Kimberlites, Orangeites, and Related Rocks || Kimberlites and Orangeites

Working towards precision in nomenclatural problems is not simply play for bureaucratic minds but. on the contrary. it is a suitable exercise which should lead to less ambiguous communication.

Gian Gaspare Zujfa (1991)

KIMBERLITES AND ORANGEITES

The principal objective of this chapter is to compare, contrast, and illustrate the mineralogy and petrology of archetypal kimberlites, also known as group I kimberlites, with those of the group of diamond-bearing rocks which, in this work, are termed "orangeites." The latter rocks have previously been termed "micaceous kimberlites" or "group II kimberlites." The discussion, in conjunction with detailed mineralogical studies described in Chapter 2, will demonstrate conclusively that kimberlites and orangeites cannot be derived from the same parental magma and thus are not genetically related.

The second objective is to present suggestions for a revised textural-genetic classification of kimberlites and to show that this scheme is applicable to orangeites and melilitoids, although it should be clearly realized that none of these is cogenetic.

1.1. ETYMOLOGY OF GROUP I AND II KIMBERLITES

Diamonds derived from kimberlites were first found in South Africa in October 1869 on the farms Bultfontein and Dorstfontein (Dutoitspan) and then in July 1870 at Koffiefontein and Jagersfontein (Roberts 1976). They were found in muddy material, excavated from small quarries located adjacent to shallow water-filled depressions known in South Africa as "pans." It was not recognized at that time that the pans were the surface expression of cylindrical intrusions of igneous rock. Subsequently, the discovery of significant quantities of diamonds on the farm Vooruitzigt in May 1871 led to the discovery of three other diamond-rich deposits in the same area. Exploitation of these deposits resulted in the establishment of four major diamond mines. It was around these mines that the town of Kimberley was established. The story of these early discoveries and the development ofthe region is described in detail by Roberts (1976), Lenzen (1980), and Wilson (1982).

The diamonds occurring in the pans were initially considered to be of alluvial origin, but, as the deposits were excavated, it became clear that they were derived from a highly altered decomposed rock. This material, locally termed "yellow ground," was found to give way, with increasing depth, to fresher competent rock termed "blue ground." By 1872 it was recognized that the deposits were not alluvial in origin and occurred in cylindrical pipe-like structures. The blue ground was eventually recognized as an altered igneous rock and the primary source of the diamonds.

1 R. H. Mitchell, Kimberlites, Orangeites, and Related Rocks© Plenum Press, New York 1995

2 CHAPTER!

The diamond-bearing rocks were not given a petrographic name until 1887, when Henry Carvill Lewis, at a meeting of the British Association for the Advancement of Science in Manchester, stressed the unique character of the rock. Lewis (1887, 1888) described the rock as a porphyritic mica-bearing peridotite and recognized it as a type of volcanic breccia. Following the type locality nomenclature rules of the day, this rock was named, from its occurrence at Kimberley, kimberlite.

Many other occurrences of kimberlite were quickly discovered as knowledge of the geological character of the original deposits at Kimberley was disseminated. By the end of the nineteenth century kimberlites had been located throughout the Cape Province, the Orange Free State, and the Transvaal. Typically, prospectors referred to any igneous rock containing diamond as kimberlite. Curiously, geologists followed this practice, and identification of a rock as kimberlite came to be based more upon the presence of this trace accessory mineral than on the major mineral assemblage present. This practice survives to this day in many parts of the world. Thus, diamond-bearing olivine melilitites occurril)g on the eastern flanks of the Anabar Shield are termed "kimberlites" by Russian petrologists (Kornilova et al. 1983). Similarly, in the Arkhangelsk diamond province, rocks containing diamond are termed "kimberlite," whereas similar diamond-free rocks are known as "picrites" (V. Tretyachenko, pers. comm.).

The Diamond Fields of South Africa by Percy Wagner, published in 1914, was the first comprehensive summary of the occurrences of kimberlite in South Africa. This extremely influential work contained some important petrological observations, but also unfortunately set the stage for much of the confusion which was to follow concerning the nature of kimberlite.

Wagner (1914, p. 78) proposed that basaltic and micaceous or /amprophyric kimberlites could be recognized. The latter variety was further divided into subtypes based upon the presence or absence of augite (Wagner 1914, p. 107). Although Wagner noted substantial petrographic differences between the two major groups, it should be realized that his terminology is based primarily upon the macroscopic appearance of the rocks. The only common factor linking these petrographically disparate rocks is the presence of diamond and olivine macrocrysts.

Subsequently, Wagner (1928, p. 140) referred to the micaceous kimberlite constituting the Lion Hill Dyke (Orange Free State) as orangite (sic). As a footnote to this initial use of the term he stated: "This name will be proposed by the writer in a forthcoming publication to designate what has hitherto been known as mica-rich or lamprophyric kimberlite" (Wagner 1928, p. 148). Clearly, Wagner (1928) recognized the fundamental differences between his two varieties of "kimberlite" and believed there were sufficient grounds for the reclassification of one variety as a new rock type.

Unfortunately, Wagner died not long after the publication of the 1928 paper, and the promised article was either not written or never published. Consequently, the new name never entered the petrographic lexicon. Instead, Wagner's (1914) classification continued to be widely accepted, and until recently remained unchallenged, although minor refinements were made by some petrologists (Williams 1932, Bobrievich et al. 1959a,b, Milashev 1963, Dawson 1967, Frantsesson 1968, 1970, Vladimirov etal. 1981,1990, Kornilova et al. 1983).

KlMBERLITES AND ORANGElTES 3

In the 1970s, modem mineralogical studies (see below) led to revisions in kimberlite nomenclature and will ultimately, it is expected, lead to the demise of the Wagner (1914) classification. Unfortunately, this classification is still used indiscriminantly by petrologists and geochemists who are unfamiliar with recent revisions to kimberlite terminology. Consequently, any mica-rich kimberlite is commonly referred to as a micaceous kimberlite, with the implication that it is similar to Wagner's (1914) group of South African micaeous kimberlites.

The first major revision to Wagner's (1914) classification scheme was made by Mitchell (1970), who recommended that the term "basaltic kimberlite" be abandoned because kimberlites do not contain feldspar and are neither mineralogically nor genetically related to basalts. Mitchell (1970) proposed that three mineralogical varieties of kimberlite were recognizable on the basis of the dominance of olivine, phlogopite, and calcite in their modes. These were kimberlite (equivalent to Wagner's basaltic kimberlite), micaceous kimberlite (equivalent to Wagner's lamprophyric kimberlite), and calcite kimberlite. The latter variety was introduced as a new name in recognition of the presence of primary magmatic calcite in kimberlite. Prior to this all calcite in kimberlites was considered to be secondary in origin.

Skinner and Clement (1979) and Clement et al. (1984) subsequently devised a modal classification ofkimberlites which completely superseded Wagner's (1914) terminology. Their approach was to classify kimberlites on the basis of the primary groundmass modal mineralogy. The method is based upon the premise that the ubiquitous presence and relative abundance of olivine is oflimited use for classification purposes. This conclusion stems from the observation of Skinner and Clement (1979, p. 131) that it is difficult to determine the relative amounts of phenocrystal and xenocrystal olivine in kimberlites.

Skinner and Clement (1979) noted that diopside, monticellite, phlogopite, calcite, and serpentine are the five primary major groundmass constituents of the majority of kimberlites. Hence, they proposed five basic subdivisions of kimberlite, named after the groundmass mineral that is modally dominant. Olivine was considered to be ubiquitous, and, although in modal abundance it varies widely within and between kimberlites, its presence plays no role in the classification scheme for the reasons noted. In the Skinner and Clement (1979) classification, most of Wagner's (1914) and Mitchell's (1970) micaceous kimberlites are reclassified as phlogopite kimberlites.

The Skinner and Clement (1979) Classification has proven to be of great use because it permits the petrographic comparison of kimberlites from diverse localities. A modified version of the scheme forms the basis of the kimberlite classification used in this work (1.11 ). The classification may be utilized only when it has been determined that the sample being described is actually a kimberlite.

In 1983, Craig Smith, using samples selected on a petrographic basis by E. Michael W. Skinner (De Beers), demonstrated that monticellite calcite serpentine kimberlite and phlogopite kimberlite from the Kaapvaal craton (South Africa) possess distinctive Sr and Nd isotopic compositions. On the basis of these data, Smith (1983) suggested that monticellite serpentine calcite kimberlites, termed group I kimberlites, and phlogopite kimberlites, termed group II kimberlites, were derived from asthenospheric and lithospheric mantle sources respectively. It should be particularly noted that Smith (1983)

4 CHAFfERl

introduced the terms group I and group II as a means of classifying kimberlites primarily on an isotopic basis and not upon their petrographic character.

As a consequence of these isotopic studies, Smith et al. (1985b) and Skinner (1986, 1989) proposed that kimberlites can be divided into two distinct groups, I and II, on the basis of differences in their distribution patterns, age, petrography, content of mantle-derived xenocrysts, xenoliths and megacrysts, isotopic character, and whole rock geochemistry. It was noted by Skinner (1989) that petrographically and isotopically defined group II kimberlites are found only in South Africa, whereas group I kimberlites are found throughout the world. In southern Africa, group II kimberlites are typically older (Smith et al. 1985a,b, 1994) than most of the geographically associated group I kimberlites.

Interestingly, Skinner's (1989) group II kimberlites correspond to occurrences of micaceous kimberlites (sensu Wagner 1914). In retrospect, it is unfortunate that Wagner (1914) was so influenced by the presence of diamond as a means of identifying kimberlite that he did not initially follow existing petrographic practice and propose a type locality name for this petrographically distinctive suite of "kimberlites."

Subsequent to the recognition of group I and II kimberlites there has been an interest in characterizing the mineralogy and geochemistry of group II rocks (Fraser et al. 1985, Fraser 1987, Dawson 1987, Mitchell and Meyer 1989a, Skinner 1989, Tainton and Browning 1991, Tainton 1992, Fraser and Hawkesworth 1992, Skinner et al. 1994, Tainton and McKenzie 1994, Mitchell 1994a). From these studies it has become evident that group I and II "kimberlites" are mineralogically and geochemically quite distinct and that group II rocks have closer affinities to lamproites than to group I kimberlites. On the basis of this evidence, Mitchell (1991a,b, 1994a) has suggested thatthe rocks are derived from genetically distinct parental magmas, and group II rocks should not be regarded as a variety of kimberlite but as rocks belonging to an entirely different petrological lineage. If this conclusion is correct, group II rocks should not be designated as kimberlites. Thus, Mitchell (1989, 1991a, 1994a), Mitchell and Meyer (1989a), and Mitchell and Bergman (1991) proposed the revival of Wagner's (1928) term "orangite" as a potential name for these rocks. Following Wagner (1928), the name is given in recognition of their initial discovery in the Orange Free State of South Africa. If this proposal is accepted, then there exist three distinct major primary occurrences of diamond: kimberlite, orangeite, and olivine lamproite.

