000 0.001 0.000 0.000 0.001 0.000 0.000 0.000 0.000

Sr 0.000 0.000 0.000 0.000 0.000 0.001 0.000 0.001 0.002

Ca 0.007 0.010 0.005 0.011 0.006 0.014 0.006 0.006 0.012

Na 1.009 1.005 0.991 0.997 1.020 1.007 1.021 1.006 1.013

K 0.001 0.002 0.001 0.003 0.003 0.001 0.002 0.002 0.002

Sum 5.013 5.013 4.999 5.009 5.023 5.016 5.023 5.012 5.022

% An 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01

% Ab 0.99 0.99 0.99 0.99 0.99 0.99 0.99 0.99 0.99

% Or 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

x-direction* 16780 22422 21787 23741 23543 25733 26417 27112 23778

y-direction* 15244 1917 18330 21711 15929 8818 3569 21789 21319

SiO2 67.93 68.71 68.40 66.67 68.07 68.62 67.34 64.53 66.34

Al2O3 19.78 19.68 19.57 21.28 19.02 19.37 19.53 22.59 20.85

BaO 0.00 0.00 0.00 0.00 0.00 0.00 0.02 0.01 0.00

SrO 0.00 0.04 0.15 0.00 0.00 0.14 0.00 0.16 0.11

CaO 0.47 0.51 0.18 2.20 0.22 0.21 0.53 3.64 1.76

Na2O 11.63 11.69 12.03 10.70 11.70 11.87 11.56 9.95 10.89

K2O 0.07 0.05 0.04 0.20 0.04 0.03 0.07 0.15 0.12

Total+ 99.93 100.74 100.46 101.07 99.22 100.24 99.15 101.09 100.14

Si 2.974 2.984 2.983 2.901 3.000 2.995 2.974 2.824 2.913

Al 1.021 1.007 1.006 1.091 0.988 0.996 1.016 1.165 1.079

Ba 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000

Sr 0.000 0.001 0.004 0.000 0.000 0.003 0.000 0.004 0.003

Ca 0.022 0.024 0.008 0.102 0.010 0.010 0.025 0.171 0.083

Na 0.987 0.984 1.017 0.903 1.000 1.004 0.990 0.844 0.927

K 0.004 0.003 0.002 0.011 0.002 0.001 0.004 0.008 0.007

Sum 5.010 5.005 5.023 5.009 5.007 5.010 5.014 5.019 5.014

% An 0.02 0.02 0.01 0.10 0.01 0.01 0.02 0.17 0.08

% Ab 0.97 0.97 0.99 0.89 0.99 0.99 0.97 0.83 0.91

% Or 0.00 0.00 0.00 0.01 0.00 0.00 0.00 0.01 0.01


continue Table 3.2 x-direction* 27500 29800 33149 32074 36294 36229 39015 46580 49981

y-direction* 16094 15749 3319 10519 18573 18572 3736 6781 19241

SiO2 65.72 65.62 64.28 64.35 4.72 63.43 63.38 63.36 64.12

Al2O3 21.63 21.64 22.65 22.20 3.96 22.56 22.93 22.75 22.54

BaO 0.05 0.02 0.00 0.05 0.01 0.06 0.05 0.00 0.01

SrO 0.00 0.17 0.12 0.07 0.00 0.00 0.00 0.06 0.04

CaO 2.31 2.64 3.48 3.39 0.62 3.85 3.98 3.83 3.29

Na2O 10.59 10.25 9.90 9.64 3.36 9.73 9.51 9.20 9.84

K2O 0.15 0.25 0.33 0.43 0.12 0.18 0.25 0.53 0.26

Total+ 100.47 100.66 100.83 100.19 12.82 99.81 100.20 99.78 100.14

Si 2.880 2.875 2.821 2.839 1.843 2.812 2.800 2.810 2.828

Al 1.117 1.117 1.171 1.154 1.821 1.179 1.194 1.189 1.172

Ba 0.001 0.000 0.000 0.001 0.001 0.001 0.001 0.000 0.000

Sr 0.000 0.004 0.003 0.002 0.000 0.000 0.000 0.001 0.001

Ca 0.108 0.124 0.164 0.160 0.259 0.183 0.188 0.182 0.155

Na 0.900 0.870 0.843 0.825 2.545 0.837 0.814 0.792 0.841

K 0.008 0.014 0.018 0.024 0.061 0.010 0.014 0.030 0.015

∑ 5.015 5.008 5.023 5.008 6.542 5.021 5.016 5.006 5.014

% An 0.11 0.12 0.16 0.16 0.09 0.18 0.19 0.18 0.15

% Ab 0.89 0.86 0.82 0.82 0.89 0.81 0.80 0.79 0.83

% Or 0.01 0.01 0.02 0.02 0.02 0.01 0.01 0.03 0.01

x-direction* 53705 57320 64034 67184 69215 67096 67477 72561 76162

y-direction* 11861 3787 18791 9233 9202 7841 12718 15502 4900

SiO2 62.84 63.96 64.70 64.07 65.83 63.85 64.53 65.24 64.46

Al2O3 22.43 23.13 21.93 22.76 20.67 21.86 22.51 21.91 22.10

BaO 0.00 0.00 0.00 0.01 0.00 0.00 0.02 0.00 0.00

SrO 0.15 0.02 0.00 0.12 0.06 0.11 0.00 0.00 0.19

CaO 4.00 3.94 2.92 3.81 1.62 3.01 2.30 3.04 3.21

Na2O 9.53 9.44 10.11 9.84 10.78 10.20 10.46 10.03 9.88

K2O 0.40 0.26 0.31 0.26 0.24 0.21 0.63 0.29 0.29

Totat+ 99.39 100.78 100.02 100.95 99.23 99.25 100.53 100.57 100.20

Si 2.805 2.805 2.854 2.811 2.916 2.843 2.838 2.861 2.843

Al 1.180 1.196 1.140 1.177 1.079 1.147 1.166 1.132 1.149

Ba 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000

Sr 0.004 0.001 0.000 0.003 0.002 0.003 0.000 0.000 0.005

Ca 0.191 0.185 0.138 0.179 0.077 0.144 0.108 0.143 0.152

Na 0.824 0.802 0.865 0.837 0.926 0.880 0.891 0.853 0.845

K 0.023 0.014 0.018 0.015 0.013 0.012 0.036 0.016 0.016

Sum 5.028 5.005 5.017 5.026 5.014 5.030 5.042 5.007 5.013

% An 0.18 0.18 0.14 0.17 0.08 0.14 0.10 0.14 0.15

% Ab 0.79 0.80 0.85 0.81 0.91 0.85 0.86 0.84 0.83

% Or 0.02 0.01 0.02 0.01 0.01 0.01 0.03 0.02 0.02


continue Table 3.2 x-direction* 87090 93056 72561 105725 99345

y-direction* 4081 17791 15502 9528 17370

SiO2 64.87 63.99 63.51 64.33 63.96

Al2O3 22.10 22.32 22.91 22.22 22.58

BaO 0.00 0.01 0.02 0.02 0.00

SrO 0.09 0.15 0.11 0.10 0.08

CaO 3.08 3.40 3.57 3.13 3.74

Na2O 9.96 9.62 9.64 9.98 9.55

K2O 0.29 0.43 0.28 0.39 0.33

Totat+ 100.39 100.04 100.05 100.22 100.28

Si 2.851 2.830 2.808 2.838 2.821

Al 1.145 1.163 1.194 1.155 1.174

Ba 0.000 0.000 0.000 0.000 0.000

Sr 0.002 0.004 0.003 0.002 0.002

Ca 0.145 0.161 0.169 0.148 0.177

Na 0.848 0.825 0.826 0.854 0.816

K 0.016 0.024 0.016 0.022 0.019

Sum 5.009 5.013 5.016 5.022 5.010

% An 0.14 0.16 0.17 0.14 0.17

% Ab 0.84 0.82 0.82 0.83 0.81

% Or 0.02 0.02 0.02 0.02 0.02

* in µm. + totals include FeO, MgO, MnO, TiO2

Fig. 3.5: Plagioclase composition along the profile Arvigo 12 (Fig. 3.3). Vein mineral precipitation occurs on the left site of the diagram, whereas the host rock continuous to the right. The small spike at around 70000 µm corresponds to the chlorite vein (Fig. 3.3). Chemical analyses are given in Table 3.2.


3.6.2.1. Plagioclase and its alteration products

An important mineralogical difference between fresh and altered rock is the complete

albitization of plagioclase (Table 3.2; Fig. 3.5, 3.6, 3.7). During albitization

plagioclase of oligoclase composition (An15-19) has been replaced by albite (An0.5-2,

Table 3.2; Fig. 3.5, 3.6). Figures 3.5 and 3.6 shows a one-dimensional and two-

dimensional profile, respectively, of plagioclase composition through the alteration

profile (Fig. 3.3). The profile shows a sharp decrease in anorthite component at about

25000 µm in distance to the fissure wall, which corresponds to the appearance of

turbidity in plagioclase (Fig. 3.3, 3.4). The development of turbidity in albite grains in

thin section (Fig. 3.4) is related to porosity increase (Fig. 3.7, 3.8).

Fig. 3.6: Modeled plagioclase composition along the profile Arvigo 12. X-direction is parallel to the schistosity and y-direction perpendicular to them.

The intra-granular pores have an angular to elongated shape, ranges from micrometer

size up to 10 µm in length (Fig. 3.7, 3.8). The slight decrease of anorthite component

at around 70000 µm from the fissure wall is geometrically related to the thin chlorite

vein (Fig. 3.3). The albite component in plagioclase shows an inverse pattern of the

anorthite pattern. Considering the orthoclase content of plagioclase along the profile a

slight decrease in the altered zone is visible (Table 3.2; Fig. 3.3). Additionally a

depletion of Sr in albite can be observed (Table 3.2). Other major changes of


plagioclase geochemistry are not detected. Saussuritization of plagioclase, a process

that is characterized by the replacement of plagioclase by fine-grained sericite (e.g.

Sandström et al. 2008) cannot be observed in our samples. Although the plagioclase is

totally albitized, the original albite law twinning in plagioclase has been preserved in

many of the altered grains (Fig. 3.4).

Fig. 3.7: EMPA images (TS 12.1) showing products of albitization process (Pl + Qtz + H2O = Ab + Lmt). (a) BSE-image showing the porosity (black dots) in albite during albitization of primary plagioclase. Laumontite is easy visible due to the perfect cleavage under an angel of ~90°. Using image analyzing methods a porosity of ~ 15% can be determined. (b) K distribution image. (c) Ca distribution image showing seriate - amoeboid grain boundary of laumontite. (d) Na distribution image. (e) Al distribution image. (f) Si distribution image.

Fig. 3.8: EMPA images (TS 12.1) showing a relict plagioclase grain, surrounding by albite. Considering the porosity enrichment around the plagioclase grain. (a) Ca distribution image. Plagioclase shows a zoning pattern, whereas the Ca content in the rims is higher, than in the core. (b) K distribution image showing seriate - amoeboid grain boundary of laumontite. (c) Na distribution image.


Table 3.3: Representative chlorite, biotite and muscovite composition and forming temperature for chlorite.

sample A8.1 A8.1 A8.2 A10 A10 Arvigo 4

Arvigo 8I

Arvigo 8I

Arvigo 8II

Arvigo 12.1

Arvigo 12.3

no. 8 9 1 21 24 1 13 19 36 14 14 mineral chlorite chlorite chlorite chlorite chlorite chlorite chlorite chlorite chlorite chlorite chlorite SiO2 24.30 24.33 24.38 25.04 25.39 25.86 26.41 25.83 26.79 23.19 24.43 TiO2 0.00 0.00 0.09 0.02 0.02 0.01 0.00 0.04 0.07 0.02 0.09 Al2O3 21.98 21.89 21.48 20.23 20.28 19.52 18.36 19.17 17.82 20.73 20.60 BaO - - - - - 0.07 0.06 0.03 0.00 0.00 0.00 FeO 31.65 32.26 31.30 30.54 30.10 27.10 27.63 25.86 28.48 31.54 30.40 MnO 0.42 0.45 0.46 0.44 0.52 0.31 0.39 0.35 0.23 0.59 0.42 MgO 10.36 10.33 10.61 11.93 11.95 14.29 14.11 15.32 14.22 9.57 9.59 SrO 0.00 0.01 0.03 0.02 0.04 0.00 0.00 0.00 0.03 0.04 0.00 CaO 0.00 0.03 0.06 0.04 0.09 0.04 0.04 0.02 0.06 0.10 0.06 Na2O 0.00 0.02 0.01 0.00 0.00 0.03 0.01 0.00 0.03 0.06 0.00 K2O 0.01 0.00 0.01 0.00 0.01 0.01 0.01 0.01 0.03 0.01 0.00 Total 88.72 89.32 88.44 88.27 88.40 87.25 87.02 86.64 87.76 85.83 85.59 based on 20 oxygens Si 5.244 5.234 5.277 5.408 5.460 5.545 5.695 5.545 5.747 5.220 5.449 AlIV 2.756 2.766 2.723 2.592 2.540 2.455 2.305 2.455 2.253 2.780 2.551 AlVI 2.836 2.783 2.758 2.559 2.600 2.477 2.361 2.395 2.252 2.719 2.862 Ti 0.000 0.000 0.014 0.004 0.003 0.002 0.000 0.006 0.011 0.003 0.015 Fe2+ 5.713 5.803 5.666 5.517 5.413 4.859 4.983 4.642 5.109 5.937 5.670 Mg 3.333 3.311 3.422 3.842 3.830 4.568 4.536 4.904 4.547 3.211 3.187 Mn 0.077 0.082 0.084 0.081 0.094 0.056 0.071 0.064 0.042 0.112 0.080 Ba - - - - - 0.006 0.005 0.003 0.000 0.000 0.000 Sr 0.000 0.001 0.004 0.002 0.005 0.000 0.000 0.000 0.004 0.005 0.000 Ca 0.000 0.006 0.015 0.009 0.021 0.009 0.009 0.006 0.014 0.024 0.014 Na 0.000 0.010 0.005 0.000 0.000 0.012 0.004 0.000 0.012 0.026 0.001 K 0.002 0.000 0.003 0.001 0.003 0.003 0.003 0.002 0.008 0.003 0.001 T (°C)* 382 383 376 355 347 333 309 333 301 386 349 averagea: 333°C ± 32°C

* chlorite temperature is calculated by the empirical calibration of Cathelineau (1988). a average of 39 analysis in 10 samples

sample Arvigo 12.3

Arvigo 12.3

Arvigo 12.3

Arvigo 12.1

Arvigo 12.3

Arvigo 12.3

Arvigo 12.3

Arvigo 12.3

Arvigo 12.1

Arvigo 12.3

no. 5 6 13 10 7 5 6 13 10 7

mineral biotite biotite biotite musco-

vite musco-

vite biotite biotite biotite musco-

vite musco-

vite based on 22 oxygens SiO2 35.31 34.91 34.63 49.47 47.70 Si 5.410 5.383 5.370 6.605 6.416 TiO2 3.06 2.98 2.43 0.52 0.83 AlIV 2.590 2.617 2.630 1.395 1.584 Al2O3 18.26 18.72 18.85 29.96 31.63 AlVI 0.707 0.786 0.815 3.319 3.431 BaO 0.11 0.15 0.10 0.22 0.22 Ti 0.353 0.345 0.283 0.052 0.084 FeO 23.10 22.41 23.36 2.70 2.41 Fe2+ 2.960 2.890 3.030 0.302 0.271 MnO 0.48 0.40 0.42 0.02 0.05 Mg 1.567 1.560 1.497 0.371 0.276 MgO 6.86 6.79 6.48 1.86 1.38 Mn 0.063 0.053 0.055 0.002 0.005 SrO 0.00 0.03 0.00 0.00 0.00 Ba 0.006 0.009 0.006 0.012 0.011 CaO 0.01 0.00 0.02 0.02 0.02 Sr 0.000 0.003 0.000 0.000 0.000 Na2O 0.12 0.12 0.13 0.25 0.34 Ca 0.002 0.000 0.003 0.003 0.003 K2O 9.32 9.19 9.22 10.42 10.02 Na 0.034 0.034 0.038 0.065 0.090 Total 96.63 95.70 95.63 95.45 94.59 K 1.821 1.807 1.823 1.775 1.719


3.6.2.2. Biotite-chlorite

Chloritization of biotite increases gradually from the fresh rock, where biotite is

unaltered, into the light altered zone (Fig. 3.3, 3.4). In the medium and highly altered

zone biotite is totally replaced by chlorite, which also occurs as unconsolidated and

consolidated spherulitic aggregates in fissures. The zone of chloritizated biotite

extends further away from the fracture than is indicated by the macroscopic leaching

zone (Fig. 3.3, red-dashed line), wherein the albitization process is the dominated

alteration feature.