This chapter examines the petrographic grounds for recognizing the term "orangeite" as a useful rock name. Currently, few petrologists who actively study kimberlites and related rocks question that group I and II rocks are distinctive and have different origins. Resistance to eliminating the term group II kimberlite appears to be related more to preserving the status quo than to be based upon petrological argument and evidence.

The differences between group I and II "kimberlites" have been recognized by the International Union of Geological Sciences Subcommission on the Systematics of Igneous Rocks (Woolley et al. 1995). Although the Subcommission does not sanction the term "orangeite," it finds no compelling grounds to accept or reject the term until the rocks in question have been sufficiently characterized. Woolley et al. (1995) provide a definition of kimberlite (group I) based on Mitchell's (1986) definition and data obtained during the preparation of this work. A preliminary definition of the rocks currently known as "group II kimberlite," based upon data presented in this monograph, is also presented.

KIMBERLITES AND ORANGEITES 5

1.2. DEFINITIONS OF CRYPTOGENIC AND PRIMARY PHASES

In general, kimberlites and orangeites exhibit a distinctive inequigranular texture due to the presence of large rounded-to-anhedral crystals set in a finer-grained matrix. The origin of many of these crystals has not been satisfactorily determined. Some are without doubt xenocrysts, but others may be either phenocrysts or xenocrysts. In recognition of this ambiguity, Clement et al. (1984) and Mitchell (1986) have recommended that such cryptogenic pseudophenocrystal phases be referred to as megacrysts and macrocrysts, terms devoid of genetic inferences. In this work they are defined as follows:

Megacrysts are rounded-to-anhedral crystals greater than 1.0 cm in their maximum dimension. Megacrystal kimberlites 01' mega crystal orangeites are arbitrarily defined as containing greater than 5 vol % of such crystals.

Macrocrysts are rounded-to-anhedral crystals 0.5-10 mm in maximum dimension. Many macrocrysts are merely fragments of megacrysts. Macrocrystal kimberlites or macrocrystalorangeites are arbitrarily defined as containing greater than 5 vol % of these crystals.

Small (<0.5 mm, commonly 1-500 Ilm), anhedral crystals which are compositionally similar to megacrysts and macrocrysts may be found scattered throughout the groundmass. Such crystals are interpreted as fragments of disaggregated megacrysts and macrocrysts. To indicate their provenance and set them apart from bonafide groundmass phases, they are termed macrocrystal clasts or microcrysts.

Aphanitic kimberlites or aphanitic orangeites are varieties in which megacrysts and macrocrysts are absent, or present, only in small quantities (<5 vol %; Apter et al. 1984).

Crystals of subhedral-to-euhedral habit, considered to be early-forming primary liquidus phases, are termed phenocrysts (>0.5 mm) and microphenocrysts (0.1-0.5 mm). Phenocrystal kimberlites are rare. Aphanitic kimberlites containing abundant olivine microphenocrysts may be described as microporphyritic kimberlites. Orangeites are typically characterized by the presence of microphenocrystal phlogopite and may also be described as microporphyritic.

The term "porphyritic" should never be used to describe the characteristic inequigranular texture of macrocrystal kimberlite or orangeite. Much of the Russian literature which refers to porphyritic kimberlite (Milashev et al. 1963, Bobrievich et al. 1964, Komilova et al. 1983, Vladimirov et al. 1990) is in reality describing macrocrystal kimberlite.

Small (typically <0.1 mm) euhedral-to-anhedral primary minerals which constitute the bulk of the fine-grained groundmass are termed groundmass phases. These crystals may be set in a very fine-grained-to-optically unresolvable primary matrix or mesostasis. Note that the mineralogy of the mesostasis and the groundmass is different. The distinction is made to emphasize the conclusion, based on petrographic evidence, that crystallization of groundmass minerals typically ceases before formation of the mesostasis.

1.3. THE HYBRID NATURE OFKIMBERLITES AND ORANGEITES

Kimberlites and orangeites are petrographically complex rocks. Both are hybrids consisting of crystals originating from three distinct sources: mantle-derived xenoliths,

6 CHAPTER!

the Cr-poor megacrystlmacrocryst suite, and primary phases crystallizing from the magma. The relative contribution to the overall mineralogy of any given rock from each of these sources varies widely and accounts for the wide petrographic variation observed in kimberlites and orangeites. Note that the Cr-poor megacrystlmacrocryst suite is absent from most orangeites.

The principal mantle-derived xenoliths found in kimberlites and orangeites are garnet lherzolite, garnet harzburgite, chromite harzburgite, spinel lherzolite, websterite, eclogite and grosspydite, metasomatized peridotites (containing potassic richterite, phlogopite, and the titanates, yimengite, and hawthorneite), and the MARID (mica-amphibolerutile-ilmenite-diopside) suite of rocks. Detailed descriptions of the mineralogy of these rocks may be found in Sobolev (1977), Dawson (1980), and Nixon (1987).

Disaggregation of these xenoliths during incorporation into, and transportation by, kimberlite or orangeite magmas results in the addition of a wide variety of xenocrysts to the magma. The majority ofthese xenocrysts can be easily identified on the basis of their compositional equivalence with the minerals in the xenoliths, e.g., chrome diopside and Cr-pyrope derived from garnet lherzolite; subcalcic knorringitic garnets from garnet harzburgite; jadeitic pyroxenes, kyanite, and grossular-rich garnets from eclogites.

Other xenocrysts, in particular olivine and phlogopite, have compositions which are identical to those of minerals considered to have crystallized from the magma as primary phases. As there are no simple textural or compositional means of identifying these minerals as xenocrysts, they are commonly included in the macrocryst suite.

The Cr-poor megacrystlmacrocryst or discrete nodule suite consists principally of single crystals of magnesian ilmenite, Cr-poor titanian pyrope, Cr-poor subcalcic diopside, enstatite, phlogopite, and zircon. Coarse-grained lamellar intergrowths of ilmenite with clinopyroxene or orthopyroxene are common, as are small inclusions of one mineral within larger crystals of another. Megacrysts are common constituents of kimberlites but are rarely present in orangeites (Skinner 1989, Smith et al. 1985b).

The compositional variation and textural relations within the suite suggest that the megacrysts represent a series of crystals precipitating from a differentiating magma. The megacryst assemblage found within any given kimberlite is considered to be a hybrid formed by the mixing of crystals derived from several episodes of crystallization of the magma which was parental to the megacrysts. Megacrysts are not in eqUilibrium with their transporting magmas at low pressures (Shee 1984, Pasteris 1980).

Megacrysts are considered to be either xenocrysts, unrelated to kimberlite (Nixon and Boyd 1973, Pasteris et al. 1979, Hops et al. 1992), or cognate crystals formed in the upper mantle from kimberlite magma (Harte and Gurney 1981, Hunter and Taylor 1984, Mitchell 1986, 1987, Canil and Scarfe 1990). Reviews of the origin and nature of the Cr-poor megacryst assemblage are given by Mitchell (1986) and Schulze (1987).

Primary phases include phenocrysts, microphenocrysts, and crystals of subhedralto-euhedral habit which represent minerals crystallizing in situ to form the groundmass and mesostasis. The assemblage of these minerals together with their compositional variation are considered to reflect the character of their parental magma. By detailed studies of these assemblages it is possible to classify correctly rocks which, petrographically, are superficially similar but actually of diverse provenance.

KIMBERLITES AND ORANGElTES 7

1.4. pmLOSOPHY AND PRINCIPLES OF CLASSIFICATION

Igneous rocks are commonly classified using a combination of modal and textural criteria. While such methods are appropriate for common igneous rocks, they are considered by Mitchell and Bergman (1991), Scott Smith (1992), Mitchell (1994a), and Woolley et al. (1995) to be inappropriate for kimberlites, lamproites, orangeites, and lamprophyres. Many of these rocks cannot be unambiguously identified using standard modal-textural classifications, either in the field or the laboratory, as they are very similar in their macroscopic and petrographic appearance. Altered ultramafic alkaline rocks present particular challenges, and it is often extremely difficult, if not impossible, to classify correctly such rocks using simple petrographic criteria.

The similarity of modal and textural criteria has commonly resulted in rocks belonging to different magmatic series being incorrectly classified. Clearly, this situation is inappropriate for petrogenetic purposes, and modally based classifications of rocks have led to suggestions of consanguineity where none actually exist, e.g., the original description of olivine lamproites as kimberlites (McCulloch et al. 1983) and the classification of sanidine-rich lamproites as minettes (Middlemost et al. 1988).

Mitchell (1994a) has noted that a particular assemblage of minerals arises from the operation of petrogenetic processes and is not a fortuitous random association. In addition, the type and composition of minerals crystallizing from a given magma type must be controlled by, and reflect, the composition of that magma. Hence, it is suggested that ambiguities of nomenclature may be resolved by incorporating into classification schemes other significant descriptive criteria, i.e., mineral composition and accessory mineral assemblages, believed to be of petrogenetic significance. This suggestion stems from the belief of Mitchell (1994a) that classifications should have a genetic significance if progress is to be made in understanding the origins of, and relationships between, diverse ultramafic alkaline igneous rock suites.

1.4.1. Modal versus Genetic Classifications

Modal classifications are based upon the texture and volumetric percentages of the dominant, or major, minerals present in a rock. No consideration is given to the composition of the minerals, with the exception of the feldspars, and accessory phases are typically ignored. This approach to nomenclature has no genetic significance, and a given rock is named without consideration of the character of coexisting comagmatic rocks.

Genetic or mineralogical classifications use textural and modal information in conjunction with compositional data for some or all ofthe minerals present. Classification is based upon assemblages of typomorphic or characteristic minerals, some of which would be relegated to accessory status in purely modal classifications. This approach has genetic implications, in that rocks are assigned to a petrological clan or suite of consanguineous rocks of widely varying texture and modal mineralogy. The mineral compositional data permit discrimination between rocks which, on a modal petrographic basis, would be given the same name, but which are actually genetically different, having crystallized from different parental magmas.