The chloritization reaction of biotite is accompanied by the formation of K-

feldspar and Ti-phases, like ilmenite and titanite (Fig. 3.4). Chloritization appears to

take place preferentially along cleavage planes of biotite. Representative analysis for

biotite and chlorite are given in Table 3.3. Biotite has a FeO content of 21.2 - 24.6

wt% and a MgO content of 6.3 - 7.0 wt%, reflecting values of XMg 0.21 - 0.24.

Chlorite chemistry shows contents of MgO 9.4 - 16.6 wt% and FeO 23.2 - 32.3 wt%

reflecting values of XMg 0.23 - 0.42 that is higher than in biotite. The most evident

change occurs in the concentration of Ti, which is a major element in biotite with an

average value of 3.12 wt% (Table 3.3) and in contrast incorporated only in traces in

chlorite.

3.6.2.3. Muscovite

The muscovite content has been preserved during alteration. Muscovite composition

varies in a limited range, but no systematic pattern over the whole range from

unaltered to altered rock is obvious. The levels of MgO (0.9 - 2.1 wt%) and FeO (1.9 -

2.9 wt%) reflect a celadonite component of 7-10 mol%. The paragonite content is 4 -

7 mol% and significant higher than the margarite component (<1 mol%) (Table 3.3).

3.6.2.4. K-feldspar

The primary K-feldspar has generally been preserved during alteration. Additional K-

feldspar is formed during chloritization of biotite. Chemical composition of rock

forming K-feldspar ranges from Or85-94 and Ab15-06, respectively whereas the anorthite


component is less than 0.5 mol% (Table 3.4). In contrast secondarily formed adularia

as euhedral crystals in fissures, succeeding fissure quartz show a higher Or

component (Or93-99, Ab07-01, An00-01, Table 3.4).

Table 3.4: Representative chemical composition of K-feldspar and adularia. sample Arvigo

12.1 Arvigo

12.1 Arvigo

12.2 Arvigo

12.3 A8.1 A10 A4.1

no. 8 20 26 2 15 18 9 mineral k-feldspar k-feldspar k-feldspar k-feldspar adularia adularia adularia SiO2 64.40 64.67 64.75 64.34 64.97 64.14 64.34 Al2O3 18.98 18.46 18.35 18.47 18.57 18.46 18.39 BaO 0.59 0.56 0.20 0.52 - - - SrO 0.00 0.06 0.07 0.00 0.00 0.00 0.00 CaO 0.05 0.02 0.04 0.05 0.01 0.00 0.01 Na2O 1.59 0.97 0.61 1.20 0.59 0.12 0.36 K2O 14.01 15.05 15.39 14.48 15.78 16.26 16.17 Total* 99.62 99.84 99.82 99.11 99.99 99.06 99.31 based on 8 oxygens Si 2.974 2.992 2.996 2.990 2.996 2.992 2.994 Al 1.033 1.006 1.000 1.011 1.009 1.015 1.009 Ba 0.011 0.010 0.004 0.009 0.000 0.000 0.000 Sr 0.000 0.002 0.002 0.000 0.000 0.000 0.000 Ca 0.002 0.001 0.002 0.002 0.001 0.000 0.000 Na 0.143 0.087 0.055 0.108 0.053 0.011 0.032 K 0.825 0.888 0.908 0.859 0.928 0.968 0.960 Or % 0.85 0.91 0.94 0.89 0.95 0.99 0.97 An % 0.00 0.00 0.00 0.00 0.00 0.00 0.00 Ab % 0.15 0.09 0.06 0.11 0.05 0.01 0.03

* total include FeO, MgO, MnO, TiO2

3.6.2.5. Quartz

Quartz is a rock-forming mineral in the Arvigo gneisses. Considering the model

mineral evolution during alteration, quartz is consumed during alteration (Table 3.1).

However, fissure quartz as the first mineral precipitated in the fissure can give

important information about fluid composition/evolution and mineral evolution

during quartz growth by its fluid and solid inclusions (cf. Previous work).

3.6.2.6. Epidote

Secondarily formed epidote occurs as dominant phase in veins and fissures

overgrowing quartz and adularia (Fig. 3.2). Epidote is overgrown by prehnite and

zeolites. The textural relationship in sample Arvigo 1 (Table 3.5) suggests a co-

genetic growth of epidote and prehnite. Epidote forms green to dark-green sheaf like


aggregates up to 20 mm in length, but common size ranges between 5 and 10 mm.

The chemical composition of epidote (ideal composition: Ca2(Fe3+,Al)3Si3O12(OH))

shows a pistacite component (Fe3+/(Fe3++Al)) that ranges from 15 to 30%. Except the

main constitutes of epidote only MnO and SrO occurs in epidote up to 0.64 wt% and

0.30 wt%, respectively. No mineral zoning in epidote is visible.

Table 3.5: Analysis of coexisting epidote-prehnite (Sample Arvigo 1). sample Arvigo1 Arvigo1 Arvigo1 Arvigo1 Arvigo1 Arvigo1 Arvigo1 Arvigo1 no. 13 14 18 19 16 17 20 21 mineral epidote epidote epidote epidote prehnite prehnite prehnite prehnite SiO2 37.63 37.45 37.35 37.77 42.52 42.44 42.22 42.54 TiO2 0.05 0.05 0.00 0.05 0.09 0.12 0.04 0.02 Al2O3 23.71 23.56 23.25 23.64 21.54 21.20 21.40 21.14 Fe2O3 12.72 12.24 12.64 10.84 3.81 3.98 4.05 3.66 FeO 0.10 - 0.38 - - - - - MnO 0.16 0.12 0.64 0.08 0.05 0.02 0.00 0.03 MgO 0.00 0.01 0.03 0.00 0.00 0.00 0.00 0.00 SrO 0.29 0.29 0.30 0.23 0.00 0.00 0.00 0.00 CaO 23.02 23.40 22.28 23.61 26.58 26.58 26.41 26.78 Na2O 0.02 0.02 0.00 0.00 0.04 0.01 0.00 0.02 K2O 0.00 0.00 0.00 0.01 0.01 0.00 0.00 0.02 Total* 97.70 97.19 96.90 96.24 94.74 94.35 94.12 94.21 based on 8 cations and 12.5 oxygens based on 7 cations and 11 oxygens Si 3.002 3.000 3.011 3.014 2.992 3.000 2.991 3.008 Al 2.229 2.224 2.209 2.223 1.786 1.766 1.787 1.762 Ti 0.003 0.003 0.000 0.003 0.005 0.006 0.002 0.001 Fe3+ 0.764 0.772 0.767 0.743 0.225 0.222 0.227 0.225 Mg 0.000 0.001 0.004 0.000 0.000 0.000 0.000 0.000 Fe2+ 0.007 - 0.025 - - - - - Mn 0.011 0.008 0.044 0.005 0.003 0.001 0.000 0.002 Sr 0.013 0.013 0.014 0.011 0.000 0.000 0.000 0.000 Ca 1.968 2.008 1.925 2.019 2.004 2.013 2.005 2.029 Na 0.003 0.003 0.000 0.000 0.005 0.001 0.000 0.003 K 0.000 0.000 0.000 0.001 0.001 0.000 0.000 0.002

* totals includes traces of BaO

3.6.2.7. Prehnite

Prehnite forms colorless to pale green fan-shaped radiating aggregates, so-called bow-

tie structures (Phillips and Rickwood 1975), or sheaf like aggregates (Fig. 3.9). The

aggregates are up to 10 cm in diameter overgrowing quartz and epidote. Prehnite used

to be a substrate mineral for zeolites. As observed in epidote, prehnite composition is

limited to their major elements (ideal composition: Ca2(Fe3+,Al)2Si3O10(F,OH)2) and

minor elements occur only in traces (Table 3.6). A small content of fluorine due to the

OH ↔ F substitution is detectable (Table 3.6). Significant compositional variations

within prehnite occur at the octahedral site in the Al2O3 - Fe2O3 ratio (Table 3.6; Fig.


3.9). The total iron content (as Fe2O3) varies from 0.0 to 9.5 wt. The content of Fe3+

derby, controls the appearance of prehnite; colorless to pale green prehnite has low

iron contents, whereas dark green prehnite is high in the iron content. Crystal

aggregates often show a zoning pattern in Fe and Al (Table 3.6; Fig. 3.9), suggesting

a Fe3+ ↔ Al substitution during growth, whereas the Fe content decreases to the rim

or with time, respectively.

Table 3.6: Representative prehnite analysis. sample A 10 A 10 A 10 A 10 no. 2 3 4 5 rim core rim core SiO2 43.75 42.95 43.78 43.21 Al2O3 23.33 19.12 23.92 21.20 Fe2O3 0.69 6.59 0.25 3.83 FeO 0.58 0.54 - 0.24 CaO 26.84 26.35 27.32 26.76 Na2O 0.01 0.01 0.02 0.00 K2O 0.01 0.02 0.01 0.00 F 0.05 0.11 0.05 0.01 Total 95.29 95.75 95.27 95.32 -O≡F 0.02 0.05 0.02 0.00 Total* 95.27 95.70 95.24 95.32 based on 7 cations and 11 oxygens Si 3.031 3.031 3.022 3.025 Al 1.905 1.590 1.946 1.749 Fe3+ 0.036 0.350 0.013 0.202 Fe2+ 0.034 0.032 - 0.014 Ca 1.992 1.992 2.021 2.007 Na 0.001 0.001 0.003 0.000 K 0.001 0.002 0.001 0.000 F 0.011 0.025 0.011 0.002

* totals includes traces of MnO, MgO, SrO BaO

Fig. 3.9: EMPA images showing element distribution in a prehnite aggregate indicating a Fe ↔ Al substitution during growth. (a) Fe distribution map showing an iron-enrichment in the core, which decrease to the rim. (b) Al distribution map showing an alumina-depletion in the core.


3.6.2.8. Zeolites

Scolecite and laumontite are by far the dominant zeolite species, whereas heulandite,

chabazite, stilbite and epistilbite can only be found sporadically. Samples from

Arvigo show the zeolite succession laumontite after scolecite (Fig. 3.2). However, one

sample indicates an inverse growth pattern in which laumontite is the first zeolite that

has formed. However from succession following chronology with increasing age can

be compiled for Alpine fissures: scolecite, laumontite, heulandite, chabazite and

stilbite (Weisenberger and Bucher 2009).

Laumontite (Ca4(Al8Si16O48) •18 H2O) often forms radiating aggregates. It forms

thin, elongated fibers or prisms elongated along the c-axis with a square cross-section.

Laumontite forms as {110} prism and commonly be twinned on {100} to form

“swallow tail” or “V” twins. It is white with a length between <1 to 15 mm. The

cleavage is perfect, with a slight pearly luster on the broad cleavage surface.

Considering the chemical composition calcium is the dominant extra-framework

cation (average value of 97 %), with minor amounts of sodium and potassium (Fig.

3.10, Table 3.7), which consists less than 5 % of the extra-framework cations.

Maximum value for K and Na are 7 and 1 %, respectively. Other elements occur only

in traces. The extra-framework cation K increases during growth from core to the rim

(Fig. 3.10). The Si/(Si+Al) content varies slightly between 0.67 and 0.69 (Fig 3.11). It

can be observed that the content of alkalies increases with increasing Si and

decreasing Ca and Al during growth, which can be expressed by the coupled

substitution Si4+ + (Na+, K+) ↔ Al3+ + Ca2+.

Scolecite (Ca8(Al16Si24O80) •24 H2O) occur as white fibrous crystals with vitreous

or slightly silky luster, forming characteristic radiating sprays (Fig. 3.2). The

transparent to translucent crystals are slender prismatic with square cross section.

Scolecite is between 1 and 20 mm in length, but the common length range is between

3 and 6 mm. 23 microprobe analyses from 3 different samples indicate no major

chemical changes (Fig. 3.10, 3.11). Calcium is the dominant extra-framework cation

(Fig 3.10, Table 3.7), which in average occupies 98 % of the extra-framework cation

sites. Minor amounts of Na up to 4 %, but in average 2 % can be observed. In contrast

K and Sr are not significantly incorporated in the framework structure of scolecite

(Fig 3.10, 3.11) and therefore occur only in traces, like Ba, Mg and Fe (Table 3.7).


The small substitution range can also be recognized in the Si/(Si+Al) ratio, which

shows a small range between 0.61 and 0.62 (Fig. 3.11).

Fig. 3.10: Triangular plots of extra-framework cation (Ca+Sr+Mg-Na-K) distribution of zeolites. Dashed framed area marks the chemical composition of zeolite found in granites and gneisses in the Swiss Alps (Weisenberger and Bucher 2009).

Heulandite ((Na,K)Ca4(Al9Si27O72) •24 H2O) forms crystals up to 12 mm in length,

but the average size of the crystals is 1 to 4 mm. The monoclinic crystals occur in

tabular habit parallel {010} and elongated in its typical coffin-shaped appearance.

Crystals are transparent to translucent and colorless. Representative analyses for

heulandite are given in Table 3.7. The average composition of heulandite from Arvigo

is Ca3.27Na0.19K1.45Sr0.30(Al8.84Si27.16O72) •22 H2O, which is nearly identical to the

heulandite composition (Ca3.37Na0.07K0.88Sr0.55(Al8.42Si27.49O72) •22 H2O) determined

by Armbruster et al. (1996) from Gibelsbach in the western Aar Massif and from

other Alpine fissures (Weisenberger and Bucher 2009). Heulandite can be classified

by its geochemistry as heulandite-Ca (Coombs et al. 1998), with higher amounts of Sr

and K (Fig. 3.10, Table 3.7). Calcium, with an average value of 63 % of all extra-

framework cations, is the important extra-framework cation. Beside Ca, K marks a

major element in heulandite, which yields an average value of 28 % (maximum to 31

%). Mentionable are the significant Sr content ranges between 5 and 7 %, which can

be incorporated on the extra-framework cation sides and distinguish heulandite from

the other zeolites found in Arvigo and in Alpine fissure, except chabazite, which

comparably shows an enrichment of Sr (Weisenberger and Bucher 2009). The

Si/(Si+Al) content ranges between 0.74 and 0.76 (Fig.11), which is in agreement with


the definition of heulandite (Coombs et al. 1998), that can be distinguished from

clinoptilolite, 0.8 < Si/(Si+Al).

Fig. 3.11: R2+ - R+ - Si compositional diagram of zeolites. Si/Al ratio increases with chronologic order. Dashed framed area marks the chemical composition of zeolite found in granites and gneisses in the Swiss Alps (Weisenberger and Bucher 2009).

Table 3.7: Representative analysis of zeolite species. sample Arvigo12I Arvigo12I

I Arvigo2 Arvigo1 Arvigo 13 Arvigo 13 A8.1 A8.1

no. 7 6 4 6 3 11 13 14 mineral laumontite laumontite laumontite scolecite scolecite scolecite heulandite heulandite

SiO2 52.50 51.43 51.52 45.64 45.64 45.82 57.84 59.46 Al2O3 21.88 19.97 21.24 24.65 24.72 25.33 16.44 15.79 CaO 11.42 10.90 11.86 14.14 13.55 13.97 6.41 6.29 SrO 0.11 0.00 0.04 0.00 0.00 0.04 1.40 1.14 BaO 0.01 0.00 0.02 0.00 0.00 0.01 Na2O 0.04 0.01 0.04 0.10 0.04 0.04 0.28 0.15 K2O 0.36 0.70 0.16 0.00 0.01 0.01 2.90 2.32 H2O 13.37 - - - - 13.69 - - Total 99.73 83.08 84.88 84.62 83.98 98.98 85.40 85.26 formula unit composition Si 16.124 16.415 16.112 24.312 24.413 24.187 26.921 27.449 Al 7.920 7.512 7.830 15.476 15.584 15.761 9.018 8.592 Ca 3.758 3.727 3.975 8.070 7.764 7.901 3.196 3.112 Sr 0.020 0.000 0.007 0.000 0.000 0.013 0.376 0.305 Ba 0.001 0.000 0.002 0.000 0.000 0.003 - - Na 0.000 0.006 0.024 0.103 0.039 0.039 0.256 0.134 K 0.141 0.285 0.064 0.000 0.004 0.005 1.723 1.366 O 48 48 48 80 80 80 72 72 H2O 13.696 - - - - 24.113 - - TSi 0.669 0.686 0.673 0.611 0.610 0.605 0.749 0.762 E% 0.06 -3.14 -2.86 -4.73 0.00 -0.73 -1.58 3.09

* totals includes traces of FeO, MgO, MnOTiO2. TSi – Si/(Si+Al) E% - a measure of charge balance, = (100*((Al)-(Na+K)+2(Mg+Ca+Sr+Ba)/(Na+K)+2(Mg+Ca+Sr+Ba))


In general all zeolite in Arvigo, as well as in fissure hosted in granites and gneisses in

the Central Alps (Weisenberger and Bucher 2009), are Ca-dominated and the Si/Al

ratio increases with decreasing temperature/time (Fig. 3.11).

3.6.3. Porosity

Porosity increase during fracture related alteration is a multiple process, by the

volume increase changes during albitization process (Eq. 1, 2) and the removal of

primary and secondary phases, like chlorite, which where formed during biotite

dissolution. The porosity of gneiss varies between 1.0 and of 1.9 %. The porosity in

the altered rock zone varies due to the removal of phases between 3.8 % at the

alteration front of albitization and increases continuously up to 6.2 % in the medium

altered rock (M, Fig. 3.3). The highly altered zone (H, Fig. 3.3), which is

characterized by the removal of chlorite exhibit a porosity of up to 14.2 %. The

porosity, which generated during albitization, is in submicroscopic scale and often not

connected. This results in a non-considering of porosity during impregnation and

digital analysis methods, which gives uncertainties of up to approximately 15 %.

Considering the process of albitization a volume change of ~16 % (Eq. 1, 2) can be

calculated in albite by using molar volume, which is in acceptable agreement with

porosity determination using image analysis (Fig. 3.7) resulting in an value of 12.7 %.

1 oligoclase + H2O => 4.19 albite + 0.81 Ca2+ + 1.62 AlO2- + 1,62 SiO2,aq + H2O (1)

∆Vsolids = (4.19 VAb – 1 VOlg)/(1 VOlg) (2)

whereas following mineral composition were used:

oligoclase = Na4.15Ca0.85Al5,85Si14.15O40

albite = Na0.99Ca0.01Al1.01Si2.99O8

Using the volume of plagioclase (Table 3.1) a porosity increase of 8.5 % can be

related to the albitization process on whole rock scale. Figure 8 shows a relictic


plagioclase grain that did not completely replaced by albite and can give therefore an

insight into the porosity evolution during replacement of a grain. Thereby the porosity

in albite can be detected with 5 % whereas around the relictic plagioclase grain the

porosity is enriched and increase up to 34 % (Fig. 3.8).

3.6.4. Whole rock geochemistry and mass changes

Changes of mass are common during hydrothermal alteration. A method for mass-

balance analyses was described by Gresens (1967), which is based on the assumption

that some elements are immobile and therefore conserved during the alteration. The

ratio of mobile elements in the fresh and altered rock is then compared to the ratio of

the immobile elements, in order to calculate the mass or volume change during

alteration.

Table 3.8: Major and trace element composition through the alteration profile, including density (see Fig. 3.3).

sample Arvigo 12 Arvigo 12 Arvigo 12 Arvigo 12 Arvigo 12 Arvigo 12 Arvigo 12 Arvigo 12 no. A I* A II A III A IV A V A VI A VII A VIII wt. % wt. % wt. % wt. % wt. % wt. % wt. % wt. % SiO2 57.10 58.04 56.49 56.28 56.40 56.36 56.68 56.02 TiO2 0.66 0.62 0.57 0.57 0.57 0.55 0.54 0.58 Al2O3 21.65 21.29 22.67 22.94 22.94 22.42 22.90 22.96 Fe2O3

tot 3.74 3.96 3.71 3.75 3.77 3.96 3.70 3.93 MnO 0.06 0.07 0.06 0.06 0.06 0.07 0.07 0.07 MgO 1.05 1.14 1.08 1.08 1.07 1.07 1.03 1.11 CaO 1.93 1.04 2.69 2.68 2.73 2.75 2.70 2.67 Na2O 5.37 6.22 5.68 5.75 5.84 5.93 5.91 5.82 K2O 4.71 4.43 4.13 4.25 4.05 3.83 3.98 4.26 P2O5 0.36 0.32 0.32 0.29 0.28 0.29 0.25 0.31 L.O.I. 2.44 1.66 1.45 1.20 1.27 1.34 1.31 1.06 Totals 99.21 98.95 99.01 99.00 99.12 98.72 99.21 98.94 ppm ppm ppm ppm ppm ppm ppm ppm V 59 53 51 50 52 52 53 57 Cr 31 25 26 24 25 27 25 27 Ni 18 15 15 19 13 15 15 17 Cu 10 < 5 < 5 8 < 5 1 5 15 Zn 54 55 54 55 53 54 52 58 Rb 144 126 122 141 132 124 129 152 Sr 105 118 291 298 304 313 307 302 Zr 323 323 280 265 275 266 264 276 Ba 801 773 701 665 641 610 633 638

(g/cm3) (g/cm3) (g/cm3) (g/cm3) (g/cm3) (g/cm3) (g/cm3) (g/cm3)

density 2.884 2.797 2.849 2.848 2.850 2.825 2.832 2.874 *includes fissure minerals


Bulk rock chemistry and density measurements along the alteration profile (Fig. 3.3)

are presented in Table 3.8. Using the simplified graphical method after Grant (1986)

to solve Gresens’ (1967) equation, and the assumption that TiO2 is conservative

during the hydrothermal alteration, Fig. 3.12 represents the isocon diagram. Anyway,

titanite as Ti-bearing phase is rarely found as fissure minerals (Weiß and Forster

1997; Wagner et al. 2000a, b), suggesting a slight Ti mobility.

Fig. 3.12: Isocon diagram and histogram for the chemical loss and gain during alteration. Fresh rock is based on sample Arvigo 12 AVIII, whereas the altered rock composition is based on analyses Arvigo 12 AII (Fig. 3.3, Table 3.8). Isocon diagram showing constant mass (CM), constant volume (CV) and Isocon line. Elements below the lines are depleted in the altered rock relative to the fresh rock. Co = concentration of original element; Cf = concentration of transformed element. Histogram showing oxide mass changes compared to their respective mass in the fresh rock. Mfi = weight concentration of component i in transformed rock; Moi = weight concentration of component i in original rock; Mo = mass of the original rock; (Mfi-Moi)/Moi = mass change in relation to original element mass; (Mfi-Moi)/Mo = mass change in relation to original rock mass. Diagrams were constracted by using the program GEOISO (Coelho 2006)

Elements plotted above the isocon have been enriched relative to the fresh rock,

whereas elements below the isocon line have been depleted during the alteration

process. The slope of the obtained isocon is 1.068 (Fig. 3.12), equivalent to a mass

loss of 6.8% (Grant 1986). Changes of the rock volume during alteration can be

calculated using the mass ratio of immobile elements and the rock densities of the

fresh and altered rock. Using the density and the mass ratio of immobile elements of

0.932 (inverted slope of the isocon), a volume loss of 3.9 % is obtained. Isocons

representing constant mass and constant volume instead of constant TiO2 are included

in Fig. 3.12 for comparison.

Considering gain and loss during alteration, CaO, Sr and Rb are the elements that

shows the highest grade of depletion in respect to their element mass (Fig. 3.12) and


decreased by 64 %, 63 % and 25 %, respectively. K2O, SiO2, MgO, Fe2O3 MnO and

Al2O3 are depleted but in minor amounts, 3 %, 3 %, 4 %, 6 %, 7 % and 13 %,

respectively. By defining TiO2 as immobile during the alteration process, Na2O

behaves also conservative, without any changes (Fig. 3.12). Ba, Zr and P2O5 increase

during alteration, 13 %, 10 % and 9 %, respectively. Nevertheless Al2O3, SiO2, CaO,

Fe2O3 and K2O are the significant elements, relative to the rock mass, whereas the

other element changes are not significant, because they occur only in traces (Fig.

3.12). Given a mass loss of 6.8 % and the changes in major elements, the mass-

balance equation for the hydrothermal alteration of the rock is (Eq. 3):

100 g fresh rock + fluid => 93.2 g of altered rock + 3.0 g Al2O3 + 1.7 g SiO2 + 1.7 g

CaO + 0.3 g Fe2O3 + 0.1 g K2O (3)

3.7. DISCUSSION

3.7.1. Mineral reactions

The most apparent mineralogical changes in the altered rock are the albitization (Fig.

3.3, 3.4, 3.7, 3.8) and the chloritization of biotite (Fig. 3.3, 3.4). Chloritization of

biotite extents farther away from the fracture than is indicated by the brightening of

the rock due to the albitization process (Fig. 3.3). This could be happens either

because biotite is more easily altered than plagioclase, or due to fluids that are more

easily transported along the connected sheet silicate clusters.

The proposed biotite chloritization reaction (Eq. 4, 5) is based on the assumption

of conserved Al and Ti (Ferry 1979; Tulloch 1979; Parry and Downey 1982) and

caused in observed mineral changes in samples from Arvigo, as well as on average

biotite and chlorite composition. By reason that the Fe3+/Fetotal ration in biotite and

chlorite is unknown, all Fe has been assumed to be Fe2+.

1.8 biotite + 4.6 H2O + 1.3 Mg2+ + 0.6 Ca2+ + 0.4 H+ => 1.0 chlorite + 1.1 K-feldspar

+ 0.6 titanite + 0.2 Fe2+ + 0.2 SiO2 + 2.1 K+ (4)


2.05 biotite + 3.6 H2O + 0.9 Mg2+ + 0.5 SiO2 => 1.0 chlorite + 2.0 K-feldspar + 0.8

ilmenite + 0.1 Fe2+ + 1.7 K+ + 0.6 H+ (5)

whereas following mineral composition were used:

biotite = K1.8(Fe3.0Mg1.6)(Al0.8Ti0.4)(Si5.4Al2.6O20)(OH)4

chlorite = (Mg4.2Fe5.2Al2.6)(Si5.6Al2.4O20)(OH)16

adularia = KAlSi3O8

titanite = CaTiSiO5

ilmenite = FeTiO3

This solid-solid reaction is confirmed by petrographic observations (Fig. 3.4).

However the reactions yields a volume increase of 6 % and 8 %, respectively, using

the mineral molar volume of Eq. 4 and 5. Textural observation suggests that the

chloritization process is volume conservative and therefore elements have to be

transported away to achieve the fully pseudomorphic replacement (Fig. 3.4). Whether

the chlorite or the K-feldspar component is dissolved cannot be achieved due to the

fact that the stoichiometric coefficient of the products and reactants varies in the rock

sample. Therefore the iso-volume reactions 6 and 7 display two endmember versions

whereas chlorite and K-feldspar, respectively is dissolved to achieve iso-volume

conditions.

1 biotite + 0.22 Mg2+ + 3.6 H+ => 0.43 chlorite + 0.98 K-feldspar + 0.39 ilmenite +

0.35 Fe2+ + 0.82 K+ + 0.04 SiO2 + 0.25 Al3+ + 0.35 H2O (6)

1 biotite + 0.11 H2O + 0.45 Mg2+ + 3.58 H+ => 0.49 chlorite + 0.77 K-feldspar + 0.39

ilmenite + 0.07 Fe2+ + 1.03 K+ + 0.36 SiO2 + 0.19 Al3+ (7)


Considering the chloritization reaction of biotite (Eq. 6, 7), the reaction takes place in

the presence of fluid (H2O) and Mg2+ and whereas SiO2, K+, Fe2+ and Al3+ and are

released.

The albitization process (Eq. 1) is a common equilibrium process under fluid

saturated conditions over a wide PT range from diagenesis (Saigal et al. 1988, Lee et

al. 2003) to greenschist (Leichmann et al. 2003) and even amphibolite facies

metamorphism (Clark et al. 2005). During the albitization of plagioclase, dissolution

of oligoclase occurs with coeval formation of albite. Ca2+, Al3+, SiO2, which were

released during the process (Eq. 1), are transported in solution to the fracture or to

open space in the adjacent rock (Fig. 3.7), where these elements precipitate as Ca-Al-

silicates. Due to the limited solubility of Al3+, it seems likely that Al3+ does not

migrate over significant distances and precipitates in proximate parts.

Oligoclase (An15-19) from Arvigo samples (Table 3.2, Fig. 3.5, 3.6) has been

replaced by albite (An0.5-2). The product of the reaction that pseudomorphically

replaced plagioclase crystals by albite produce porosity due to the differences in

molar volume between the solid phases (Eq. 2). However, microporosity in

plagioclase/albite increases with alteration as seen in Fig. 3.7 and 3.8, whereas the

orientation of the pores are related to the crystallographic orientation (Fig. 3.4).

Preserved albite twinning in albitized plagioclase can be seen in Fig. 3.4, which

implies an epitactic overgrowth as previously described in altered plagioclase (Engvik

et al. 2008) as well as in K-feldspar (Walker et al. 1995; Cole et al. 2004). This

implies that the albitization process is controlled by dissolution-reprecipitation

mechanism along a moving interface (e.g. Putnis and Putnis 2007, Engvik et al.

2008), resulting in porosity generation.

Primary muscovite and K-feldspar does not show any alteration texture without a

signs of alteration and therefore they have not been regarded in the alteration scheme

in Fig. 3.13.

3.7.2. Mass changes and element mobility

Mass changes and element mobility on whole rock scale is limited to few elements

(Fig. 3.12). Significant loss is marked by the major elements CaO, Al2O3, SiO2,


Fe2O3, MgO and K2O (Fig. 3.12). In contrast the loss of trace elements Rb and Sr and

gain of Ba, Zr and P2O5 is relatively high, but their quantitative changes are

evanescent in respect to the whole rock mass (Fig. 3.12). Mass changes can basically

be linked to the two major alteration reactions of biotite chloritization and albitization

of plagioclase (Eq. 1, 6, 7). Element mobility is summarized in Fig. 3.13 representing

an alteration scheme of redistribution of elements during alteration between primary

minerals, secondary minerals and hydrothermal fluid.