Note that this type of genetic classification does not imply that we know how a particular magma type was formed or how individual rocks in a comagmatic series

8 CHAFfER 1

GENETIC MODAL GENETIC

Ti - PHLOGOPITE TO Ti - TETRAFERRI - + MICA (45) • AI - PHLOGOPITE TO Ti-BIOTITE PHLOGOPITE

( Fe ,K ) AI Si30e + ALKALI FELDSPAR (35) • (Na,KIAISi30e TITANIAN POTASSIUM RICHTERITE + AMPHIBOLE (4) • ARFVEDSONITE

PRIDERITE + OPAQUES (I) • ILMENITE

+ + + SAMDINE PHLOGOPITE LAMPROITE I I "MINETTE" I I MINETTE

Figure 1.1. Modal versus genetic classifications of a feldspar and mica-bearing hypabyssal rock.

originated. Classification is based upon directly observable characteristics of the mineral phases present and not upon inferred genetic criteria. Thus, the nomenclature is independent of revisions to, or modifications of, hypotheses advanced to explain the petrogenesis of a given comagmatic series.

The differences between modal and genetic classifications may be appreciated by considering the names given to a hypabyssal or effusive rock consisting of mica (45 vol %), alkali feldspar (35 vol %), clinopyroxene (10 vol %), and accessory minerals (5 vol %). In a modal classification this rock would be termed a "minette" on the basis of the dominance of mica and feldspar. However, Figure 1.1. demonstrates how the rock may be termed either a "sanidine phlogopite lamproite" or a "minette" when the composition of the minerals and the accessory phases are taken into consideration. Clearly, incorrect classification would have significant repercussions if the object of the exercise is for petrogenetic or exploration purposes.

For the identification of a given rock, genetic classifications of the type proposed here commonly require more information than is available from standard transmitted/reflected light studies. It is necessary to employ electron microbeam methods for the analysis and identification of minerals, as experience has shown that many typomorphic phases cannot be readily identified by thin-section petrography. The use ofbackscattered electron imagery coupled with energy dispersive X-ray spectrometric microanalysis (Mitchell 1995) is a prerequisite of these investigations. Use of electron microbeam techniques may be considered by some exploration geologists, or petrologists who do not specialize in alkaline rocks, to be a significant impediment to rock identification. However, these analytical techniques are now routine, and their application to the problem is mandatory if advances are to be made in the classification of alkaline rocks leading to a better understanding of their petrogenesis or economic potential.

1.4.2. Petrological Clans

The concept of a petrological clan was introduced by Scottish petrologists at the end of the nineteenth century and popularized by Reginald Daly in his influential text Igneous Rocks and Their Origin (Daly 1914). It was first realized at this time that differentiation and crystallization of particular magma types led to the formation of characteristic suites of rock types. Rocks which are so related form a consanguineous or comagmatic series.

In this work a petrological clan is regarded as a group of rocks derived from a particular type of parental magma which has been produced repeatedly in time and space.


Individual petrological provinces are composed of comagmatic rocks derived from specific batches of this magma type (Mitchell 1994a). Note that rocks of diverse modal character and chemical composition may be included in such a clan as they are genetically related.

The term "clan" as a genetic concept, as used above, follows the original (eleventh to twelfth century) form of the Scottish Highland clan system, in which membership of a clan was defined in terms of actual or purported descent from a common ancestor (Donaldson 1974, pp. 161-164, Oxford English Dictionary 1989, The Concise Scots Dictionary 1985). The word "clan" is derived from the Gaelic "clann," meaning children.

This definition of clan is different from that given in standard geological reference works, namely "a group of igneous rocks that are closely related in composition" (American Geological Institute Glossary of Geology, The Encyclopedia of Igneous and Metamorphic Petrology). This definition is particularly unsatisfactory because it has no petrogenetic significance and "closely related" is not defined. Rocks which are "closely related" in composition may be merely heteromorphs. Of greater significance is the fact that rocks belonging to a comagmatic series are certainly not closely related in composition, e.g., olivine lamproite and sanidine richterite lamproite are both members of the lamproite clan, yet they differ significantly in composition.

1.4.3. The Lamprophyre Clan Rock (1986,1990) has linked kimberlites (and orangeites), lamproites, and lampro

phyres (sensu lato) into a supergroup of rocks termed the "lamprophyre clan." Rock used the term to refer to a group of rocks he believed looked superficially similar, which are commonly associated in the field and have a number of petrological characteristics in common, e.g., richness in volatiles, porphyritic texture, occurrence as minor intrusions. This usage follows a looser and deri vative version of the original Scottish clan concept, namely that a "clan" is group united by a common trait and not related by blood (Donaldson 1974, Oxford English Dictionary 1989). Thus, Rock (1990) recognized that members of his "lamprophyre clan" have distinct origins.

Mitchell (1994a,c) has discussed Rock's lamprophyre clan concept at length and noted that it is a misnomer, since members of the clan are genetically unrelated. The clan merely unites rocks which have crystallized under volatile-rich conditions. Rocks belonging to different comagmatic series which are united by such a common characteristic are better considered as an example of the facies concept (Cas and Wright 1987). Consequently, Mitchell (1994c) proposes the recognition of a lamprophyre facies as a means of conveying the concept that some members of a petrological clan crystallized under different, i.e., volatile-rich, conditions than other members of that clan. Thus, Mitchell (l994c) recommends that the term "lamprophyre clan" be abandoned because there is no "lamprophyre magma type." Woolley et al. (1995) have also recommended that kimberlites and lamproites should not be regarded as members of a "lamprophyre clan."

Mitchell (l994c) recommends that the adjective "lamprophyric" be used to describe a facies of rocks derived from a particular parent magma. This usage retains the original meaning of "lamprophyric" as a description of observable characteristics of a particular group of rocks within a petrological clan which set them apart from other associated rocks.

10 CHAPTER 1

Recognition of a lamprophyre facies does not imply that we know how rocks belonging to the facies originated. However, recognition of the facies ultimately has genetic significance, as it serves to draw our attention to rocks which must have formed by processes specific to and/or different from other members of the comagmatic suite. Although it is possible to recognize a lamprophyric facies of the kimberlite, orangeite, melilitite, or lamproite clans, there are no sound petrogenetic reasons for gathering these rocks under a single petrological banner, as they may form by diverse processes in very different environments.

1.4.4. Mineralogical-Genetic Nomenclature within Petrological Clans

Mineralogical-genetic classifications of the type used in this work were first proposed for kimberlites by Skinner and Clement (1979). The methodology was subsequently applied to lamproites by Scott Smith and Skinner (1984), Mitchell (1985), and Mitchell and Bergman (1991). Practical experience in using this style of nomenclature by both academic and exploration petrologists during the last decade has demonstrated its effectiveness in the description of kimberlites and lamproites (Jaques et al. 1986, Mitchell 1986, Mitchell and Bergman 1991, Scott Smith 1992).

In mineralogical-genetic classifications, rocks are named on the basis of the nature of the parental magma of the clan. This magma type is stated in a root name (equivalent to a genus), e.g., kimberlite or lamproite. Individual rocks, or subdivisions of the clan, are described by mineral names (equivalent to a species) given as prefixes to the root name, e.g., leucite diopside lamproite. These compound names reflect the modal abundance of the major phases present. Prefixes are given in order of increasing modal abundance. Skinner and Clement (1979) recommended including in the name those minerals which are present in amounts exceeding two thirds of the volumetric abundance of the dominant mineral. Thus, a monticellite serpentine kimberlite would consist predominantly of serpentine in combination with monticellite amounting to more than two thirds of the serpentine abundance. Multiple prefixes may be added according to the limits defined by Skinner and Clement (1979). Strict application of this style of classification requires accurate determination of the mode. However, in practice, prefix names are more commonly given simply in order of increasing modal abundance. This approach reflects, in part, the fact that it is commonly very tedious, difficult, and time consuming to determine the modes of fine-grained altered ultramafic alkaline rocks.

It should be stressed that mineralogical nomenclature may only be" utilized when a rock has been identified as belonging to a particular petrological clan on the basis of the overall mineralogy and/or other criteria. Mineralogical-genetic classifications of some alkaline rocks have only become possible because of recent detailed petrological studies of type suites of rocks derived from a particular magma type, e.g., the lamproite clan (Scott Smith and Skinner 1984, Jaques et al. 1986, Mitchell and Bergman 1991). These studies have enabled petrologists to recognize the extent of modal and compositional variation in a given suite of consanguineous rocks. Prior to these studies it was difficult, if not impossible, to characterize many isolated samples of alkaline ultrabasic igneous rocks obtained during grassroots exploration or regional mapping programs.


In practice, mineralogical classification of a rock is usually based upon samples of hypabyssal facies material, as crystallization is usually sufficiently slow to allow development of typomorphic mineral assemblages. The classification of other textural varieties is far more difficult as rapidly quenched lapilli usually have not crystallized diagnostic groundmass minerals. In addition, crater and diatreme facies rocks, especially in tropical environments, are particularly prone to alteration and/or weathering.

The existing mineralogical-genetic classifications of kimberlites and lamproites have recently been endorsed by an lUGS Subcommission on the Systematics oflgneous Rocks (Woolley et al. 1995). Further discussion of the mineralogical nomenclature of orangeites and kimberlites may be found in Sections 1.10 and 1.11, respectively.

1.5. MINERALOGICAL COMPARISONS BETWEEN KIMBERLITES AND ORANGEITES

Mitchell (1986, 1991a, 1994a) and Mitchell and Bergman (1991) have suggested that the mineralogical and compositional differences between the rocks currently known as group I and II kimberlites, are so profound that they must represent rocks derived from different magma types. If this contention is correct, it follows that the rocks should possess readily identifiable petrographic and mineralogical characteristics permitting them to be classified according to the genetic principles described above.

Important mineralogical features of the rocks are compared in Table 1.1, from which it is apparent that they differ greatly with respect to their megacrystlmacrocryst and primary mineral assemblages. Table 1.1 is based upon previous mineralogical studies of archetypal kimberlites, summarized by Mitchell (1986) and Skinner (1989), and the new data for orangeites and kimberlites presented in Chapter 2 of this work.

Orangeites differ from kimberlites in that they

• Do not characteristically contain Cr-poor Ti pyrope, magnesian ilmenite, subcalcic diopside, and enstatite megacrystslmacrocrysts. Although some orangeites in the Prieska area 0.8.9) contain Mg-ilmenite and Cr-poor pyroxene macrocrysts (Skinner et al. 1994), and others contain Cr-poor Ti pyrope megacrysts (Moore and Gurney 1991, Bell and Rossman 1992), they are not abundant.