Fig. 3.13: Flow chart illustrating the exchange of ions and element mobility during hydrothermal alteration of gneisses from Arvigo. The diagram refers to alteration in rocks containing plagioclase altered to albite and chloritization of biotite. An external fluid is required to supply of H2O, CO2 and O2 for alteration. Primary Muscovite and K-feldspar regarded in the alteration scheme, due to the fact that they are not involved into the alteration. Polygons in the wall rock field represent primary minerals (Pl, Bt and Qtz) and their secondary alteration products (Ab, Kfs, Chl, Ttn and Ilm) that remain into the wall rock during alteration. Secondary minerals plotted into the fracture field, that are found as euhedral fissure minerals in the Arvigo fissures, but it is not excluded that these secondary minerals are not precipitated in open space in the wall rock. Black arrows represent migration of elements, which are based on alteration reaction and average mineral composition discussed in the text. Dashed black arrows represents element migration, that are limited onto framework conditions: chlorite and K-feldspar formation in the fissure will happen, if the volume increase during chloritization of biotite can not balanced by the precipitation in open spaces in the adjacent to the wall rock. SiO2 required for zeolite formation, may either derive from primary quartz, or is provided in solution in the hydrothermal fluid. Prehnite and epidote will incorporates iron, if Fe2+ is oxidized to Fe3+.

The decrease of Ca2+ is compatible with albitization of plagioclase (Eq. 1), during

which Ca2+ is mobilized (Fig. 3.12, Table 3.8). The change in Ca2+ is strongly

connected to Sr2+, due to similarity in size and charge that allows Sr2+ to substitute


Ca2+ in the plagioclase lattice (Sun et al. 1974). Therefore it is obvious that Sr2+ is

also mobilized during plagioclase dissolution. Most of the Ca2+ content was

transported out in solution of the wall rock (Fig. 3.12, Table 3.8) and precipitated in

the fracture as Ca-Al-silicates (epidote, prehnite and zeolites) and calcite, depending

on fluid composition (CO2-H2O) and temperature. Nevertheless around 1 wt% CaO is

still stored in the altered rock. The remaining content of Ca2+, which is incorporated

in albite only in traces (Table 3.2) is assumed to be hosted in titanite formed during

chloritization (Eq. 6) and in secondary phases, which precipitated already in the open

space of the wall rock (Fig. 3.7). Sr2+ is preferred to integrate into heulandite (Table

3.7).

The decrease of Al (Fig. 3.12) is noteably high in fact that Al is generally

relatively immobile compared to other elements during fluid-rock interaction

(Carmichael 1969; Ragnarsdottir and Walther 1985; Verdes et al. 1992). Al often has

been assumed to be immobile and was therefore often used as a constant reference

frame for mass balance calculations (e.g. Thompson 1975; Grant 1986). Nevertheless,

field evidences (Fig. 3.3), mass changes (Fig. 3.12, Table 3.8) and mineral reaction

(Eq. 1) suggest Al3+ mobility during fluid-rock alteration. Considering the albitization

process, by which porosity is generated, Al3+ leached out to the fracture and

precipitated as Ca-Al-silicate. The volume conservative chloritization reaction marks

an additional source for the Al3+ deficiency and Al mobility. Therefore, Al3+ have to

be transported away (Eq. 6, 7) to achieve the fully pseudomorphic replacement (Fig.

3.4), without volume expansion.

The decrease of SiO2 during alteration is linked with the leaching of silica during

chloritization and albitization (Eq. 1, 6, 7).

Considering Eq. 6 and 7, Fe is released during chloritization and Mg has to be

added to balance the chloritization reaction. Nevertheless mass balance calculation

indicating a loss in both elements (Fig. 3.12). The Mg loss is related to the chlorite

removal out of the wall rock into the fracture (Fig. 3.3, Table 3.1). Fe (Eq. 6, 7)

migrates to the fracture, where it is oxidized and precipitates in epidote and prehnite

(Table 3.5, 3.6).

The volume conservative chloritization (Eq. 6, 7) consume Mg2+. Mg could be

added by an external fluid or due to migration of biotite from the adjacent rock.


However, previous research (Parry and Downey 1982; Parneix et al. 1985; Drake et

al. 2008) has shown that the Mg content of chlorite decreases systematically as more

biotite is replaced, which indicates that Mg is a very mobile element during biotite

chloritization.

The decrease in K+ is consistent with the chloritization process (Eq. 6, 7) during

which K+ leached out from the wall rock (Fig. 3.2). Primary K-feldspar and

muscovite can be excluded as K source by textural evidences (Fig. 3.4).

The increase of volatiles (LOI, Table 3.8) is not astonishing, if we consider the

hydration reaction of biotite and the formation of hydrous Ca-Al-silicates during

albitization.

No changes are assessed in Na+ concentration (Fig. 3.12) during alteration. These

agree with volume calculations during albitization and measured microporosity in

altered plagioclase, regarding the chemical change in plagioclase.

Elements that were release during albitization and chloritization are transported

out to the fissure and precipitates their in secondary minerals. In general solute

transport in porous material is accomplished by three principal mechanism: advection,

aqueous diffusion and hydrodynamic dispersion, whereas advection is the dominant

mechanism (Steefel 2008). Therefore a temperature gradient between the fissure-fluid

and the fluid that reacts with the minerals represents a reliable driving force that

enhance advection of the solution into fissure direction.

3.7.3. Mineral stability and mineral equilibria

3.7.3.1. Prehnite and epidote

Prehnite and epidote are common minerals in low-grade metamorphic rocks (e.g.

Fricke 1952; Kuniyoshi and Liou 1976; Tulloch 1979; Liou et al. 1983; Liou 1985;

Cho et al. 1986; Rose and Bird 1987; Bevins et al. 1991; Freiberger et al. 2001).

Temperature and/or fO2 conditions, during which prehnite and epidote are formed, are

reflected in the chemical composition. The occurrence of pumpellyite, which is often

associated with prehnite and epidote, could not be confirmed and therefore can give

an indication about pressure conditions (Kuniyoshi and Liou 1976). The quite


common inhomogeneity of the Al2O3-Fe2O3 ratio, which is common on thin section

scale in prehnite, may be due to (1) partial re-equilibration during progressively

changing P-T-fO2 conditions in process of Ca-Al-silicates formation, whereas the iron

content in prehnite increases with decreasing temperature and increasing fO2, whereas

the Al contents decrease (Kuniyoshi and Liou 1976; Liou et al. 1983), (2) successive

discrete hydrothermal events (Freiberger et al. 2001), or (3) local chemical influence

of host minerals. However elevated fO2 conditions necessary for the oxidizing process

of Fe2+, released during biotite dissolution, seems like the cause for the Fe3+

enrichment in the core (Fig. 3.9).

Fig. 3.14: Predicted temperature of the formation of coexisting prehnite and epidote, using the Fe3+ - Al partitioning curves, determined by Rose and Bird (1987). Distribution of Fe3+ between coexisting epidote and prehnite is expressed as the mole fraction of Ca2FeAlSi3O10(OH)2 in prehnite and as pistacite component Ca2Fe3Si3O12(OH) in epidote. Dashed lines represents limits on the compositional range for coexisting prehnite and epidote (I: prehnite => zoisite + grossular + quartz; II: laumontite + prehnite => clinozoisite + quartz). Solid lines presents isotherms, based on constant log K by using thermodynamic properties of the reaction: Al-prehnite + epidote => Fe-prehnite + clinozoisite (adapted from Rose and Bird 1987).

Prehnite stability determined for metabasites reaches up to 400 °C and up to 300 MPa

(Liou et al. 1985; Frey et al. 1991). However the absence of pumpellyite, which is

stable between 100 and 800 MPa, suggests that pressure conditions during the


formation of Ca-Al-silicates were below 100 MPa (Kuniyoshi and Liou 1976; Liou et

al. 1985; Frey et al. 1991).

Sample Arvigo 1 (Table 3.5) reveals the coexisting prehnite-epidote assemblage

and therefore it can be used to get information about the formation temperature.

Figure 3.14 shows the formation temperature of coexisting prehnite and epidote, using

the approach from Rose and Bird (1987). Rose and Bird (1987) suggested that the

iron partitioning of coexisting prehnite and epidote is a function of temperature.

Using the iron partitioning treatment after Rose and Bird (1987) a formation

temperature between 330 and 380°C (Fig. 3.14) for coexisting prehnite and epidote

can be diagnosed. Isotherms in Fig. 3.14 are based on thermodynamic calculations by

iterative solutions for the composition of coexisting prehnite, epidote and grandite

garnet (Rose and Bird 1987). The compositional limits on the stability of coexisting

prehnite and epidote are represented by the two dashed lines (Fig. 3.14). Data from

Arvigo (Fig. 3.14) requires that the coexisting prehnite and epidote pairs represent

non-equilibrium Fe3+-Al partitioning and are metastable with respect to the reaction:

Prh = Zo + Grs + Qtz. The evaluated temperature is consistent for prehnite stability in

active geothermal systems (275-350°C; Bird et al. 1984) and in hydrothermal

experiments (376°C, Liou et al. 1983).

Considering the determined temperature and formation temperature of chlorite,

which are in the same range and therefore well agree with the textural appearances

that suggests a contemporaneous growth.

3.7.3.2. Chlorite

Several chlorite thermometers, applying structural and chemical criteria are available

from the literature (e.g. De Caritat et al. 1993). The empirical calibration based on

AlIV content (Cathelineau and Nieva 1985; Cathelineau 1988) was tested for low-

grade basic rocks within a regional metamorphic context (Bevins et al. 1991).

However De Caritat et al. (1993) has shown that the content of AlIV is not dependent

on the geochemical composition of the host rock and therefore the Cathelineau (1988)

thermometer is frequently used including chlorite hosted in granites and gneisses (e.g.

Rahn et al. 1994; Orvosová et al. 1998). Although the precision of such empirical


thermometers is difficult to assess, Vidal et al. (2001) showed that the variation of

AlIV in chlorite with is temperature thermodynamically sound.

Chemical composition of chlorite in Arvigo was examined on 40 grains and the

corresponding formation temperature varies from 27 to 380 °C in a wide range, with

an average value of 333 °C (Table 3.3). Considering the temperature distribution (Fig.

3.15), two distinct groups of chlorite are evident. The first group shows a formation

temperature around 310 °C and the higher temperature group varies from 330 to 380

°C. Chlorite, which pseudomorphic replaces biotite in the rock matrix, trends to

higher formation temperatures, in contrast to spherulitic chlorite precipitated in the

fissure and open space, respectively, which is related to lower temperatures.

Fig. 3.15: Distribution of calculated chlorite temperatures for Arvigo samples using the calibration after Cathelineau (1988).

3.7.3.3. Zeolites

Zeolites mark beside apophyllite, the youngest secondary mineral formed in the

Alpine fissure in Arvigo. The general chronology of the Arvigo zeolite assemblages

is: scolecite, laumontite, heulandite and stilbite and is comparable with recent

evaluations of zeolite bearing fissures in the Alps and thermodynamic phase modeling

in the system CaAl2Si2O8–SiO2–H2O (Weisenberger and Bucher 2009).

Reaction isograds for Ca-zeolites are well determined by experimental methods

(e.g. Liou 1971; Thompson 1970; Cho et al. 1987; Frey et al. 1991). In general the

maximum temperature and pressure limits of zeolite stability are in agreement with

observations on geothermal systems (Kristmannsdóttir and Tómasson 1978; Frey et

al. 1991). Nevertheless, there is a significant deviation between temperature noted at

the position of a given zeolite isograd reaction and temperature resulted from phase


equilibrium calculations (Kristmannsdóttir and Tómasson 1978; Frey et al. 1991;

Neuhoff et al. 2000), which arises difficulties in attempt to use experimental

observations of phase equilibrium to assess thermobarometric conditions in zeolite-

facies rocks that usually reflects higher temperatures as it observed in natural systems.

According to this discrepancy various variables like pH, chemical composition of the

water, pCO2, the presence of additional extra-framework cations like Sr, Na and K, the

amount of H2O incorporated in the zeolite channel structure, order-disorder and fluid

pressure, respectively, can affect the thermodynamic equilibrium conditions and

consequently the reaction isograds (e.g. Thompson 1970; Liou 1971; Cho et al. 1987;

Frey et al. 1991; Neuhoff et al. 1999).

Fig. 3.16: Temperature - fO2 phase-diagram (50 MPa) displaying the stable mineral assemblages for the bulk composition of host rock material of profile Arvigo 12. Thermodynamic calculation where done with iron-free prehnite.

The mineral evolution and the evolution of porosity were modeled by using computed

assemblages stability diagrams with the Theriak/Domino software of de Capitani and

Brown (1987). Considering an increase in fO2 during alteration, which is implied by

the iron zoning in prehnite (Fig. 3.9), Figure 3.16 represents a T-fO2 phase diagram. It


indicates that at constant temperature a change of fO2, which is externally controlled

by the infiltrated fluid, can change the stability of zeolites.

The formation of zeolite occurs in a low-pressure regime (<200 MPa) (Bish and

Ming 2001). The occurrences of scolecite and laumontite indicate maximum

formation temperatures of 280 - 300 °C by using equilibrium phase modeling in the

system CaO-Al2O3-SiO2-H2O-CO2 (Fig. 3.17). However, in situ temperature

measurements in active hydrothermal systems, like in basaltic rocks on Iceland

suggest lower temperatures (e.g. Kristmannsdóttir and Tómasson 1978).

An important factor, which controls the formation of zeolite, is the composition

of the fluid from which the secondary minerals precipitated. Regarding the low

frequency of zeolites in the Lepontine Alps, this lack can be related to CO2 dominated

fluids (Poty et al. 1974; Mullis et al. 1994; Stalder 2007). Zen (1961) noticed that

zeolite mineral assemblages could be obtained by the increase of the chemical

potential of H2O relative to that of CO2, at constant temperature and pressure. At

relatively low CO2 activities, calcium zeolites are destabilized relative to assemblages

contain calcite, quartz and clay minerals (Zen 1961; Senderov 1973) that is supported

by the fluid inclusion evolution to CO2 free fluids. However, the lack of zeolite

inclusions in quartz, suggests that quartz growth was finished before zeolite formation

starts and no information about fluid compositional at the time of zeolite formation is

available. Nevertheless thermodynamic modeling points out that the fluid has to be

low in CO2 (Fig. 3.17). The effect of CO2 bearing fluids on the stability of zeolites

and other Ca-Al-silicates can be seen in the calculated thermodynamic phasediagrams

for different Ca/Al ratios and pressure conditions in the system CaO-Al2O3-SiO2-

H2O-CO2 (Fig. 3.17). For calculation the ideal mixing model for H2O-CO2 was used.

Zeolite species are stable in fluids dominated by H2O with low CO2 concentrations.

With increasing CO2 activity zeolite species are replaced by kaolinite (e.g. Val

Bedretto, Stader et al. 1998) at lower temperature and other Ca-Al silicates, calcite

and quartz at higher temperature (Fig. 3.17). Stilbite is stable only at very restricted

XCO2 less than 0.04 at 10 MPa and low temperature, whereas heulandite, laumontite

and scolecite are stable at higher XCO2 (Fig. 3.17). Scolecite stability is controlled by

pressure conditions and occurs only at lower pressures. The Ca/Al ratio can also

effect the stability variations of zeolites. However the occurrences of heulandite may

also be controlled by additional extra-framework cations like Sr, Na and K, which


prefer to incorporate into heulandite in Alpine fissure (e.g. Weisenberger and Bucher

2009). Generally zeolite stability increases with decreasing pressure and at pressure

condition at about 100 MPa, zeolites will destabilize at XCO2 lower than 0.05 (Fig.

3.17). However, in hydrothermal systems the fluid pressure is unlikely equal to the

total pressure. The assumption of very low CO2 is supported by the absence of

kaolinite, which would be present at lower temperatures with XCO2 > 0.04.

Fig. 3.17: Equilibrium T- XCO2 diagrams at P = 10 MPa, 50 MPa and 100 MPa, for the CaO-Al2O3-SiO2-H2O-CO2 system for different Ca/Al ratios.


3.7.4. Mineral evolution

Bulk rock composition of the unaltered rock of sample Arvigo 12 (Fig. 3.3, Table 3.8)

has been recalculated to atomic proportions. To simplify the diagrams, Ti and Mn

have been ignored. The presence of chlorite, epidote and prehnite indicate that Fe

occurs in di- and trivalent states and that some provision for the redox state is need to

be made. However the fact, of Fe zoning in prehnite, which is interpreted as change in

oxygen fugacity, specification of the redox state is not possible. Thereby modeling

with different redox state conditions was done with the result that the ratio 1/1 of

Fe3+/Fe2+ reflects the best fit with observed secondary mineral inventory. The

alteration is driven by hydrothermal process and therefore H2O was set in excess.

Figure 3.18 gives a PT diagram that is appropriate for hydrothermal alteration

conditions. The corresponding predicted assemblage evolution is shown in Fig. 3.19.

According to Fig. 3.18, epidote would start to form at temperature conditions of

450°C, which is higher than the temperature estimation by using iron distribution in

epidote and prehnite. However, the CO2 rich fluid at higher temperature could be the

reason for the delay and formation of epidote at lower temperature.

Fig. 3.18: Assemblage stability diagram for unaltered rock of sample Arvigo 12 (Fig. 3.3, Table 3.8). Note: note all assemblage fields are labeled.


Chloritization occur at temperature between 350 and 325 °C, depending on pressure

conditions and is in good agreement with the empirical calibrated chlorite

thermometers by Cathelineau (1988) that yields an average chlorite formation

temperature of 333 °C.

The phasediagram points out that K-feldspar is affected by a hydration reaction

and forms muscovite. But in contrast to the computed diagram, the observed mineral

inventory (Table 3.1), where K-feldspar occurs as rock-forming mineral, as well as

fissure adularia, did not reflect the stable assemblage, and therefore K-feldspar

behaves metastable in the hydrothermal Arvigo system. Considering the chloritization

reaction (Eq. 6, 7), which releases Al2O3 during reaction, the phasediagram has to be

regarded with care.

Fig. 3.19: Predicted assemblage evolution during hydrothermal alteration, calculated along a linear exhumation path diagonal through Figure 3.18.

Metasomatic reactions depend on the compositions of the fluid phases leaving and

entering the system, and cannot be thermodynamically treated by merely considering

the solid phases in the same way that isochemical metamorphic reactions are


traditionally used to determine PT conditions. Nevertheless the phase-diagram reflects

a good approach to model the mineral evolution, because prehnite and the zeolite-in

reaction reflect a plausible mineral evolution, if we consider field observations and

compare them with other hydrothermal systems. In any case, which kind of zeolite is

formed is also related to cation substitutions and might increase the stability in

contrast to the pure Ca endmember, which lack in the thermodynamic data-base, like

the incorporation of Sr into heulandite.

Fig. 3.20: Porosity evolution during hydrothermal alteration. Minerals that were precipitated during hydrothermal alteration were assumed to precipitated in the fissure, whereas elements necessary for formation were moved out from the host rock, producing porosity. One path reflects the calculation, that assumed the total removal of chlorite and one reflects the porosity path, whereas chlorite occurs in the rock matrix.

Using the predicted assemblage evolution along the cooling path (Fig. 3.18, 3.19) and

assuming that the elements for the secondarily formed Ca-Al-silicates in the fissure

were derived from the adjacent wall rock, the porosity evolution can be calculated by

removing the molar volume portion of the secondary mineral, from the initial molar

rock volume (Fig. 3.20). Figure 3.20 therefore shows the temporal porosity evolution

along the PT path (Fig. 3.18).


One path is modeled by assuming that all chlorite remains in the wall rock,

whereas the second path requires complete chlorite removal into the fissure space and

resulted porosities of 9.2 and 11.3 vol. %. In contrast porosity calculation using modal

mineral composition gives a porosity of ~17 %. The difference could be related to the

2-dimensional analysis of porosity of an anisotropic texture or to the meta-stability of

K-feldspar in phase diagram calculation and therefore the non-consideration of K-

feldspar into the porosity calculation.

3.7.5. Fluid accessibility and composition

The alteration process, forming hydrous Ca-Al-silicates requires a significant amount

of H2O. Fluids have to infiltrate the wall rocks, where fractures act as fluid channels

on outcrop scale (e.g. Austrheim 1987; Bons 2001) as well as on microscale (e.g. Fitz

Gerald and Stünitz 1993; Oliver 1996).

Fracturing is caused by brittle deformation that is younger than the main Alpine

deformation and related to the uplift of the Central Alps 10-20 Ma ago (Steck 1968;

Purdy and Stalder 1973). Two distinct fracture directions can be observed in Arvigo,

which differ in mineralogy and suggest a change in the stress field with time. The

later formed fissures are mineralized with zeolites. Recent dating of the latest fissure

minerals in the Central Alps (Weisenberger and Bucher 2008) suggests a younger age

(∼2 Ma) of zeolite formation in the Central Alps.

Biotite as well as chlorite occur as connected cluster due to foliation and provide

migration pathways for the fluid through the whole altered zone and increase into

fissure direction. In contrast, the albitization process produces porosity that is constant

over the whole sharp alteration front (Fig. 3.3). In general, the porosity of the

albitization is the volume occupied by the fluid phase and is generated at the reaction

interface where the volume of dissolved plagioclase is less than the volume of albite

that reprecipitates. However, only the interconnected porosity provides permeability.

Nevertheless, the porosity which is generated during volume loss of albitization is

enriched at the grain boundaries (Fig. 3.8) and marks a prominent fluid channel,

whereas intragranular porosity may not affect the fluid permeability due the poorly


connected pores. Both alteration reactions, chloritization of biotite and albitization,

enhance the permeability.

If we include secondary minerals into the volume equation and neglecting that

some elements are transported away or added from the fissure wall, a volume increase

has to be assumed. This process of precipitation of secondary in the adjacent area of

the wall rock can impregnate earlier formed porosity in the wall rocks, as well as the

fracture. This decrease hereby the permeability to the point that the fluid low is

interrupted and the alteration process is terminated due to the absent of fluid. This is

seen at the chlorite vein in Fig. 3.3, where the adjacent plagioclase is depleted in Ca

(Fig. 3.5, 3.6), but less depleted than the albite crystals in the appreciable alteration

zone. These Ca remaining in plagioclase can be related to inaccessibility due to the

impregnation of the fluid channel by secondary chlorite.

Early CO2 dominated fluids may derive from decarbonation processes or

oxidizing of organic matter of Mesozoic metasedimentary rocks, which are integrated

in the nappe-stack of the Lepontine Alps (e.g. Poty et al. 1974; Mullis et al. 1994).

Magmatic waters seem to be an unlikely source regarding the geological setting and

the absence of magmatic intrusions. If the CO2 was formed in lithological units below

or above and migrated to the fissure remains as open question. The change in fluid

composition could be caused due to infiltration of meteoric waters or by the lack of

the sources for decarbonation, due to the erosion of the metasedimentary units above

the Simano nappe. However, if we consider the element mobility (Fig. 3.13) the

change in element concentration and mineralogy does not need infiltration of

chemically exotic fluids.

3.8. CONCLUSION

Low-grade mineral assemblages are the key to the appreciation of water-rock

interaction in hydrothermal and geothermal systems located in granites and gneisses.

The Arvigo locality is a example for a crystalline basement unit consisting of granites

and gneisses, which is significantly affected by fracture-related hydrothermal

alteration. Fissures and gashes formed by semi-brittle deformation were generated


during exhumation and uplift of the Alpine orogen. These fractures and cavities were

filled with fluids and new minerals crystallized in the open space.

The Arvigo fissures contain the assemblage epidote, prehnite, chlorite and

various species of zeolites. In general epidote is overgrown by prehnite, chlorite and

zeolites. Ca as extra-framework cation dominates all zeolites, whereas the specific

zeolite formed in the fissures depends on the temperature. Following Ca-dominated

zeolites precipitated from the low-CO2 aqueous fluid with decreasing temperature:

scolecite, laumontite, heulandite, chabazite and stilbite.

The composition of coexisting prehnite/epidote reveals temperature conditions

between 330 and 380 °C for the pre-zeolite assemblage using the Rose and Bird

(1987) calibration. The iron zoning pattern in prehnite suggest elevated oxygen

fugacity during early growth of prehnite. AlIV occupancy on the octahedral site in

chlorite (Cathelineau 1988) suggests temperature conditions of 333 ± 32 °C. Zeolite

formation takes place at temperatures below 250°C.

Fluid induced mineral reactions occurred during the hydrothermal alteration of

rock-forming minerals in the wall rock. The reactions are marked by the albitization

of plagioclase accompanied by chloritization of biotite, forming a reaction front

propagating from central fractures into the gneiss matrix. A first replacement reaction

changes biotite into chlorite within a 3 to 7 cm thick zone of the host rock. The

plagioclase replacement reaction releases components for zeolite formation and forms

a sharp reaction front in the gneiss at about 2 to 2.5 cm from the central fracture.

The albitization reaction is associated with a volume decrease for the solids.

Thereby albite remains as daughter phase during in the wall rock and exhibit a

porosity increase of ~16 %, whereas the anorthite component get dissolved. We

conclude that much of the produced volume is transferred to the central extension

fracture by laumontite precipitation in the open fracture. The porous product albite

suggests that the propagation of the reaction front through the gneiss matrix occurred

via a dissolution-precipitation mechanism. Chloritization is accompanied by the

release of K+, Fe2+, Al3+ and SiO2 to be volume conservative.

Temperature controlled advection can be assumed to control the transport of

dissolved elements into fissure direction. The mineral evolution along an exhumation

path is conforming to petrographic and mineralogical observations.


Although the calculations are isochemical the results are consistent with the

observations including wall rock minerals and fissure minerals. Theses suggest that

elements that were transferred out of the wall rock, precipitates in the fissure and did

not transported away.

The remarkable and astonishing lack of zeolites in late fissures in the Lepontine

Alps, compare to the Arvigo locality could be related to pCO2 above critical threshold

value that makes zeolite formation impossible. The calculated phase diagrams in the

system Ca-Al-Si-O-C-H encourage the fluid evolution to CO2 poor fluids with time,

which is observed in fluid inclusions in quartz, that were formed prior to the zeolite

formation.

Mass balance calculations for the whole rock suggest a mass loss of 6.8 % and

depletion in Al2O3, SiO2, CaO, Fe2O3 and K2O in the altered wall rock. These

elements are subsequently found as major components in epidote, prehnite, calcite,

adularia, chlorite and zeolites as fracture filling minerals. The mass transfer is

associated with an increase in porosity, caused by the volume decrease during

albitization and the removal of chlorite in the wall rock.

The mineral paragenesis in low-grade rocks, often result from fluid-rock

interaction alteration. The mineralogical, geochemical and textural signatures, caused

by low-grade metamorphism, can be interpreted by relatively simple paragenetic

schemes, which can be linked to the tectonic and thermal history of the rocks as well

on the fluid evolution.


We are grateful to Giovanni Polti and Alfredo Polti SA for permission to do field

work in the active quarry. Special thanks go to the technicians and staff of the

Institute of Geosciences, Mineralogy – Geochemistry, University of Freiburg and

particularly Hiltrud Müller-Sigmund for her useful advise during EMP analyses and

her patience with us at the electron microprobe. Andreas Leemann from the Swiss

Federal Laboratories for Materials Testing and Research for impregnation of rock


samples. A special thanks deserved to the Friedrich Rinne foundation for the financial

support.

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TIMING AND MINEARL EVOLUTION OF LOW-TEMPERATURE FLUID-ROCK INTERACTION 132

4. TIMING AND MINERAL EVOLUTION DURING

LOW-TEMPERATURE FLUID-ROCK INTERACTION

ON UPPER CRUSTAL LEVEL: 40Ar/39Ar

APOPHYLLITE-(KF) DATING AND APATITE FISSION

TRACK ANALYSIS ON ALPINE FISSURES (CENTRAL

ALPS/SWITZERLAND)


4.1. ABSTRACT

The mineral assemblage quartz, laumontite and apophyllite-(KF) occur in a fissure

within the Southern Aar granite located in the Aar Massif (Switzerland). They were

formed during exhumation of the Alpine orogen and laumontite and apophyllite

marks the latest fissure minerals in the Central Alps. A combined study of 40Ar/39Ar

age dating, apatite fission track (FT) and chemical characterization of tunnel and

surface samples are present to carry out the position of low-temperature water-rock

interaction in respect to the Alpine history.

Apatite FT analysis yields an exhumation rate of 0.45 mm a-1, a cooling rate of 13

°C Ma-1 and a geothermal gradient of 28 °C km-1. Combining these with the 40Ar/39Ar

plateau age for apophyllite of ∼2 Ma, a minimum formation temperature and depth of

70 °C and 2800 m, respectively can be assumed. Temperature-time evolution of

fissures in the Aar Massif and thermodynamic mineral evolution indicate that

laumontite were formed between 7 and 2 Ma before present at temperatures between

150 and 70 °C.

Elements for laumontite formation derived during dissolution of primary

minerals. Changes of laumontite chemistry could be an effect of temperature drop or a

change in fluid chemistry that would be supported by later apophyllite formation.

Keywords: laumontite, apophyllite-(KF), 40Ar/39Ar, apatite fission track, granite

4.2. INTRODUCTION

Fluid-rock interaction is an important process in the upper crust, with respect to

porosity evolution, permeability and fluid migration. The fluid composition monitors

the water-rock interaction and controls the dissolution of primary minerals and the re-

precipitation of secondary minerals in open spaces (e.g. Nordstrom et al., 1989;

Bucher and Stober, 2001). Thereby the formation of zeolites and apophyllite is

widespread in basaltic rocks (e.g. Walker, 1959; Belsare, 1969; Sukheswala et al.,


1974; Keith and Staples, 1985; Young et al., 1991; Neuhoff et al., 1997;

Weisenberger and Selbekk, 2009) as well as in upper continental crust, particularly in

hydrothermal fractures and veins in granites and gneisses (e.g. Borchardt et al., 1990;

Borchardt and Emmermann, 1993; Armbruster et al., 1996; Freiberger et al., 2001;

Fujimoto et al., 2001; Ciesielczuk and Janeczek, 2004, Weisenberger and Bucher,

2009). The formation of zeolites requires a H2O dominated fluid (Zen, 1963;

Senderov, 1973; Weisenberger & Bucher, 2009) and is restricted to low temperature

(<250 °C), low pressure (<200 MPa), water-saturated environments.

Considering deep continental fluids, they have the potential to form zeolites

(Stober and Bucher, 2004). High-pH waters from the NEAT tunnel in the basement of

the Swiss Aar Massif (Fig. 4.1)(Seelig et al., 2007), water from the crystalline

basement at Stripa, Sweden, (Nordstrom et al., 1989), Bad Urach (Stober and Bucher,

2004) and from the Black Forest basement (Bucher and Stober, 2000) are all

oversaturated in respect of zeolites. Therefore a detailed study of fissure minerals can

give important information about the hydrogeochemical evolution in the upper

continental crust.

The formation of zeolites and apophyllite-(KF) marks the last step of the long (20

Ma) history of fracture generation and mineralization as result of the uplift and

exhumation of the Alpine orogen (e.g. Weisenberger & Bucher, 2009).

The Gotthard-NEAT tunnel is a good example to study subsurface samples,

approximately 2000 m below the surface, which are usually not accessible. For

instance laumontite is the most widespread zeolite in Alpine fissures, exposed

underground in tunnels sections or in active quarries. Because the mineral

decomposes by dehydration at room temperature and decays to a powdery mass,

laumontite occurs rarely in surface outcrops (Weisenberger and Bucher, 2009). The

timing of low-temperature water-rock interaction is commonly difficult to establish

because of the paucity of suitable material for geochronology. The appearance of

apophyllite-(KF) following laumontite during late stages of Alpine fissure

mineralization gives the opportunity to get information on the age of this event.