• Contain primary microphenocrystal and groundmass diopside (see 2.2). • Do not contain magnesian ulvospinel and spinels belonging to kimberlite spinel

evolutionary trend 1 (see 2.4.3). • Do not contain monticellite. • Are characterized by micas which evolve from phlogopite to tetraferriphlogopite.

Ba-rich aluminous micas belonging to the phlogopite-kinoshitalite solid solution series are absent.

• Contain K-Ba titanates (2.5) and zirconium silicates (2.12). • Rarely contain groundmass sanidine and potassium richterite. • Contain perovskites and apatites which are characteristically enriched in rare earth

elements and strontium relative to the compositions of these minerals in kimberlites (see 2.6, 2.7).

12 CHAPTER 1

Table 1.1. Comparison of the Mineralogy of Kimberlites and Orangeites

Olivine Macrocrysts

Phenocrys ts

Mica Macrocrysts Microphenocrysts Groundmass

Spinels

Monticellite

Diopside

Perovskite

Apatite

Carbonates

Serpentine

Sanidine K -richterite K-Ba hollandite K2Ti 130 27

Mn ilmenite Zr-silicates Leucite REE-phosphates

Barite Quartz Macrocryst suite

Kimberlite Orangeite

Abundant-principally xenocrysts Common to rare-principally xenocrysts

Common (FoS7_90) subhedral/euhedral Minor (F091-93)

subhedral/euhedraVdog's tooth

Minor, phlogopite-cryptogenic Rare, phlogopite Common, phlogopite-kinoshitalite

reticulate laths

Abundant, large (O.D1-O.1 mm) Typically Mg-chromite zoned to Mgulvospinel (Trend I). Atoll spinels very common. Trend 2 spinels rare, only in varieties with macrocrystal mica

Common, may be pseudomorphed by carbonate and/or serpentine

Primary diopside absent, may occur in contaminated groundmasses

Common, rounded to euhedral SrO- «I wt %) and (REE)zOrpoor «7 wt %)

Common to rare, euhedral prisms or acicular radiating aggregates in serpentine---calcite segregations. srO«1 wt %) and (REE}z0rpoor(<1 wt %).

Simple assemblages, common calcite, minor dolomite. Rare Sr-REE carbonates very evolved types

Abundant secondary and common primary in segregations

Absent Absent Very rare, only evolved types Absent Rare Very rare, only evolved types Absent Absent

Rare Absent Characteristic

Common, phlogopite cognate Common, phlogopite Common, phlogopite-

tetraferriphlogopite Poikilitic plates Minor to rare. small (<0.001-0.02 mm)

Euhedral Mg-chromite common, rarely zoned to Ti-magnetite (Trend 2). Atoll spinels rare. Mg-ulvospinels absent

Absent

Microphenocryst. Common to rare. Commonly resorbed. Zoned to Tiaegirine

Rare, subhedral to poikilitic srO- «1-6 wt %) and (REE}z03 rich (3-16 wt %)

Common euhedral prisms and poikilitic plates srO- (3-22 wt %) and (REEhOrrich (<1-10 wt %)

Common calcite, common Sr-Mn-Fe dolomites, minor witherite, ancylite, strontianite, norsethite

Common secondary

Rare groundmass Rare groundmass Common Common Common Common Rare pseudomorphs in poikilitic mica Minor monazite, daqingshanite, Sr-

REE phosphate Common Minor, groundmass Absent to extremely rare


• Contain a more varied assemblage of carbonates, including norsethite, strontianite, and witherite, than is found in kimberlites (2.13).

Significant petrographic differences are also apparent, e.g., the modes of orangeites are dominated by microphenocrystal and ground mass phlogopite, whereas phlogopiterich kimberlites are rare. Groundmass spinels and perovskites are typically less abundant and finer grained than spinels and perovskites in kimberlites (Skinner 1989). Illustrations of the petrographic characteristics of orangeites and kimberlites are given in Sections 1.1 0 and 1.11, respectively.

Geochemical differences (Chapter 3) include the distinctive isotopic compositions of the rocks (Smith 1983) and the elevated abundances of incompatible elements (Sr, Ba, REE, Zr) in orangeites relative to those of kimberlites.

Mitchell (1994a) has stated that, given the very different mineralogical character of micaceous (group II) kimberlites relative to archetypal monticellite-bearing primary diopside-free group I kimberlites, it is unlikely that, if this group of rocks were to be discovered today, they would be termed "kimberlites."

It is concluded, on the basis of the evidence presented in Table 1.1, that orangeites are not merely petrographic variants of archetypal kimberlites and thus must be derived from a distinct magma type. If the two rock groups are not genetically related there is no rational petrological reason for continuing to refer to group II rocks as kimberlites. Perpetuation of the name will only suggest false petrogenetic relationships between these rocks and bona fide kimberlites. Consequently, different names should be given to the two groups of "kimberlites" to reflect their fundamentally different origins.

Hence, Mitchell (1994a), following Wagner (1928), has proposed that the term "orangeite" be used to describe the rocks currently known as group II kimberlites, and the terms group I and II kimberlites be abandoned. Group I kimberlites are best referred to simply as "kimberlites." The group I prefix is not necessary as the clan name alone conveys all of the required nomenclatural (and genetic) information.

The introduction of a new name is justified in that it recognizes a newly identified, distinct magma type. Rocks crystallizing from this magma belong to the orangeite clan and may be described by compound names of the type described in Section 1.4.4, e.g., apatite orangeite, richterite orangeite. Orangeite is chosen as a name for the clan because the type localities of these rocks occur primarily within the Orange Free State of South Africa.

All of the mineralogical criteria required to identify orangeites and distinguish them from kimberlites, lamproites, and lamprophyres are readily observable by using a combination of petrographic and electron microbeam methods. Recognition is not based upon inferred or interpreted genetic criteria. However, the latter serve to confirm that orangeites and kimberlites are generated from different magma types whose sources are located at different depths within the mantle (see 3.8.1, 4.5, 4.6).

The value of the recognition of an orangeite clan may be appreciated by consideration of rocks recently described from southern Africa by Tainton (1992), Tainton and Browning (1991), Clarke et al. (1991), and Skinner et al. (1994). These studies showed that differentiation of certain group II kimberlites (Pniel, Sover North, Sweetput-Soutput, Besterskraal) leads to the formation of rocks containing groundmass potassium feldspar

14 CHAPl'ER 1

and potassium titanium richterite. These rocks might be termed "sanidine richterite lamproites" if they were classified in isolation from their consanguineous antecedents with which they are modally gradational. However, it is clear from field relationships and their mineralogy that they are genetically related to less-evolved olivine- and phlogopiterich rocks. Using existing terminology they could justifiably be termed "sanidine richterite group IT kimberlites." This designation, although meaningful to kimberlite petrologists who are aware of the mineralogical distinctions between the two groups, is guaranteed to breed confusion among nonspecialists, who might not realize that there are "kimberlites" and "kimberlites." Consequently, reference to these late differentiates as sanidine richterite orangeites makes clear the genetic distinctions between these rocks, kimberlites, and lamproites.

1.6. DEFINITIONS OF ORANGEITES AND KIMBERLITES

1.6.1. Orangeites

There is no previous detailed mineralogical definition of orangeite, as mineralogical characteristics of these rocks have been included in definitions ofkimberlites (sensu lato). A preliminary definition is given in Woolley et al. (1995). On the basis of the data presented here and in Chapters 2 and 3, the following definition is suggested:

Orangeites are a clan of ultrapotassic peralkaline volatile-rich (dominantly H20-rich) rocks, characterized by the presence of phlogopite macrocrysts and microphenocrysts, together with groundmass micas, which vary in composition from phlogopite to tetraferriphlogopite. Rounded olivine macrocrysts and euhedral primary olivines are common, but are not always major constituents. Characteristic primary groundmass phases include diopside, commonly zoned to, and mantled by, titanian aegirine, spinels ranging in composition from Mg-chromite to Ti-magnetite, Sr- and REE-rich perovskite, Sr-rich apatite, REE-rich phosphates (monazite, daqingshanite), potassian harlan titanates belonging to the hollandite group, potassium triskaidecatitanates, Nb-rutile, and Mn-ilmenite. These are set in a mesostasis which may contain calcite, dolomite, ancylite, and other rare earth carbonates, witherite, norsethite, and serpentine. Evolved members of the group contain groundmass sanidine and potassium richterite. Zirconium silicates (wadeite, zircon, kimzeyitic garnets, Ca-Zr silicate) are common as late-stage groundmass minerals. Quartz may occur rarely as a mesostasis mineral. Barite is a common secondary mineral.

Orangeites may be distinguished from kimberlites by the absence of monticellite, magnesian ulvospinel, and Ba-rich micas belonging to the barian phlogopite-kinoshitalite series. In addition, orangeites, in common with kimberlites and lamproites, do not contain melilite, alkali feldspar, plagioclase, kalsilite, or nepheline.

1.6.2. Kimberlites

Previous definitions of kimberlite (Clement et al. 1984, Mitchell 1979, 1986) included the mineralogical characteristics of orangeites. The definition presented below is based upon the definition of Mitchell (1986), which has been modified by the elimination of minerals characteristic of the orangeite clan and incorporation of recent mineralogical studies of groundmass micas in kimberlites (see 2.1.9). This definition of kimberlite has


been endorsed by the lUGS Subcommission on the Systematics of Igneous Rocks (Woolley et al. 1995).

Kimberlites are a group of volatile-rich (dominantly C02) potassic ultrabasic rocks commonly exhibiting a distinctive inequigranular texture resulting from the presence of macrocrysts (and in some instances megacrysts), set in a fine-grained matrix. The megalmacrocryst assemblage consists of anhedral crystals of olivine, magnesian ilmenite, Cr-poor titanian pyrope, diopside (commonly subcalcic), phlogopite, enstatite, and Ti-poor chromite. Olivine macrocrysts are a characteristic constituent in all but fractionated kimberlites. The matrix contains a second generation of primary euhedral-to-subhedral olivine which occurs together with one or more of the following primary minerals: monticellite, phlogopite, perovskite, spinel (magnesian ulvospinel-Mg-chromiteulvospinel-magnetite solid solutions), apatite, and serpentine. Many kimberlites contain late-stage poikilitic micas belonging to the barian phlogopite-kinoshitalite series. Nickeliferous sulfides and rutile are common accessory minerals. The replacement of earlierformed olivine, phlogopite, monticellite, and apatite by deuteric serpentine and calcite is common. Evolved members of the group may be poor in, or devoid of, macrocrysts and/or composed essentially of second-generation olivine, calcite, serpentine, and magnetite, together with minor phlogopite, apatite, and perovskite.