Apophyllite as tool of age dating by 40Ar/39Ar techniques is not used widely. However

Fleming et al. (1999) and Molzahn et al. (1999) evaluated the feasibility of using

apophyllite for geochronology by the 40Ar/39Ar method on secondary mineralization


in the Transantarctic Mountains and demonstrated that apophyllite dating can produce

geologically meaningful ages.

The aim of this study is to assess information about the timing of low-

temperature mineralization in the upper continental crust and carry out the relation to

the temporal evolution of fissure mineralization during fluid-rock interaction at

Alpine exhumation stage.

To assign these information geochemical characterization of fissure minerals are

presented, as well as Ar/Ar age determination on apophyllite and apatite FT analysis

to obtain the local exhumation rate and therefore estimate the formation depth of late

stage fissure minerals.


The Aar Massif is one of the external massifs of the Central Swiss Alps (Frisch et al.,

1990) situated in the Helvetic zone (Fig. 4.1). The Aar massif is formed by Hercynian

intrusives, emplaced into a polymetamorphic basement (Fig. 4.1) of Paleozoic to late

Proterozoic age (Grünenfelder et al., 1964; Gulson and Rutishauser, 1976). The

Upper Carboniferous Central Aar granite is a lens-shaped batholith, consisting of

granites and granodiorites, and is exposed over an area of about 550 km2. The

Southern Aar granite is only exposed in the eastern part and can be traced over 20 km

in W-E direction and 1-2 km in N-S direction (Fig. 4.1). The Southern Aar granite has

earlier been considered as marginal southern facies of the Central Aar granite (Huber

1948), but modern age determination indicates that the Southern Aar granite is not

genetically related to the Central Aar granite. Schaltegger and Corfu (1992)

determined (U-Pb- method on zircon and allanite) the age of the emplacement of the

Southern Aar granite took place at around 350 Ma. This age predates the late

Hercynian Central Aar granite, which was emplaced in a short period of 2-4 Ma at

around 298 Ma (Schaltegger and Corfu 1992).

The Aar Massif was subject to Cenozoic Alpine greenschist-facies

metamorphism. The N-S transection is marked by different isograds with increasing

metamorphic grade from the north to the south (Bambauer and Bernotat, 1982; Frey


et al., 1980; Frey and Mählmann, 1999): the first appearance of green biotite (Steck

and Burri, 1971), disappearance of stilpnomelane (Jäger et al., 1967) and the

transformation isograd of microcline/sanidine (Bambauer and Bernotat, 1982;

Bernotat and Bambauer, 1982; Frey and Mählmann, 1999). Peak metamorphism

exceeds zeolite facies all-over.

FIG. 4.1: (a) Detailed geological map of the Eastern Aar Massif (modified after Labhart, 1977) including sample locality and track of the Gotthard new railway base tunnel (NEAT). (b) Outline of Switzerland and the position of the central external massifs in gray. Rectangle mark the section a.

During late-orogenic exhumation the considerable increase of erosion rate and

denudation forced the evolution of shear zones related to backthrusting, parallel

normal faulting and the opening of fissures and gashes. These act as pathways for

fluids which seep through and react with the surrounding rocks to finally form

secondary fissure minerals (Berger et al., 2005; Mullis 1995, 1996). An exhumation


rate of 0.5 mm a-1 for the Reuss valley in the Central Aar Massif is proposed by

Michalski and Soom (1990) for the past 27 Ma (apatite and zircon FT), with a cooling

rate of 13 °C Ma-1, which are in good agreement with uplift rates of 0.3 - 0.6 mm a-1

during the last 6-10 Ma (Schaer et al., 1975). Using exhumation rates and trapping

temperatures of early fluid inclusions (Mullis, 1996) the first opening of fissures and

precipitation of fissure minerals in the Aar- (Zinggenstock) and Gotthard Massif (La

Fibbia) is determined to around 20 Ma ago.

4.4. SAMPLES AND METHODS

4.4.1. Analytic

Whole rock analyses were performed by standard X-ray fluorescence (XRF)

techniques at the Institute of Geosciences (Mineralogy and Geochemistry) at the

University of Freiburg/Germany, using a Philips PW 2404 spectrometer. Pressed

powder and Li-borate fused glass discs were prepared to measure contents of trace

and major elements, respectively. The raw data were processed with the standard XR-

55 software of Philips. Relative standard deviations are < 1 % and < 4 % for major

and trace elements, respectively.

Quantitative mineral analyses were performed at Institute of Geosciences

(Mineralogy and Geochemistry), University of Freiburg, using a CAMECA SX 100

electron microprobe equipped with five WD spectrometers and one ED detector with

an internal PAP-correction program (Pouchou and Pichior, 1991). Major and minor

elements in zeolites were determined at 15 kV accelerating voltage and 10 nA beam

current with a defocused electron beam of 20 µm in diameter with counting time up to

20 s. Na and K were counted first to minimize the Na and K loosed during

determination. Since zeolites lose water when heated, the crystals were mounted in

epoxy resin to minimize loss of water due to the electron bombardment. Natural and

synthetic standards were used for calibration. The charge balance of laumontite

formula is a reliable measure for the quality of the analyses and correlates with the

difficulties related to the thermal instability of zeolites in microprobe analysis. A

useful error test investigates the charge balance between the non-framework cations


and the amount of tetrahedral Al (Passaglia, 1970). Analyses are considered

acceptable if the sum of the charges of the extra-framework cations (Ca2+, Sr2+, Ba2+,

Na+, and K+) is within 10% of the framework charge (Al3+).

Isotopic dating was carried out at the 40Ar/39Ar laboratory of the Department of

Mineralogy, University of Geneva, Switzerland. Crystals of apophyllite were crushed

and clear, inclusion free chips were packed in copper foil. The samples were

irradiated for 3 hours at 1 MW in the Oregon State University CLICIT facility, and J

values were calculated via the analysis of Fish Canyon Tuff sanidines, which were

spaced by <1cm throughout the columnar irradiation package. Stepwise degassing

was performed using a 30W CO2-IR laser, and extracted gas was purified in a UHV

extraction line equipped with SAES AP10 and GP50 getters, prior to analysis. Isotope

ratios were measured with a GV instruments ARGUS multi-collector mass

spectrometer, equipped with four high-gain (10-12 Ohms) Faraday collectors for the

analysis of 39Ar, 38Ar, 37Ar and 36Ar and one single 10-11 Ohms Faraday collector for

the analysis of 40Ar. Blanks were measured between every three degassing steps and

before every new sample. Data reduction was performed using the program

ArArCALC (Koppers, 2002) and corrections were applied for post-irradiation decay

of 37Ar (T0.5 = 35.1 days) and 39Ar (T0.5 = 269 years). The mass discrimination factor

during analysis was 1.00436 based on ongoing measurements of 40Ar/36Ar ratios in

quantitatively calibrated air shots from an air pipette. Correction factors for

interfering Ca- and K- derived isotopes have been calculated from 10 analyses of two

Ca-glass samples and 22 analyses of two pure K-glass samples, and are: 36Ar/37Ar(Ca)=2.603E-4 ± 2.373E-9, 39Ar/37Ar(Ca)=6.501E-4 ± 7.433E-9 and 40Ar/39Ar(K)=1.547E-2 ± 7.455E-7.

For the FT measurement, apatite grains were separated from sample 115900 and

SueArGr (8-10 kg rock material) using standard crushing, magnetic and heavy liquid

techniques. Separated apatites (fraction 63-300 µm) were mounted with epoxy on

glass slides, polished, and etched for 20 s with 5N HNO3 to reveal the spontaneous

fission tracks. Mounts were covered with U-free white mica sheets and sent to

irradiation at the FRM-II reactor facility in Garching/Germany, together with other

samples and top and bottom CN5 dosimeter glasses. After irradiation, the white micas

were removed and etched for 45 minutes in 40 % HF to reveal the induced tracks.

Fission tracks were recorded and confined track lengths and Dpar measured using


transmitted and reflected light at 1600x magnification on a Zeiss Axioplan

microscope, equipped with a computer-driven stage and a digitizing tablet at the

Geosciences Department of Basel University. 40 grains were counted per sample in

order to reduce the age error of the expectedly young ages. Central ages (Galbraith

and Laslett, 1993) were calculated using the IUGS-recommended zeta calibration

approach (Hurford and Green, 1983). Statistical χ2 tests were applied to search for

internal variation assuming the single grain ages of mono-population samples to be

Poissonian distributed. Failure of this test (P(χ2) < 5%) may indicate the presence of

internal age variation due to partial annealing or chemical inhomogeneities.

FIG. 4.2: Photograph and schematic illustration showing fissure mineral assemblag at the 115935 (Table 4.1) locality in the NEAT tunnel from the sample KB868. (a) Photograph of fissure assemblages from the same fissure than sample KB868: clear apophyllite “cubs” overgrown by laumontite needles. (b) Schematic sketch on thin section scale of sample KB868 fissure mineral succession Qtz → Lmt → Apo. (c) Representative microphotographs under crossed polarized light.


4.4.2. Samples

All sample sites are located in the Southern Aar granite (Fig. 4.1, Table 4.1). Sample

KB868 is collected from a fissure during excavation of the Gotthard NEAT tunnel,

15935 m south of the north portal (Fig. 4.1, 4.2). The sample exhibits a fissure

assemblage with the chronological order: quartz, adularia, laumontite and apophyllite

(Fig. 4.2). Additionally chlorite and milarite are found in the same fissure (P.

Amacher pers. com.). Euhedral fissure minerals, up to 12 mm in size (Fig. 4.2) grew

on a thin leached matrix (< 1 cm in thickness). Leaching is indicated by higher

porosity than in the fresh tunnel sample 115900 (Table 4.1) from the same lithological

unit of the Southern Aar granite. A third sample (SueArGr) is collected from a surface

outcrop of the Southern Aar granite (Fig. 4.1, Table 4.1) The vertical offset between

the tunnel- (115900, KB868) and the surface specimen (SueArGr) is 1623 m (Fig. 1,

Table 4.1).

TABLE 4.1. Sample description. Indicated x and y coordinates corresponds to the Swiss coordinate net (units: km). Sample No x y altitude [m] Description Assemblage* SueArGr 700 588 174 999 2 123 Rock sample of the Southern Aar granite

collected on a surface outcrop. XRF- analysis, apatite FT

Qtz, Kfs, Pl, Ms, Chl, Ep, Ap, Ttn, Rt, Py, (Bt)

115 900 700 062 173 863 ∼500 Drill core of the Southern Aar granite collected in the Gotthard-NEAT tunnel; 115900 m in distance to the north portal. XRF- analysis, apatite FT.

Qtz, Kfs, Pl, Ms, Chl, Ep, Ap, Zrn, Rt

KB868 (115 935)

700 079 173 832 ∼500 Fissure assemblages in the Southern Aar granite: quartz, adularia, laumontite and apophyllite; additionally chlorite and milarite are found in the same fissure. Fissure: 40 x 60 x 30 cm. 15935 m in distance to the north portal in the Gotthard-NEAT tunnel

Apo, Lmt, Kfs, Qtz, Chl, Mil

*Abbreviations according to Bucher and Frey (2002); Mil = milarite, minerals in parentheses are metastable

4.5. RESULT

4.5.1. Petrography and geochemistry

The Southern Aar granite consists predominantly of albite, K-feldspar, quartz,

muscovite and chlorite. The rock is deformed with K-feldspar megacrysts up to 1 cm


in length in a matrix of plagioclase, quartz, muscovite and chlorite as mafic

components, usually not exceeding grain sizes of some mm (Fig. 4.3). Both rock

samples (115900, SueArGr) exhibit the same mineralogy (Table 4.1), whereas the

sample from the surface (SueArGr) is strongly be weathered.

FIG. 4.3: Representative microphotographs of mineral assemblages of rock sample 115900 and SueArGr. (a) Recrystallization quartz, saussuritizated plagioclase, chlorite and K-feldspar in sample 115900; plane polarized light. (b) Fracture between saussuritizated plagioclase filed with epidote in sample SueArGr; crossed polarized light. (c) Epidote flakes in between chlorite sheets in sample 115900; crossed polarized light.


TABLE 4.2. Bulk rock geochemistry SueArGr 115 900 wt % SiO2 68.86 67.50 TiO2 0.42 0.47 Al2O3 16.23 15.46 Fe2O3

tot 2.62 3.16 MnO 0.06 0.06 MgO 1.19 1.29 CaO 2.10 2.66 Na2O 4.15 3.78 K2O 3.56 4.03 P2O5 0.22 0.23 LOI 1.05 0.91 Totals 100.67 99.80 ppm V 37 44 Cr 20 33 Ni 16 23 Cu 4 22 Zn 55 56 Rb 161 145 Sr 572 693 Zr 162 248 Ba 1071 1208

Plagioclase is altered to sericite, whereas the fine muscovite flakes are concentrated in

the cores of the plagioclase grains. Saussuritization of plagioclase causes the

formation of epidote that occurs as small inclusion therein, as well as interstitial

filling in chlorite and in small (< 100 µm) veins (Fig. 4.3). Quartz crystals (Fig. 4.3)

show a characteristic fabric of dynamic recrystallization by subgrain rotation that was

caused during Alpine deformation. The assemblage K-feldspar and chlorite suggests

that the mineralogy of the Southern Aar granite has been altered and retrogressed to

temperatures below 400°C (Bucher and Frey, 2002). A few altered relictic biotite

grains are present. Pseudomorphic replacements of chlorite after biotite often show

preserved sagenitic intergrowth. The K-feldspar is microcline that shows a tartan

plain pattern and often exhibits perthitic exsolution that indicates that the samples

come from locations north of the microcline/sanidine transformation isograd

(Bambauer and Bernotat, 1982; Bernotat and Bambauer, 1982; Frey and Mählmann,

1999). Accessory minerals include apatite, titanite, pyrite, allanite and zircon.

The peraluminous Southern Aar granite (Table 4.2) shows a slightly lower SiO2

content than the Central Aar granite (69.85 ± 3.60; Schaltegger, 1990) that is derived

from calc-alkaline magmatism during the Hercynian orogenesis (Schaltegger and


Corfu, 1992). MgO, TiO2 and Fe2O3 contents are very low, whereas the K2O und

Na2O contents are elevated. Ca values are 2.10 and 2.66, respectively, slightly higher

to the Central Aar granite of the Reuss valley (1,77 ± 0.78; Schaltegger, 1990) and

significant higher than accordant units in the Grimsel area (Schaltegger, 1990). The

trace elements are dominated by Ba and Sr. The bulk rock geochemistry of both

samples (115900 and SueArGr) shows no major differences (Table 4.2).

4.5.2. Mineralogy and geochemistry

4.5.2.1. Laumontite

Laumontite is a monoclinic (space group C2/m) zeolite. It forms thin, elongated

fibbers or prisms elongated along the c-axis with a squared cross-section (Fig. 4.2).

Twinning occurs on {100} to form “swallow tail” or “V” twins. It is white with a

length between <1 to 12 mm.

FIG. 4.4: Chemical variation in laumontite from sample KB868 as function of the Si/Al ration and extra-framework cations.

The composition of laumontite was obtained on sample KB868 (Table 4.3) and is

close to endmember composition Ca4(Al8Si16O48) •18 H2O (Armbruster and Kohler,

1992). Ca is the dominant extra-framework cation (average value of 96 mole%), with


Na and K typically below 5 mole% (Table 4.3, Fig. 4.4). Additionally Ba is

incorporated as extra-framework cation up to 3 mole%. Sr occurs only in traces (Fig.

4.4). The contents of Na, K and Ba increase from core to rim of zoned crystals (Fig.