Kimberlites are best identified using the typomorphic assemblage of primary minerals referred to in the above definition. It is particularly important to make a distinction between cryptogenic, macrocrystal subcalcic diopside and primary groundmass diopside and note that kimberlites do not contain the latter. When present, groundmass diopside is a secondary phase, the crystallization of which is induced by the assimilation of siliceous xenoliths (Clement 1982, Scott Smith et al. 1984). Macrocrystal diopside is included in the definition in recognition of its common, but not characteristic, occurrence in kimberlites. Reference to this and other members of the macrocryst suite should be deleted from the definition if they are subsequently proven to be entirely of xenocrystal origin. For this reason bonafide xenocrysts such as diamond are not included in the definition.

Note that perovskites and apatites in kimberlites are poor in Sr and REE relative to the compositions of these minerals in orangeites. Potassium feldspar and potassium richterite are not found in the groundmass of kimberlites. The majority of kimberlites examined to date also lack the suite of complex K-Ba titanates, K-Zr silicates, and Sr-, Ba-, Mg-, and REE-rich carbonates which occur in the groundmass of orangeites. A similar accessory mineral assemblage has been found only in highly evolved calcite kimberlites, e.g., Benfontein, Wesselton sills (Mitchell 1994b). However, these kimberlites contain kinoshitalite and magnesian ulvospinel, i.e., typomorphic minerals which serve to distinguish them from orangeites.

Scott Smith (1989) has reported from kimberlite pipes 2 and 5 of the Andra Pradesh (India) province a Ca-aluminosilicate pseudomorphing a lath-shaped mineral which is considered to have been originally melilite. Murthy et at. (1994) have found a similar mineral in pipe 10 of this field and claim this to be fresh melilite; however, the composition is not in accord with that of gehlenite. These particular Andra Pradesh kimberlites also contain pectolite and have unusual olivine morphologies (Scott Smith 1989). Pseudoprimary pectolite in kimberlites elsewhere has been shown to result from contamination (Scott Smith et ai. 1983), a conclusion which suggests that if melilite were

16 CHAPTER 1

originally present its formation may also have been due to contamination. Although these occurrences suggest that melilite may occur in some kimberlites, it is concluded here that as yet there are no proven occurrences of melilite in kimberlites or orangeites.

1.7. AGE AND DISTRIBUTION OF ORANGEITES

Figure 1.2 (after Skinner 1989, Skinner et al. 1994) illustrates the age and distribution of orangeites and kimberlites in southern Africa. Skinner et al. (1992) have estimated that there are 229 occurrences of orangeite (termed group II kimberlites) as compared with 580 occurrences of archetypal (or group I) kimberlite. Figure 1.2 shows that most orangeites are distributed within a 400 x 1250 km belt trending southwest from Doko1-wayo (Swaziland) to Eendekuil, near Sutherland (Cape Province), most of the occurrences being within the Kaapvaal craton as clusters of intrusions in the southwestern part. Only in the Preiska area, located at the southwest margin of the craton, do a few atypical orangeites occur in an off-craton setting (Clarke et al. 1991, Skinner et al. 1994). In contrast, kimberlites are more widely distributed and characteristically occur in on- and off-craton settings.

o !

500 km ,....0500 ....... ·0 ........ -!

S:> 90 ~~~~ r' O~S /,/ ~ .

J ~50 .r'

h.. .r~ 0 I

i I 900,. Kaapvaal 1200 (fe I O~~D ~oo ~ ~.

i \ ~/-.' ........ ..."..) .~- ~ 1:5 e . i

nb A.e .20 115 (X) Craton . I oo~ ".145:J' r-'

(\'~ ..... _. __ .J1S i09"·~::';~ CO 100 120, ~o 0 F~tl90

It~ 90 .'" C3'c) .J

~137 e liO

SOUTH AFRICA o KIMBERLITES e ORANGE ITES

~ Craton Margin 165= Age in M.Y.

Figure 1.2. Distribution of orangeites and kimberlites in southern Africa relative to the Kaapvaal craton. Data points may represent clusters of several intrusions (after Skinner 1989).


Orangeites are unknown elsewhere in the world (Skinner 1989), whereas kimberlites, identical in geochemical and mineralogical character to archetypal southern African kimberlites, are commonly found on all continents. Rocks which have been ''termed micaceous kimberlites," e.g., the Zagodochnaya kimberlite (Yakutia) and the Tunraq kimberlite (Canada), are now considered in this work merely to represent kimberlites which have been modally enriched in macrocrystal micas. Such rocks have been termed "micaceous kimberlites" primarily on their macroscopic appearance. Unfortunately, without further detailed mineralogical and geochemical studies, this appellation is commonly taken to imply that the rocks are equivalent to the southern African orangeites.

Orangeites in southern Africa range in age from about 200 Ma to 110 Ma (Skinner 1989). Within the belt of orangeite intrusions, ages appear to decrease progressively from the northeast to the southwest (Figure 1.2). Skinner (1989) and Ie Roex (1986) consider this trend is suggestive of passage of the craton over a hot spot, although Mitchell (1986, this work) and Bailey (1993) disagree with this interpretation (see 4.4.3).

It is significant that the age trend is not well defined in the main clusters of Lower Cretaceous orangeites occurring in the southwestern part of the craton (Figure 1.2). Intrusions appear to have been emplaced contemporaneously over a period of 25 Ma in different fields rather than sequentially. These youngest occurrences include the mineralogically atypical orangeites found at the margins of the craton in the Preiska area (Skinner 1989, Skinner et al. 1994, Smith et al. 1994).

As an alternative explanation of the data it is suggested that there were two episodes of orangeite magmatism rather than one continuous period of activity. Thus, Lower Cretaceous (125-110 Ma) orangeites form a broadly contemporaneous group of intrusions, among which there are no obvious temporal and spatial correlations. The relatively few Upper Jurassic (165-145 Ma) intrusions in the central part of the craton may be regarded as separate events rather than representatives of the initial portions of a chronological trend. The 200-Ma Dokolwayo intrusion is temporally anomalous relative to all other orangeites in that it predates Karroo volcanism. The age and status of this intrusion as an orangeite require reinvestigation (see 1.8.7).

In contrast to orangeites, southern African kimberlites exhibit a very wide range in age and tectonic setting (Figure 1.2). While the majority of those in on-craton tectonic settings are 100-80 Ma, three other significantly older periods of kimberlite magmatism have been found: Kuruman 1600 Ma, Premier 1200 Ma, and Zimbabwe 500 Ma (Shee et al. 1989). Kimberlites emplaced in mobile belts surrounding the Kaapvaal craton give consistently younger ages, on the order of 54-72 Ma (Dawson 1989).

The Mayeng and Frank Smith kimberlites of the Bellsbank area have ages of 117 and 114 Ma (Apter et al. 1984, Smith et al. 1985a), respectively. These ages are similar to those of the Bellsbank and Newlands orangeite dikes, 119 Ma and 114 Ma, respectively (Smith et al. 1985a). Skinner (1989) has suggested that in cases where the differences in age between geographically closely related orangeites and kimberlites, e.g., Bellsbank, Preiska, are not great, the orangeites are less micaceous than orangeites from elsewhere, and the slightly younger kimberlites exhibit isotopic signatures transitional between those of orangeites and kimberlites.

18 CHAPTER 1

1.8. OCCURRENCES OF ORANGEITES

It is not the object of this work to provide a gazetteer or detailed descriptions of all occurrences of orangeite. Unfortunately, few modem textural-genetic studies of orangeite intrusions have been undertaken. Brief descriptions of the nine major districts (Figure 1.3) referred to in this work are given to illustrate the overall style of magmatism.

1.8.1. Finsch

The Finsch intrusion (Clement 1982), located near Postmasburg, Cape province, 170 km west of Kimberley (Figure 1.3), was discovered in 1960 and is the second largest pipe in South Africa after the Premier kimberlite. The surface expression is roughly circular with an area of 17.9 ha (Figure 1.4). The original area at the time of emplacement is estimated to be about 100 ha. From the current erosion surface to a depth of about 140 m the intrusion is emplaced in banded ironstones of the Asbesheuwels subgroup of the Precambrian Griqualand West Supergroup. These are underlain by 30 m of interbedded siliceous, dolomitic, and ferruginous sediments known as the Passage Beds. The lower

r· .. /

Kaapvaa/ ./

/' J Craton

GIBEON • . \ r· )/ (\ 13 3

I" . '\.. ..... .r-. '...J • PRETORP 5 NAMIBIA i L._ . .-" SWARTRUGG;NS 9 J~HAN~~~~~~:Y~'~

I 55 *-.;; ~'<; 0 \. 1'\ J r---q;~; /4 8 13· KROONSTAD

, ') POSTMASBURG.. Y? 0 0 I., ,,",-,~ . ../' I!INSCH .....,.." .BOSHOP.WINBURG

~KIMBE!!:P37 2.-' »-__ _.

PRIESK~O KO~FIEFONTE~ ..,.,/

BRITSTOWN .J 4~

VICTORIA WEST

SOUTH AFRICA

o 500 km

Figure 1.3. Distribution of orangeite fields and locations referred to in the text. Numbers adjacent to field boundaries indicate number of intrusions in the field (after Skinner et al. 1992),


100 m L..........J

Datum Surface

. ': :.',: .... ' .... " : .. ,": .

' ... '. : • : • II :: :1

t~-f~ I I .J.J

.c Satellite~ L Orangeite -.I "~ I 1 I

if;;um '% .

A N

1590,136m. AMSL

ORANGEITE

Figure 1.4. Size and shape of the Finsch diatreme (after Clement 1982).

19

part of the pipe (170-600 m) is emplaced in the Chuniespoort (Ghaa Plateau) dolomite. The Rb-Sr age of the intrusion is 118 rna (Smith 1983).

Figure 1.4 shows that the sides of the diatreme dip steeply at about 800 and the cross-sectional area decreases regularly with depth. Mining has not yet reached any complex root zone.

The pipe consists of eight discrete intrusions between the surface and 348-m level (Figure 1.5) of the open pit. These are numbered according to the sequence in which they were discovered. F1, a pelletal-textured volcaniclastic heterolithic orangeite breccia, is volumetrically the most important intrusion. Other significant intrusions are F2, F3, and F7, although their relative importance depends upon the level at which they are observed in the pipe. Thus, F7 is of minor importance between 196 and 220 m relative to F2, whereas the reverse is observed at 348 m. Clement (1982) has determined that the F6 dolerite-poor hypabyssal intrusion cuts the F5 dolerite-rich hypabyssal intrusion. Both intrusions are restricted in the open pit to a small projection outward from the pipe margin

20 CHAPTER I

F6

Figure 1.5. Composite plan illustrating the distribution of intrusions (FI-F8) and megaxenoliths (floating reef) in the Finsch diatreme between the 196- and 220-m levels of the open pit (after Clement 1982). Megaxenoliths consist of Ventersdorp lava (stippled) and Karoo sediments (black).

and have been truncated by the Fi volcaniclastic unit. F3 is also a marginal earlier hypabyssal intrusion whose age, relative to F5 and F6, is not known. F2, F4, F7, F7 A, F7B, and F8 all cut Fl and are internal hypabyssal intrusions. F2 is characterized by the presence of abundant globular segregations. F4, F7 A, and F7B are all thin curvilinear internal dikes. The hypabyssal units appear to have been emplaced in the sequence F7, F2, F4. F2 is younger than F8, but age relations between F7 and F8 are not known. From these data it appears that an early period of hypabyssal activity, i.e., embryonic pipe activity, was succeeded by diatreme-producing volcaniclastic activity, and that hypabyssal units subsequently intrude the diatreme facies rocks during the waning stages of activity. Given the increasing abundance of hypabyssal rocks with increasing depth, it would appear that the deep levels of the pipe will consist largely of a hypabyssal root zone-dike complex.

The Fi volcaniclastic unit contains many un metamorphosed xenoliths of Karroo age basalts, dolerites, shales, and sandstones, although no Karroo rocks are preserved in situ in the vicinity of the pipe. Extremely large xenoliths of basaltic lava, known locally as floating reefs, are concentrated near the margins of the pipes (Figure 1.5). These are also encountered in exploratory drill holes at depths of 600 m. Clement (1982) estimated that lava inclusions at Finsch must have sunk as much as 1000 m within the diatreme.

The Finsch intrusion is rich in diamonds and has an average recovery grade of about 1 metric carat/tonne (Shee etal. 1982). Clement (1982) noted that there is no variation in the diamond grade of the Fl unit from the surface to 348-m depth. The diamonds contain inclusions belonging primarily (86%) to the peridotitic suite (Gurney et al. 1979a). Mantle-derived ultramafic inclusions are not common at Finsch. Shee et al.


(1982) noted the presence of two diamondiferous peridotites in a suite of 80 xenoliths found at a single location in a minor body of kimberlite intersected at the 172-184-m level. The Fl unit is devoid of ultrabasic xenoliths, although high-Cr, Ca-poor knorringitic garnets occur in the heavy mineral concentrates produced during diamond extraction (Gurney and Switzer 1973). Megacrystal Ti-pyrope, subcalcic diopside, and Mg-ilmenite are not present at Finsch (Clement 1982).

Despite the size and economic importance of the Finsch intrusion there have been few published studies of the geology or petrology of the rocks. The weathered material occurring in the upper levels of the pipe is described by Ruotsala (1975). Descriptions of the morphology of the pipe and the various intrusions are given by Clement (1982). Geochemical and isotopic studies have been undertaken by Smith (1983), Smith et al. (1985b), Fraser (1987), Fraser and Hawkesworth (1992), and Kirkley et al. (1989). Prior to this work the only mineral compositional data were those obtained by Fraser (1987) in a limited study of olivine, spinel, and phlogopite.

1.8.2. Barkly West Region

A large number of dikes and small pipes occur about 100 km northwest of Kimberley in the district of Barkly West, Cape Province. The intrusions form a NNW-to-SSE-trending zone from Bellsbank through Sover and Newlands to Pniel (Figure 1.6).

1.8.2.1. Bellsbank

The Bellsbank orangeite field (Figure 1.7) consists of several subparallel dikes, known locally as fissures, striking N300E and approximately parallel to the escarpment of the Campbell Rand Plateau. Since their discovery in 1952, mining of portions of the Main and Bobbejaan Fissures has been undertaken intermittently. The Intermediate and Water Fissures have lower diamond grades and are infrequently mined. The grade of the West Fissures is unknown.

The dikes intrude the Campbell Rand dolomite and the underlying Black Reef Group quartzite and shale of the Griqualand West Supergroup. The orientation of the dikes is strongly controlled by the regional fracture pattern (Bosch 1971, Tainton 1992). The Rb-Sr phlogopite age of the dikes is 121 Ma (Smith et al. 1985a).

The Bellsbank Main Fissure is approximately 4.2 km in length and consists of a discontinuous series of en-echelon thin dikes «1 m), dipping 85° east (Bosch 1971, Tainton 1992). At the extremities of each segment of the dikes they decrease in width from a single dike to four or five veinlets before pinching out. Three small "blows" are developed on the Main Fissure system, the North, Middle, and South Blows, and a fourth (East Blow) is found 200 m to the east of the Middle Blow on the adjacent Kosmos Fissure. The North and Middle Blows contain macrocrystal orangeites which do not differ in petrographic character to the orangeites of the Main Fissure. Xenoliths of country rock are absent. The North Blow consists of four pipe-like swellings which narrow into a dike 0.5 m in width at a depth of 100 m. The South Blow orangeite contains abundant xenoliths of granite and schist derived from the underlying basement, in addition to rounded xenoliths of dolomite, but is in other respects similar to that of the Main dike. Karroo age xenoliths are not present, and there is no evidence in the xenolith suite to support the

22

Kalahari Supergroup

m Karoo Sequence

')~ ...... ,\ 1.1 t~ Postmasburg Group

~ Asbesheuwels Subgroup

~ Campbell Rand Subgroup

· .. • ' •• : Ventersdorp Supergroup · . .

CHAPTER!

20· r· ... ·,.·'\

20·----~-----'~·~_r--f_t_

CAPE TOWN

o I

loookm I

Figure 1.6. Geology of the Barkly West District and locations of orangeites discussed in the text (after Tainton 1992).

contentions ofTainton (1992) and Bosch (1971) that the South Blow represents a diatreme that extended to the original surface. All the blows are best regarded as hypabyssal facies root zones.

The Bobbejaan Fissure dips 85° east, is 2.3 km in length, and has an average width of 0.5-0.8 m but may be up to 1.2 m in places. Figure 1.8 illustrates the structure ofthe dike with respect to depth and shows that it consists of a series of en-echelon lenses. Compared with the Main Fissure, this dike is relatively continuous, and only a few lateral offsets of up to 15 m are found. Blows are not developed on this dike.


D 1m

Ghaap Group

Kalahari Supergroup

Calcrete and alluvium

-----Road

--Farm Boundary

! 1000 m!

C/) w a:: ::J C/)

~ IL.

l(/) w 3:

Figure 1.7. Distribution of dikes of orangeite in the Bellsbank field (after Bosch 1971, Tainton 1992).

23

The Intennediate, Water, and West Fissures are blowfree dikes similar to the Main and Bobbejaan Fissures. Their surface expression is marked by a discontinuous zone of calcrete, and little is known of their petrology.

The Main and Bobbejaan dikes are strongly weathered at their surface outcrops and where the dike is less than 0.6 m in width. Wide variations in color, from brown through green to gray along strike, reflect different intensity of alteration (Bosch 1971). Fresh

material obtained from deeper levels of the dikes consists of hypabyssal facies orangeite characterized by numerous large olivine macrocrysts set in a fine-grained black matrix consisting predominantly of microphenocrystal mica with minor amounts of primary olivine, spinel, apatite, serpentine, and dolomite. Flow differentiation has commonly

24

No.6 LEVEL

L- ~ No.1 LEVEL

&

No.7 LEVEL

& No.5 LEVEL __ B_

No.3 LEVEL

U/ SHAFT SECTION LOOKING

CHAPTER 1

A SECTION LOOKING WEST SURFACE B C SOUTH 0 ~--~~~~~~~~~~~~

I I

I I

I I

I I

I

/ /

I I

I

\... ,/E;'::",=.,"=/;;=:= No.1

......... U---:;' e NO.

2

---- ... ",- ......... , ./' 'I No.3 , \ \ , \ \ .I

/ \ \ ,/ \ " ."""," ,,"'" J,-----""~t~~=='= ... = No.4

.I~' I' /' , I

," / I I

/ / ==="¢!.:=:==== No.5

,,,,,.I " '.... -,,(

- ----, I I

o 50 m ... 1 __ ...................... 1

Figure 1.8. Plans and sections illustrating the distribution of en-echelon lenses of orangeite along a portion of the Bobbejaan dike. Bellsbank field (after Clement et al. 1973).

concentrated the larger macrocrysts at the centers of the dikes. The macrocrysts display undulose extinction and may have recrystallized to strain-free neoblasts.

Tainton (1992) noted that with increasing alteration the abundance of dolomite in the matrix of the rock increases relative to that of serpentine. In the Bobbejaan Fissure this replacement has gone to completion and only dolomite is present in the matrix.

The Intermediate Fissures consist of macrocrystal orangeites of similar character to those of the Main Fissure, but are poorer in spinel, apatite, and perovskite (Tainton 1992). The West Fissures are relatively poor in macrocrystal olivine, although complex parallel aggregates of olivine prisms are present in addition to subhedral phenocrysts. Olivines of

KIMBERLITES AND ORANGEITES 2S

this morphology are common in lamproites (Mitchell and Bergman 1991). The groundmass consists of diopside and phlogopite set in a matrix of serpentine. Tainton (1992) considered that the serpentine matrix was originally sanidine and that the West Fissures are more evolved than others at Bellsbank and are of "lamproitic affinity."

Mineralogical and geochemical studies of the Bellsbank orangeites may be found in Bosch (1971), Fesq etal. (1975), Kable etal. (1975), Boctor and Boyd (1982), Smith et al. (1985b), Kirkley et al. (1989), and Tainton (1992).

1.8.2.2. Sover Two kimberlite dikes occur on the farms Sover, Doornkloof, and Mitchemanskraal

(Figures 1.6 and 1.9). They are emplaced in the lower Karroo shales, which are underlain at depths greater than 100m by Ventersdorp lavas. The dikes, labeled A and B by Bosch

- Orangeite Dyke

OE;;;';:'I Dolerite} Karoo Sequence Shale

-------- Road --- Farm Boundary

A N

Figure 1.9. Distribution of dikes and intrusions in the Sover field (after Bosch 1971. Tainton 1992).