4.4, 4.5). Often a second laumontite generation is observed (Fig. 4.5), enriched in K

and Na. The Si/(Si+Al) ratio varies slightly between 0.67 and 0.69 (Fig. 4.4), with an

average of 0.68. Extra-framework cations vary with the Si/(Si+Al) ratio (Fig. 4.4).

With increasing Si/(Si+Al) ratio the Ca content decreases, while Na, K and Ba

increase. This can be expressed by the coupled substitution of Si4+ + (Na+, K+, Ba2+) =

Al3+ + Ca2+.

TABLE 4.3. Laumontite chemistry Sample no. KB 868.3 KB 868.3 KB 868.3 KB 868.2 KB 868.1 KB 868.1 Analysis no. 4 6 7 10 15 16

wt.% SiO2 53.12 52.78 53.42 53.47 52.76 53.85 Al2O3 21.52 22.12 21.58 21.06 20.83 21.33 CaO 11.72 12.24 11.68 11.49 11.33 11.56 SrO 0.00 0.04 0.00 0.02 0.00 0.00 BaO 0.25 0.06 0.33 0.42 0.48 0.45 Na2O 0.09 0.03 0.12 0.16 0.18 0.16 K2O 0.15 0.07 0.00 0.15 0.02 0.19 Totala 86.85 87.36 87.20 86.78 85.62 87.55 Si 16.228 16.044 16.250 16.357 16.352 16.334 Al 7.748 7.925 7.737 7.593 7.609 7.625 Ca 3.836 3.986 3.807 3.766 3.762 3.757 Sr 0.000 0.007 0.000 0.004 0.000 0.000 Ba 0.030 0.007 0.039 0.050 0.058 0.053 Na 0.053 0.018 0.071 0.095 0.108 0.094 K 0.058 0.027 0.000 0.059 0.008 0.074 O 48 48 48 48 48 48 E%b -1.22 -1.51 -0.34 -2.57 -2.03 -2.09 Si/(Si+Al) 0.68 0.67 0.68 0.68 0.68 0.68

aTotals include traces of Mg, Mn and Fe. b E % = (100*((Al)-(Na+K)+2(Mg+Ca+Sr+Ba)/(Na+K)+2(Mg+Ca+Sr+Ba)), measure of charge balance

4.5.2.2. Apophyllite-(KF)

Apophyllite (KCa4Si8O20(F,OH) •8 H2O) occurs as overgrowth on laumontite in

Alpine fissures (Fig. 4.2). It forms transparent tetragonal pseudo-cubes with truncated

spikes of rhomboid faces that end in a pyramid. Sizes are up to 1.5 cm. The apex is

truncated in which case the appearance is cubic (Fig. 4.2).


TABLE 4.4. Apophyllite chemistry Sample no. KB 868.1 KB 868.1 KB 868.2 KB 868.2 KB 868.3 KB 868.3 Analysis no. 2 3 10 11 16 18 wt % SiO2 50.47 50.86 50.63 50.64 50.18 50.96 Al2O3 0.99 0.76 0.85 0.96 0.76 0.95 CaO 24.25 24.33 24.31 24.38 24.29 24.35 Na2O 0.27 0.11 0.11 0.14 0.10 0.25 K2O 4.32 4.63 4.68 4.67 4.57 4.18 F 1.95 2.06 2.02 2.06 2.02 2.06 -O≡F 0.82 0.87 0.85 0.87 0.85 0.87 Totala 81.43 81.94 81.77 82.03 81.11 81.95 Si 7.857 7.884 7.868 7.849 7.866 7.875 Al 0.182 0.139 0.156 0.175 0.140 0.173 Ca 4.045 4.041 4.048 4.048 4.080 4.032 Na 0.081 0.033 0.033 0.042 0.030 0.075 K 0.858 0.916 0.928 0.923 0.914 0.824 Total 13.022 13.020 13.035 13.046 13.032 12.985 O 20 20 20 20 20 20 F 0.960 1.010 0.993 1.010 1.001 1.007 K/Cab 0.178 0.190 0.193 0.192 0.188 0.172

aTotals include traces of Ba, Fem Mg, Mn and Sr. b ratio of weight

FIG. 4.5: K (a) and Na (b) element concentration map for laumontite in sample KB868. The electron microprobe concentration map clearly shows 2 distinct laumontite generations. Noticeable the sector zoning in the older generation.


The chemical composition of apophyllite from the NEAT tunnel is given in Table 4.4.

It is near the stoichiometric composition of apophyllite-(KF) end-member. The

individual crystals are relatively homogenous in terms of chemistry, with small

differences in concentrations of minor components like Al and Na, with average

values of 0.14 and 0.05 mole%, respectively and no other elements occur in

apophyllite.

4.5.3. Ar/Ar age

The apophyllite-(KF) crystals (TW003-APO, ATW-APO) from sample KB868 were

selected as age marker for secondary mineral formation. The 40Ar/39Ar total fusion

and incremental-heating results for the samples are summarized in Table 4.5 and Fig.

4.6. The K/Ca ratio measured from nucleogenic Ar isotopes has the value of 0.19 and

0.18, respectively and is in good agreement with the values measured by electron

microprobe 0.17-0.19 (Table 4.3).

For sample ATW-APO a 40Ar/39Ar total fusion age of 2.11 ± 0.06 Ma is obtained

(Fig. 4.6). The error plateau age of the sample is 2.04 ± 0.14, by excluding the first

three heating steps to improve the visual fit. These shows different ages, suggesting

that these steps have either excess radiogenic argon, which they captured from

hydrothermal fluids, impurities on surface of the crystal, or lost radiogenic argon from

close to grain boundary sites.

For sample TW003-APO the weighted plateau age of 1.96 ± 0.08 Ma (Fig. 4.6)

over 10 out of 11 steps and a 40Ar/39Ar total fusion age is 1.83 ± 0.05 Ma. The very

low 36Ar content suggest that almost all of the 40Ar from the sample is radiogenic and

not from fluid inclusions or argon which was trapped during crystallization.

Both samples overlap in age by using the plateau age, which gives the best

statistical requirement and an apparent age of ∼2 Ma is reasonable.

TABLE 4.5. 40Ar/39Ar increment heating ages of apophyllite Sample no. Total fusion

age (Ma) Plateau age (Ma ±2σ)

MSWD 39Ar % of total

Isochron age (Ma ±2σ)

40Ar/36Ar intercept

J

ATW - APO 2.11 ± 0.06 2,04 ± 0.14 9.74 99.50 2.22 ± 0.41 225 ± 118 0.0008215

TW003 - APO 1.83 ± 0.05 1.96 ± 0.08 1.61 56.01 1.46 ± 0.81 407 ± 189 0.0008222


FIG. 4.6: Plots of 40Ar/39Ar incremental-heating spectra for apophyllite.

4.5.4. Apatite fission track analysis

Two samples were collected for apatite FT analysis along a vertical section in the

Southern Aar granite (Fig. 4.1, Table 4.1). The results are presented in Table 4.6 and

Fig 4.7. Apatite FT ages are 9.6 Ma (SueArGr) and 6.1 Ma (115900), respectively,

with 1σ age errors of less than 10 %. Both samples pass the χ2 test, indicating that the


single grain ages belong to one single age population. The confined length

measurements give mean FT lengths of 14.1 and 13.7 µm that corresponds to an

undisturbed steady cooling behavior.

TABLE 4.6: Apatite fission track data. For sample location see Fig. 4.1 and Table 4.1. Sample No

Mineral and No. Crystals

Spontaneous rs

(Ns)

Induced ri

(Ni)

Pχ2

Dosimeter

rd*

(Nd)

Central FT Age (Ma)

(-2σ/+2σ)

D(par) (+/-S.d.)

Mean Track

Length

S.d. of distribution (No. Tracks)

115 900

apatite (40)

0.007 (133)

0.315 (6018)

87 % 16.08 (12518)

6.1 (-1.0/+1.2)

2.33 (± 0.29)

13.70 1.21 (32)

SueArGr

apatite (40)

0.011 (194)

0.316 (5580)

92 % 01605 (12491)

9.6 (-1.3/+1.5)

3.07 (± 0.48)

14.08 1.25 (100)

(i) Track densities are (x107tr cm-2), *=(x105 tr cm-2) numbers of tracks counted (N) shown in brackets; (ii) analyses by external

detector method using 0.5 for the 4π/2π geometry correction factor; (iii) apatite ages calculated using dosimeter glass CN5 with

ζCN5 =344 ± 5; (iv) P(χ2) is probability for obtaining χ2 value for v degrees of freedom, where v = no. crystals – 1; (v) track

length and D(par) data are given in 10-6m, S.d. = 1σ standard deviation.

Dpar measurements (Table 4.6) revealed a significant difference in Cl content (and

thus in Dpar) between the two samples suggesting that the closure temperature of the

surface sample (SueArGr) is slightly higher than the one of the sample from the

tunnel (Donelick et al., 2005). However, preliminary estimations revealed that the

influence of such closure temperature variation on the estimation of the depth of

apophyllite formation would be negligible in comparison to the variation introduced

by the age errors.

FIG. 4.7: Apatite fission track length data.


4.6. DISCUSSION

4.6.1. Mineral reaction

Secondary minerals formed during precipitation of oversaturated hot fluids in respect

of the secondary minerals. Structural and textural evidences, like leaching zones and

porosity increase in the altered wall rock (Weisenberger and Bucher, 2009) imply that

primary minerals of the host rock are dissolved along the fractures and supply

elements necessary for zeolite formation (Eq. 1).

Ca2+ + 2 AlO2- + 4 SiO2,aq + 4 H2O ⇒ CaAl2Si4O12 •4 H2O (Lmt) (1)

Those hot aqueous fluids reach a high degree of super saturation with respect to

zeolites as observed in the NEAT tunnel (Seelig et al., 2007) and in other deep

continental fluids (Urach geothermal site, German continental deep drilling site KTB;

Stober and Bucher, 2004).

Sources of Ca, Al and Si in the wall rock, which are necessary for the laumontite

formation, are albite, clinozoisite, quartz and calcite (Eq. 2, 3). Clinozoisite (epidote),

calcite and albite are present in the host rock as result of Alpine greenschist facies

metamorphism due to the consumption of prealpine plagioclase. Whereas plagioclase

is considered to be the source for elements that form zeolites in rocks of higher Alpine

metamorphism like in Arvigo and in the Gotthard Massif (Weisenberger and Bucher,

2009).

2 Ca2Al3Si3O12(OH) (Czo) + 6 SiO2,aq + CO2 + 11 H2O ⇒ 3 CaAl2Si4O12 •4 H2O

(Lmt) + CaCO3 (Cc) (2)

Keep in mind that the reactions involve a transport step between dissolution and

precipitation. The additional silica necessary for the formation of zeolites during

clinozoisite dissolution may either be derived locally from dissolution of primary

quartz or albite (Eq. 3) or from externally derived SiO2,aq.


This plausible reaction mechanism co-precipitates laumontite and calcite, which

is a common assemblage in Alpine fissures (Weisenberger and Bucher 2009).

However, calcite does not occur in our sample, which may indicates that calcite

saturation was not obtained by the fluid.

Ca2Al3Si3O12(OH) (Czo) + NaAlSi3O8 (Ab) + 2 SiO2,aq + 7 H2O + H+ ⇒

2 CaAl2Si4O12 •4 H2O (Lmt) + Na+ (4)

This reaction consumes albite and clinozoisite and forms laumontite, and is

accompanied by an increase in pH and the total of dissolved solids (TDS). The

proposed reaction is supported by high Na+, high pH, and high degrees of over-

saturation with respect to zeolites in deep groundwater reported from the NEAT

Gotthard rail base tunnel (Seelig et al., 2007).

During laumontite growth an increase of Na, K and Ba can be observed (Fig. 4.4,

4.5). This change in chemistry can either be related to change in formation

temperature, or to change in fluid chemistry during growth.

Since bulk Ca is low to very low in granites of the Aar Massif dissolution of

widespread matrix and fissure fluorite (Stalder et al., 1998) provides some of the Ca

necessary for zeolite development, and also F for late apophyllite growth:

4 CaF2 (Flt) +8 SiO2,aq + 12 H2O + K+ ⇒

KCa4Si8O20(F) • 8 H2O (Apo) + 8 H+ + 7 F- (5)

Reaction (5) consumes K+ and dissolves silica in addition to fluorite. The reaction

releases F- to the water as a by-product of apophyllite (or laumontite) formation. The

proposed reaction mechanism is supported by the presence of leached fissure fluorite

in the Aar granite and by ultra-high fluoride concentrations in hot deep groundwater

reported from the Gotthard rail base tunnel (Seelig et al., 2007). Seelig et al. (2009)

reported pronounced fluoride concentrations in the tunnel waters raging from 5 to 29

mg L-1, whereas fluoride derived mostly from biotite alteration and fluorite leaching.


4.6.2. Depth and temperature estimation

Although many deep continental fluids are oversaturated with respect to zeolites (e.g.

Bucher and Stober, 2001; Seelig et al. 2007), the zeolite forming process could never

been observed in situ. Therefore we used the approach to assign a minimum depth and

temperature by comparing exhumation rate and the age of apophyllite.

The variation in the apatite FT ages reveals the vertical uplift of the two samples

that differs in altitude. The ages gives the time which elapsed between passing the 120

°C isotherm and present time. Thus, a cooling rate and uplift rate can be calculated

using the present rock temperature in the tunnel of 43 °C and a mean annual surface

temperature of ∼0 °C. Therefore a cooling rates of 12.5 °C Ma-1 can be assessed,

which is in agreement to already known cooling rates (Michalski and Soom, 1990).

Using the apatite FT age and the vertical height difference (Table 4.1) of the two

sample localities an exhumation rate of 0.45 mm a-1 can be assessed agreeing with

already known exhumation rates (Schaer et al., 1975; Michalski and Soom, 1990),

observed along the Reuss valley ∼10 km west of the studied sample localities.

This implies that apophyllite with an Ar/Ar age of ∼2 Ma is formed ∼900 m

below the NEAT tunnel level, or -400 m below sea level. This depth can be supposed

as minimum depth of laumontite formation.

The geothermal gradient obtained by the product of cooling rate and the inverse

of the exhumation rate results in value of 28 °C km-1 for the geothermal gradient in

the studied area and coincidence with a achieved geothermal gradient in the Aar

Massif by Reinecker et al. (2008), ranging from 25 - 30 °C km-1.

Taken the rock temperature of 43°C at the sample locality in the NEAT tunnel

and the geothermal gradient the apophyllite Ar/Ar age reflects a minimum formation

temperature for laumontite of ∼70 °C. However, whether the Ar age represents the

formation temperature or a closure temperature below the apophyllite formation

temperature is still unknown. Nevertheless, the appearance of apophyllite in basalts

associated with zeolites (e.g. Betz, 1981; Keith and Staples, 1985) suggests formation

temperatures in the same order.

Considering the overburden of ∼1900 m with respect to the apophyllite sample

locality in the NEAT tunnel a total depth of 2800 m is presumed, which corresponds


to a hydrostatic pressure of 28 MPa. Regarding deep continental drill holes, like the

German continental deep drilling site KTB, laumontite occurs down to depth of

∼5300 m (Borchardt and Emmermann, 1993). Stober and Bucher (2004) suggested

the formation of laumontite at a depth of ∼3500 m, based on the over-saturation of

fluids. However, the laumontite saturation index are calculated on the observed

temperature of 150 °C which is twice as high as assumed minimum temperature in

this study. Therefore an upper limit has to be assessed using thermodynamic modeling

as well as the integration in the Alpine fissure history.

4.6.3. Thermodynamic approach

A thermodynamic approach was used to show the mineral evolution along a PT-path

(Fig. 4.8). Computed assemblages stability diagrams were modeled with the

Theriak/Domino software of de Capitani and Brown (1987) based on the

thermodynamic data by Bermann (1988), Evans (1990), Frey et al. (1991).