26 CHAPTER 1

(1971), strike NI5°E and N30oE, with lengths of 4 and 6.5 kIn, respectively. They appear to intersect between the Du Plessis No. 1 and No.2 mine shafts (Figure 1.9). However, Bosch (1971) claims that two preexisting intersecting fractures were filled with magma of different diamond content and the eastern dike B on the Sover farm is actually the northern extension of dike A.

The dikes, which are discontinuous lenses that regularly pinch out along strike, have an average width of about 0.8 m, reaching 1.5 m in places. Individual lenses may be slightly offset relative to each other (Figure 1.9), but the dikes do not form a well-defined en-echelon suite. Chilled margins are absent, and macrocrysts are concentrated by flow differentiation in the center of the dikes. Tainton (1992) suggests that in places the dikes are multiple, with the internal contacts being marked by veins of carbonate and serpentine.

Orangeites from this dike system exhibit a wide range in color, reflecting different degrees of alteration and/or carbonate content of the groundmass. Gray fresh rocks occur in the centers of the dikes and are isolated from the wall rocks by red-brown altered material. Country rock shale xenoliths are rare, and the xenolith suite consists primarily of rare lower crustal rocks and eclogites. Many of the latter are diamond bearing.

The intrusions are composed of macrocrystal hypabyssal orangeite. Microphenocrysts of phlogopite are set in a fine-grained groundmass of mica, spinel, perovskite, and apatite. The mesostasis consists of serpentine and dolomite.

The Sover North intrusion is an elliptical body (600 m2) located approximately 1.5 km NW of the Sover Mine (Tainton 1992). It intrudes Karroo shales which have been upturned at the margins of the intrusion. The morphology of the pipe is unknown. The main intrusion, termed SNI by Tainton (1992), consists principally of hypabyssal orangeite containing xenoliths of Ventersdorp andesite. A small intrusion consisting of xenolith-poor orangeite, termed SN2 by Tainton (1992), does not outcrop, lies at a depth of 30 m within the SNI breccia, and appears to have intruded the latter. All samples of Sover North material studied in this work are derived from the SN2 intrusion.

SN1 is hypabyssal facies lithic orangeite breccia. Xenoliths have reacted with magma, the andesites being replaced by phlogopite, amphibole, and clinopyroxene, the shales by sanidine and phlogopite. The rock consists of an assemblage of minerals (olivine, sanidine, diopside, perovskite, K-Ba titanates, rare richterite) and exhibits textures ("dog's tooth" olivines, poikilitic mica) similar to those occurring in madupitic olivine lamproites. Thus, Tainton (1992) classifies the rock as a "madupitic olivine lamproite."

The SN2 intrusion has a similar mineralogy but is relatively poor in olivine and rich in potassian richterite. Tainton (1992) describes SN2 as a "poorly macrocrystal Krichterite lamproite." Mineralogical differences between SNI and SN2 are considered by Tainton (1992) to be due to heteromorphism rather than different bulk composition. Tainton (1992) considers the main Sover dike system and the Sover North intrusion to be consanguineous. Thus, the Sover North intrusion is considered here to represent a differentiate of an orangeite magma. The mineralogy of Sover North is illustrated and discussed further in Section 1.10.

The Sover dike system has been investigated only by Bosch (1971) and Tainton (1992).

KlMBERLITES AND ORANGEITES 27

1.8.2.3. Newlands

The Newlands orangeites (Figure 1.6) occur as five small blows aligned over 600 m along a northeast-south west-trending series of fissures. The Rb-Sr age of the intrusion is 114 Ma (Smith et al. 1985a). The orangeite intrudes calcrete of the Kalahari Supergroup, which forms a veneer over Karroo shales and mudstones which are partially capped by the remnants of a diabase sill. The blows are numbered 1 to 5 from east to west. Blow 1, also termed the North Blow, consists of up to 90% shale xenoliths. Blow 2 is small and inadequately characterized due to poor exposure. Blow 3 contains many gneissic basement xenoliths. Blow 4 has a high content of lava xenoliths believed to be derived from the Karroo Stormberg series. These lavas are no longer preserved in the Karroo sequence in this region. Blow 5 (South Blow) is rich in xenoliths of shale. The fissure linking the blows is deeply weathered and replaced by calcrete. The orangeite filling the dike and forming the matrix of the blows is a macrocrystal hypabyssal orangeite. The petrographic character of the rock is illustrated in Section 1.10. The Newlands complex may be regarded as the lower levels of a diatreme root zone that is transitional to a feeder dike system.

Newlands orangeites contain a wide range of ultramafic and eclogitic xenoliths, including diamondiferous varieties. Xenocrysts of green knorringitic garnet are common at Newlands (Clarke and Carswell 1977, Schulze 1989).

Previously published studies of the geology and mineralogy of Newlands are those of Bonney (1899), Wagner (1914), and Tainton (1992). Isotopic studies have been undertaken by Smith et al. (1985a), Kirkley et al. (1989), and Tainton (1992).

1.8.2.4. Pniel

The Pniel, also known as Aaron's Prospect, intrusion was emplaced in Karroo shales (Figure 1.6). The geology of the occurrence is not well characterized as it is not of economic significance. Tainton (1992) states that it consists of a central shale-rich blow cut by numerous later dikes and sills. Wagner (1914) describes the occurrence as a complicated network of weathered dikes and sills which have indurated the shale.

The intrusion consists of macrocrystal diopside richterite orangeite. The presence of groundmass K-Ti richterite was originally noted by Erlank (1973), and the intrusion was considered to be a possible lamproite by Mitchell and Bergman (1991). Tainton (1992) describes the rock as an "amphibole-bearing phlogopite kimberlite," rather than a lamproite, because sanidine is not present in the rock. The petrography and mineralogy of the Pniel occurrence are discussed further in Sections 1.10 and 2.8.

1.8.3. BoshofDistrict

Numerous northeasterly trending dikes occur approximately 40 km to the east of the town of Boshof, Orange Free State (Figure 1.3). Most of the dikes appear to be uneconomic, although three significant diamond mines (Roberts Victor, New Elands, and Blaaubosch) have been developed at points where some of the dikes blossom into root-zone complexes. The dikes in the field are undoubtedly consanguineous as they yield similar Rb-Sr ages (Smith eta/. 1985a) of 127 ± 5Ma (New Elands), 128 ± 15Ma (Roberts Victor), and 133 ± 27Ma (Blaaubosch).

28 CHAPFER 1

Unfortunately, very little infonnation is available regarding the geology and petrology of the intrusive rocks found in this field, although there have been many detailed investigations of the eclogite xenoliths, which are particularly abundant at Roberts Victor.

1.8.3.1. Roberts Victor

The Roberts Victor intrusion (Clement et al. 1973) consists of two parallel fissures (Nos. 1 and 2) whose orientation has been controlled by the regional joint system. Number 1 fissure strikes 028° and dips 76-86° east. Three pipes and three blows have developed along No. 1 fissure. The fissures are covered by 12 m of Kalahari sand and are intrusive into Ecca shales and underlying Ventersdorp lavas. Wagner (1914) claims that the No.1 and No.3 pipes, which are separated by a distance of only 50 m at the surface, are different pipes developed upon the same fissure. Number 1 pipe contains xenoliths of Karroo shale and lavas and is devoid of xenoliths derived from units higher in the Karroo succession. In contrast, pipe No.3, in addition to Ecca shales, contains large xenoliths of Beaufort sandstone which must have been downrafted into the pipe. In contrast, Clement et al. (1973) contend that only one pipe is present as the intrusions apparently coalesce at depth. Insufficient petrological data are presented to assess the validity of this claim.

The matrix of both the fissure and the blows is a macrocrystal hypabyssal orangeite. Unfortunately many of the orangeites found within the intrusion are highly altered. Consequently, detailed mineralogical or geological studies of these rocks have not been undertaken. Some data may be found in C.B. Smith et al. (1985b),J.V. Smithet al. (1978), and Kirkley et al. (1989).

1.8.3.2. New Elands

The New Elands Mine is located on a blow developed where two dikes of orangeite intersect. Later intrusions of orangeite, relatively richer in mica with respect to the earlier dikes, cut this blow. The country rocks consist of Ecca shale and the main pipe contains xenoliths of these and downrafted Beaufort sandstone (Wagner 1914). The dikes consist of macrocryst-poor hypabyssal orangeite (see 1.10).

Detailed mineralogical studies of the New Elands orangeite have been presented by Mitchell and Meyer (1989a). Other mineralogical and geochemical data may be found in J.v. Smith et al. (1978) and C.B. Smith et al. (1985a,b).

1.8.4. Winburg District

Numerous east-west-striking dikes of orangeite occur in the Winburg district (Figure 1.3), but geological descriptions of the majority of these occurrences have not been published. A major mine, the Star Mine, is located upon a series of parallel dikes near Theron's Siding, about 15 km north ofTheunissen. Here several dikes (Star, Wynandsfontein, Burns) occur in a zone about 100 m in width and have been traced over 15 km along an east-west strike. A small pipe, the Phoenix (or Lion Hill) pipe, occurs at the western end of the Star Mine property on the Wynandsfontein dike. The dikes are typically <1 m in width and have intruded shales of the Lower Karroo Beaufort Series.

The dikes consist of macrocrystal hybabyssal orangeite (see 1.10) with each one differing subtly in petrographic character and mode. The only published detailed minera-


logical studies are by Mitchell and Meyer (1989a). Other mineralogical and geochemical information may be found in Wagner (1914) and Smith et al. (1978).

1.8.5. Kroonstad District

Two large pipes, Voorspoed and Lace (Crown), several smaller ones, and many small dikes and sills (Voorspoed, Rhenosterkop, Besterskraal, Normandiend) of orangeite have been identified in the Kroonstad district (Figure 1.3). Of the latter, the Besterskraal occurrence is particularly interesting in that it is a sanidine richterite orangeite (see Section 1.10). The age of the intrusions is on the order of 145 Ma (Skinner 1989). Modern geological and mineralogical investigations of this region have not been published.

Voorspoed

The Voorspoed Pipe occurs at the intersection of two fracture systems and crosscuts a northwest-striking antecedent orangeite dike. Clasts of this dike are found within the pipe, which is itself cut by a northeast-striking subsequent orangeite dike. The pipe and dikes are emplaced in Ecca Series shales and sandstones (Wagner 1914).

At the current level of erosion over 50% of the surface area of the pipe is occupied by a very large xenolith (6 ha) of Drakensberg lava. Figure 1.10 illustrates the size of this "floating reef' relative to the pipe. Because of the size and location of the xenolith, Wagner

o -30 m-

137 m-

340 m-

30m Level

A B G H

11~~~t\ ','."