Bulk rock composition of the unaltered rock sample 115900 (Table 4.2) has been

recalculated to atomic proportions. To simplify the diagrams, Ti and Mn have been

ignored. The presence of chlorite and epidote indicate that Fe occurs in di- and

trivalent state and that some provision for the redox state is need to be made. Thereby

modeling with different redox state conditions was done with the result that the ratio

1/1 of Fe3+/Fe2+ reflects the best fit with observed mineral inventory on observed thin

section scale (Fig. 4.3). Due to the fact that alteration is driven by hydrothermal

process H2O was set in excess. CO2 was excluded due to the fact that fluid inclusions

in fissure-quartz of the studied region shows a very low CO2 content (Mullis et al.,

1994) and the knowledge that relative low CO2 activities are ample to destabilize

zeolites (e.g. Senderov, 1973; Weisenberger and Bucher, 2008).

Figure 4.8 gives a predicted assemblage evolution along a PT path, whereas

pressure conditions were chosen to be in between hydrostatic and lithostatic pressure.

Calculations were done assuming closed system conditions, with the exception of the

variable H2O content. Albite, quartz, K-feldspar, muscovite, biotite and epidote are

the stabile phases at peak Alpine metamorphic conditions. Along the cooling path

biotite and albite deceases whereas the epidote content increase. At a temperature of


330 °C biotite is replaced by chlorite, which is in good agreement with temperature

estimates from chlorite in the NEAT tunnel using the empirical calibrated chlorite

thermometers by Cathelineau (1988). The phase diagram points out that K-feldspar is

affected by a hydration reaction and forms muscovite. But in contrast to the

calculation, the observed mineral inventory with rock-forming K-feldspar does not

reflect the stable assemblage, and therefore K-feldspar behaves metastable in the

hydrothermal system. Laumontite launch to form at temperature of ∼150 °C under the

consumption of epidote, albite and quartz that is postulated in Eq. 4.

FIG. 4.8: Predicted assemblage evolution during hydrothermal alteration, calculated along a linear exhumation path with H2O in excess.

The mineral reactions and the evolution of fissure minerals involve a transport step

between dissolution and precipitation as the later processes proceed at a spatially

different location in a fluid regime. Nevertheless calculations of the mineral evolution

suggest that reactions took place in a fluid regime with no exotic chemical

composition. Calculations by Dipple and Ferry (1992) showed that changes in major

element concentrations in rocks, which is the case if we assume that elements

necessary for laumontite formation are derived from the dissolution of primary

minerals, do not need infiltration of chemically exotic fluids, but can instead be driven


by aqueous flow along normal temperature gradients under conditions close to local

equilibrium.

4.6.4. Alpine history

Considering the Alpine history, fissure precipitates formed by fluid infiltration and

the fluid rock interaction with primary minerals during uplift and exhumation of the

orogenic units.

FIG. 4.9: Temperature-time path of fissure mineralization during Alpine exhumation in the Central Swiss Alps. Three time paths are given that coincidence with the path evaluated in this study. The shift in the higher temperature area can be explained by the south - north decline of peak metamorphism indicated by the solid triangles (1 = first appearance of oligoclase in the Gotthard Massif (Steck, 1976) 2 = transformation isograd of microcline/sanidine (Bambauer and Bernotat, 1982); 3 = first appearance of green biotite (Steck and Burri, 1971)). The Gibelsbach upper zeolite limit is defined by fluid inclusion measurement in fluorite proceeding the zeolite formation. Schematic illustrations showing the fissure formation in relation to the temperature-time path.

Figure 4.9 gives temperature-time paths during the phase of uplift in the Central

Swiss Alps and shows the time interval at which formation of zeolites is to be likely.

The formation of zeolites and apophyllite marks the last step in the Alpine fissure

history.

The opening of fissures starts after Alpine peak metamorphic conditions by

retrograde passage through the brittle-ductile transition (Fig. 4.9). Fissure quartz,


which is often a substrate mineral on which zeolites grow, formed during earlier

fissure phases. By fluid inclusion analysis different growth generation in fissure

quartzes can be attached to a temperature regime between 250 °C and 450 °C and

pressure conditions between 180 MPa and 440 MPa (Mullis, 1995). Mullis (1995)

linked the quartz population to the temperature-time path of the same area, generated

by radiometric age data (apatite and zircon FT; Rb/Sr in biotite and muscovite). This

yields the time for fissure quartz formation in the Eastern Aar Massif between 21 and

13 Ma before present, which is slightly retarded with respect to the southern Gotthard

Massif (Fig. 4.9).

The textural evidence of no zeolite inclusions in fissure quartz suggests that the

formation of zeolites initially starts after quartz formation was completed and

therefore zeolite formation in the Eastern Aar Massif have to be assume formed later

than 13 Ma before present.

Considering the zeolite locality at Gibelsbach/Fiesch (Valais, south-western Aar

Massif) whereas the assemblages quartz, green fluorite and zeolites occur, fluid

inclusions measurements in fluorite by Armbruster et al. (1996) yields formation

temperatures above 200 °C by assuming a pressure conditions of ∼100 MPa. These

implies that the zeolite formation start to form at temperatures below 200 °C and at

later time (∼10 Ma; Fig. 4.9) than the formation of fluorite was finished.

From this follows that the formation of laumontite can be limited to a time range

between ∼10 and 2 Ma and a temperature range between 200 and 70 °C. Nevertheless

thermodynamic modeling (Fig. 4.8) indicate the formation of laumontite below 150

°C which would forward limited the formation range of the laumontite formation to a

time range between 7 and 2 Ma before present.

4.7. CONCLUSION

Combining different methods, the study low-temperature fluid rock interaction leads

to an evidence for zeolite (laumontite) and apophyllite generation in respect to the

Alpine exhumation history:


(1) Zeolites (laumontite) followed by apophyllite formed as latest mineral in

Alpine fissure. Age measurements indicate an Ar/Ar age of ∼2 Ma for apophyllite,

reflecting the Pleistocene epoch.

(2) The exhumation rate of the studied area is 0.45 mm a-1 that is in the range of

other known exhumation ranges in the Eastern Aar Massif. Apatite FT yield a cooling

rate of 13 °C Ma-1 and a geothermal gradient of 28 °C km-1. Taken the exhumation

rate and the 40Ar/39Ar age of apophyllite a minimum formation temperature and depth

of laumontite of ∼70 °C and 2800 m, respectively can be determined.

(3) Considering a temperature-time path and the thermodynamic approach to

estimate the formation of laumontite in the Southern Aar Granite (Aar Massif),

laumontite are formed between 7 to 2 Ma before present in a temperature range of 150

to 70 °C.

(4) During growth of laumontite the Si/Al ratio, K, Na and Ba increases, whereas

Ca decreases, which could be an effect of temperature drop. However the overgrowth

of apophyllite do not exclude a chemical change in fluid composition during

laumontite and apophyllite growth.

(5) Elements for the formation of laumontite derived during dissolution and

transport of primary minerals (clinozoisite/epidote, albite and quartz) of the wall rock

and no exotic fluid are necessary.


We would like to thank Peter Amacher who provided high-quality mineral specimens

from the Gotthard NEAT tunnel. We are grateful to Alptransit and the geologist

Roger Rütti for sample supply from the tunnel. In addition a special thanks to the

technicians of the Institute of Geosciences (Mineralogy – Geochemistry) University

of Freiburg for the assistance in preparing and analysing samples. A special thanks

deserved to the Friedrich Rinne foundation for the financial support.


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APPENDIX i

APPENDIX

I Own Contribution

Contribution of Tobias Weisenberger on paper manuscript processing presented in

Chapter 2, 3 and 4.

Chapter 2:

Weisenberger1 T. and Bucher1 K. ZEOLITES IN FISSURES OF GRANITES AND GNEISSES OF

THE CENTRAL ALPS. submitted to “Journal of Metamorphic Geology” Own contribution: Idea: 50 %; Data generation: 100 %; Interpretation: 75 %; Manuscript writing: 60 %

Chapter 3:

Weisenberger1 T. and Bucher1 K. POROSITY EVOLUTION, MASS TRANSFER AND

PETROLOGICAL EVOLUTION DURING LOW TEMPERATURE WATER ROCK INTERACTION IN

GNEISSES OF THE SIMANO NAPPE – ARVIGO, VAL CALANCA, GRISONS, SWITZERLAND.

shortly to be submitted to “Contributions to Mineralogy and Petrology” Own contribution: Idea: 90 %; Data generation: 100 %; Interpretation: 80 %; Manuscript writing: 90 %

Chapter 4:

Weisenberger1 T., Rahn1,2 M., van der Lelij3 R., Spikings3 R. and Bucher1 K. TIMING

AND MINERAL EVOLUTION DURING LOW-TEMPERATURE FLUID-ROCK INTERACTION ON

UPPER CRUSTAL LEVEL: 40AR/39AR APOPHYLLITE-(KF) DATING AND APATITE FISSION

TRACK ANALYSIS ON ALPINE FISSURES (CENTRAL ALPS/SWITZERLAND). shortly to be

submitted to “Mineralogical Magazine” Own contribution: Idea: 90 %; Data generation: 40 %; Interpretation: 90 %; Manuscript writing: 90 % 1Institute of Geosciences, Mineralogy - Geochemistry, Albert-Ludwigs-University Freiburg, Albertstr.

23 b, 79104 Freiburg, Germany

2Eidgenössisches Nuklearsicherheitsinspektorat ENSI, 5232 Villigen, Switzerland

3Department of Mineralogy, Université de Genève, Rue des Maraîchers 13, 1205 Geneva, Switzerland

APPENDIX ii

II Publications

Other related contrbutions by the author not included in the thesis.

Peer-Reviewed Papers

[1] SELBEKK R.S. AND WEISENBERGER T. (2005) Stellerite from the Hvalfjördur area,

Iceland. Jökull 55, 49-52

[2] SPÜRGIN S., WEISENBERGER T. AND HÖRTH J. (2008) Das Leucitophyrvorkommen

vom Strümpfekopf im Kaiserstuhl – eine historische und mineralogische

Betrachtung. Berichte der Naturforschenden Gesellschaft zu Freiburg i. Br. 98,

221-244

[3] WEISENBERGER T. AND SELBEKK R.S. (2008) Multi-stage zeolite facies

mineralization in the Hvalfjördur area, Iceland. International Journal of Earth

Sciences, (DOI 10.1007/s00531-007-0296-6)

[4] WEISENBERGER T. AND SPÜRGIN S. (2009) Zeolites in alkaline rocks of the

Kaiserstuhl volcanic complex, SW Germany - new micropobe investigation and

their relationship to the host rock. Geolgica Belgica 12/1-2, 75-91

Talks und Poster Presentations on International Confereneces

[1] WEISENBERGER T. AND BUCHER K. (2006) Zeolites on fissures of crystalline

basement rocks in the Swiss Alps. In Bowmann R.S. and Delap S.E. (eds).

Zeolite `06 - 7th International Conference on the Occurrence, Properties, and

Utilization of Natural Zeolites, Socorro, New Mecixo USA, 16-21 July 2006, p.

253 (Talk)

[2] SELBEKK R.S. AND WEISENBERGER T. (2006) Zeolite facies metamorphism in the

Hvalfjördur area, Iceland. In Bowmann R.S. and Delap S.E. (eds). Zeolite `06 -

7th International Conference on the Occurrence, Properties, and Utilization of

Natural Zeolites, Socorro, New Mecixo USA, 16-21 July 2006, p. 220

APPENDIX iii

[3] WEISENBERGER T. AND BUCHER K. (2006) Zeolites on fissures of alpine

crystalline basement rocks in the Swiss Alps. Berichte der Deutschen

Mineralogischen Gesellschaft, Beih. z. Eur. J. Mineral. Vol. 18, No. 1, p.153

(Talk)

[4] BUCHER K. AND WEISENBERGER T. (2006) Zeolites on fissures of crystalline

basement rocks in the Swiss Alps. Geological Society of America Abstracts with

Programs, Vol. 38, No. 7, p. 113

[5] SELBEKK R.S. AND WEISENBERGER T. (2007) Multi-stage zeolite facies

metamorphism, Southwest Iceland. NGF Winterconference Stavanger, 8.-10.

Januar 2007. NGF Abstracts and Proceedings of the Geological Society of

Norway, no, 1, 90 (Poster)

[6] WEISENBERGER T. AND BUCHER K. (2007) Low-grade zeolite facies

metamorphism in gneisses of the Simano nappe (Arvigo, Val Calanca, Grisons,

Switzerland). Geochimica et Cosmochimica Acta 71(15) Supplement 1, A1100

(Talk)

[7] WEISENBERGER T. AND BUCHER K. (2007) Porosity increase during low-

temperature metamorphism in gneisses of the Simano nappe (Arvigo, Val

Calanca). 5th Swiss Geoscience Meeting, 17-18.11. 2007, Geneva, Switzerland,

104-105 (Poster)

[8] WEISENBERGER T. AND BUCHER K. (2008) Porosity evolution and mass transfer

during low-grade metamorphism in crystalline rocks of the upper continental

crust. 33rd IGC International Geological Congress, 06. - 14.08.2008 Oslo,

MPN03710L (Talk)

[9] WEISENBERGER T. AND BUCHER K. (2008) Ca-Al Silicate Formation During Low-

grade Metamorphism in the Upper Continental Crust. 86th Annual Meeting of

the German Mineralogical Society – DMG 14th -17th September 2008 Berlin,

S18T06 (Talk)

[10] SPÜRGIN S. AND WEISENBERGER T. (2008) Faujasite Growth During

Palagonitisation of Mg-rich Sideromelane: an Example from the Kaiserstuhl

Volcanic Complex, SW Germany. 86th Annual Meeting of the German

Mineralogical Society – DMG 14th – 17th September 2008 Berlin, S18P07

(Poster)

APPENDIX iv

Popular scientific papers

[1] WEISENBERGER T. (2005) Zeolithe in Island. Island; Zeitschrift der Deutsch-

Isländischen Gesellschaft e.V. Köln und der Gesellschaft der Freunde Island

e.V. HAMBURG, ISSUE 1, APRIL 2005, 40-45

[2] WEISENBERGER T. AND SELBEKK R. S. (2006) Die Zeolith-Fundstelle Hvalfjordur,

Island. Mineralien Welt 17, Heft 1, 50-56

[3] WEISENBERGER T., SPÜRGIN S. AND SELBEKK R. S. (2008) Die Fundstelle

Helgustadir (Island): Geologie, Mineralogie und die bedeutende Geschichte des

Isländischen Doppelspats für die Wissenschaft. Aufschluss 1, 53-63

Excursion Guide Books

[1] WEISENBERGER T. AND SPÜRGIN S. (2008) Secondary minerals in the limburgites.

In Keller J.: Tertiary Rhinegraben volcanism: Kaiserstuhl and Hegau. 9th

International Kimberlite Conference, Frankfurt/Main, field trip, 24-25

APPENDIX v

III Curriculum Vitae

Tobias Weisenberger

Date/ Place of birth: 04.12.1979 in Emmendingen/ Baden-Württemberg/

Germany

since 2006 Dissertation at the Institute of Geosciences, Mineralogy

- Geochemistry, Albert-Ludwigs-University of Freiburg.

Subject of the thesis: “Zeolites in fissures of crystalline

basement rocks”. Adviser Prof. Dr. Kurt Bucher and

Prof. Dr. Reto Gieré

2000 – 2005 Geology (Diploma) study at the Albert-Ludwigs

University of Freiburg. Subject of the thesis: "Zeolite

facies mineralisation in the Hvalfjördur area, Iceland".

Adviser Dr. Rune Selbekk

1990 – 1999 Secondary school, Gymnasium Kenzingen, Germany

1986 – 1990 Primary school Endingen, Germany

Dissertation Tobias Weisenberger

Documents

Transcript of Dissertation Tobias Weisenberger