Figure 1.10. Plan of, and vertical sections through, the Voorspoed diatreme, Kroonstad District, showing the downrafted megaxenolith (stippled) of Drakensberg lava (after Clement 1982).

30 CHAPl'ER 1

(1914) believed that the Voorspoed orangeite must have intruded a preexisting volcanic neck. However, recent drilling has demonstrated that the basaltic rocks form a real xenolith that persists to a depth of 300 m in the pipe (Figure 1.10; Clement 1982).

The orangeite in the pipe is a highly altered, relatively mica-poor volcaniclastic orangeite breccia. The antecedent dike is a hyabyssal diopside orangeite.

1.8.6. Swartruggens District

Orangeites, in the region of Swartruggens, Transvaal, occur as a series of subparallel, en-echelon dikes (Figures 1.3 and 1.11) emplaced in lavas of the Pretoria Series. The principal dike swarm has an overall strike trend of 110° over a distance of about 5 km. Several small outcrops of orangeite, found along the Elands River to the north of the main dike system, suggest the existence of other dike swamis in the area.

Within the main dike swarm individual dikes have a sinuous outcrop and may strike between 100° and 120°. The thickness ranges within short distances, from 2 cm to 2 m. Many are composite, with individual members being as thin as 0.5 cm. It is impossible to trace anyone dike as a continuous unit along strike throughout the whole system. Dikes encountered in the mines commonly do not outcrop. Indi vidual dike segments differ with respect to their degree of alteration and petrographic character. Thus, the Changehouse dike is highly altered, whereas the Main dike is relatively fresh. Individual dikes encountered at different levels in the mines are petrographically different as a consequence of the complexity resulting from multiple intrusion. Hence, each dike is not a single petrographic unit formed by the crystallization of a single batch of magma. The dike system appears to have a character similar to that of Bellsbank and is thus considered to be a discontinuous series of en-echelon lensoid bodies.

Individual dike segments have been named primarily after their discoverers with some apparently renamed by various mining companies. Consequently, dike nomencIa-

NOOITGEDACHT 405

o I

I km I

-5

A N

Figure 1.11. Distribution of orangeite dikes in the vicinity of Swartruggens. Transvaal (after Fourie 1958).


ture is confusing. All Swartruggens samples studied in this work are derived from different mining levels intersecting the Main dike occurring in the Helam Mine.

The Swartruggens dikes vary with respect to their diamond content and petrographic character. All diamond-bearing dikes are classified here as macrocryst-poor hypabyssal orangeites. Other diamond-free dikes are olivine-bearing minettes. The principal occurrence ofthe latter is known as the Male lamprophyre (N.B. "Male" [mah-lay] in Afrikaans means barren). This rock consists of serpentine pseudomorphs after oli vine set in a matrix of pyroxene, mica, and potassium feldspar. The compositions of the pyroxenes (2.2.3) are not significantly different to those of pyroxenes in the Swartruggens orangeites. However, the micas have different compositions and evolutionary trends from those of orangeite mica, suggesting that these lamprophyres have no simple genetic relationship to orangeites.

It was recognized in earlier studies (Fourie 1958) that Swartruggens orangeites differ from those in other districts in that they contain considerable amounts of quartz. Interestingly, quartz veins occur parallel to the strike of the dike swarm. Fourie (1958) states that these veins are post-dike in age, but this conclusion does not rule out their being formed from hydrothermal residual fluids. Alternatively, the quartz may be simply of secondary origin.

Age determinations have not established the relationship between the orangeite and lamprophyre dikes. Allsopp and Barrett (1975) give a Rb-Sr mica isochron age of 147 ± 4 Ma for micas from the Main dike. However, this age is unreliable as it is essentially a two-point isochron. These data, together with others for micas from the Main dike, were used by Smith et al. (1985a) to construct a Rb-Sr errorchron that gave an age of 156 ± 13 Ma. This errorchron included one sample from the Male dike. As the datum lay within the limits of the errorchron, it may be concluded that the Male lamprophyre dike may be of similar age to the orangeites. However, the extremely different mineralogy and petrographic character of the two dikes suggest that they are not consanguineous. It is suggested here that they are merely contemporaneous and have exploited the same fracture system during emplacement.

Previous studies of the geology of the S wartruggens dikes are by Haughton (1935) and Fourie (1958). Mineralogical and geochemical data are given by Mitchell and Brunfelt (1975), Allsopp and Barrett (1975), J.Y. Smith et al. (1978), C.B. Smith et at.

(1985a,b), Skinner and Scott (1979), and Kirkley et at. (1989).

1.8.7. Dokolwayo

The Dokolwayo diatreme, Swaziland (Figure 1.3), considered to be a "group II kimberlite" by Skinner (1989), occurs as a single intrusion emplaced in late Archean granite-gneiss country rocks. It lies outside the present limits of the Karroo basin but contains downrafted xenoliths of Karroo age. No accurate radiometric age determinations for the intrusion are available. Skinner (1989) gives an age of 200 Ma (method unspecified), and Hawthorne et al. (1979) suggest an age of 190 Ma, based upon field relations and xenolith provenance. No other diamondiferous rocks, apart from a few minor dikes, occur in the area (Hawthorne et al. 1979).

32 CHAPTERl

Hawthorne et al. (1979) have published a general description of the Dokolwayo diatreme and the nearby Hlane alluvial diamond deposit. The diatreme is an elongate structure (600 m x 50--100 m) with a northeasterly strike. The area increases from 2.8 ha at the surface to 3.8 ha at a depth of 120 m, narrowing again at greater depths. The upper-level bulge is capped on the northern flank by breccias lacking an igneous matrix, which are interpreted (this work) to be contact breccias (Clement 1982). The central parts, and the bulk, of the intrusion consist of several varieties of diatreme facies volcaniclastic heterolithic breccias (this work). Some units with extremely high contents of xenocrystal quartz are described as crystallinoclastic breccias (Clement 1982). Xenoliths range in size from> 1 m to microscopic clasts and consist primarily of sandstones and coals of middle Ecca (Lower Permian) age together with locally derived Archean granite-gneisses. Xenoliths of crater facies resedimented volcaniclastic (epiclastic) material are present in one of the diatreme-facies volcaniclastic units (this work). Pelletal-textured rocks are absent, and the intrusion is considered here to have been eroded to depths at which the diatreme proper enters the root zone level.

Altered hypabyssal rocks are found at the margins of the intrusion. Their textures have been described as "porphyritic" by Hawthorne et al. (1979). They consist of altered subhedral olivines set in a groundmass of spinel, phlogopite, serpentine, and calcite. Groundmass micas occur as fine-grained «0.1-0.3 mm), irregular partly-altered plates and laths. Euhedral grains are present locally, and much of the mica has been altered to serpentine. Because of the high modal amounts (30 vol %) of groundmass mica, Hawthorne et al. (1979) term the rock a "phlogopite kimberlite" (sic), using the mineralogical classification of Skinner and Clement (1979).

Examination of the altered hypabyssal facies rock during the preparation of this work showed that the groundmass micas are colorless phlogopites which are commonly intergrown in a reticulate pattern of interlocking laths. The texture is reminiscent of that observed in many archetypal kimberlites (Figure 1.74). Detailed mineralogical studies of the rocks have not been undertaken; however, on a petrographic basis, there are no grounds for considering them to be orangeites, as the micas do not have the textural or optical characteristics of orangeite micas. Thus, the petrological status of the Dokolwayo intrusion requires reexamination.

1.S.S. Prieska District

The geologically complex southwestern margin of the Kaapvaal craton in the neighborhood of Prieska (Figure 1.3) has 130 known intrusions of kimberlite and orangeite (Skinner et al. 1994). The region represents an area where a geographical association of temporally distinct kimberlite and orangeite magmatism can be found.

Skinner et al. (1994) have divided the region into five domains (Figure 1.12) on the basis of the disposition of the intrusions, relative to the craton margin, major faults, and petrographic and isotopic character of the rocks. Each domain is characterized by intrusions of predominantly one petrological and isotopic type. Thus, intrusions in domains I and III are predominantly "group 2 kimberlites" (Le., orangeites), whereas those in domains II and IV are "group 1 kimberlites" (Le., archetypal kimberlites). Intrusions in domain V have mineralogical and isotopic characteristics considered by

KlMBERLITES AND ORANGEITES

~26 KIO(l14) KII

.K32

• DOMAIN

DO O~ .. K5(126)

i1t(9, •• ~ ~t9G' , •

\ ~~ K27 •• Prieska , ~ DOMAIN III " ij\ 1": +6'> .......... K25

o 6/" ~ I/~ .,

A N

• •• //.9 ~ ......... K56K20• ~ ~! • -...K24 ..... ? .I'K46 \.. • ... ~K52 K25 .. aoK45 KI5.' ,. t • \ K41 0" 0 K42

-300

(127),. KI9 (liS) .' K40/ .. ••• K9, • .1 (114) ..

KI30 :.. K64 •• I 0 ...... DOMAIN II • .. 1 oo~..;,;.;.:...:..;.;,.;.....:.:..

K690 K66~K6 ,~" t .... K2(IOI) ;-'O--' .. K65 K35(74)

o KI (71) ......... Q, 0 K22 (S2) .........

El it! 0" 0 DOMAIN IV ,-31 0

K25 0 0 , *KI6(I03) .0 ..........

KI5 KI(74)~ 0

o 0K5 0 '/--------- - K23 . *KII K9**. ""?

K19* DOMAIN V .

o Kimberlite • Orangeite .. Unclassified

* Isotopically Transitional

o !

50km ! ~ ", ..

Major fault Inferred craton

boundary boundaries ~,,-,' Domain ,

33

Figure 1.12. Tectonic domains and the distribution of orangeites and kimberlites in the Prieska District, Cape Province (after Skinner et al. 1994).

Skinner et al. (1994) to be transitional between "group 1 and 2 kimberlites" (see below and 3.8.1). Figure 1.12 demonstrates that the orangeite intrusions occur primarily in an on-craton tectonic setting, whereas the kimberlites straddle the craton margin.

The intrusions of domains I and III are predominantly hypabyssal dikes which are commonly extensively altered. Skinner et al. (1994) have described them by optical petrographic studies, but detailed mineralogical investigation by electron microbeam

34 CHAPTERl

methods has not yet been undertaken. Only a few of the intrusions are petrographically similar to bona fide on-craton orangeites such as occur at Barkly West, the olivine macrocryst-rich intrusions of Sandrift (25/K5) and Omdraaisvlei (26/K25) being the best examples.

Many of the rocks contain groundmass spinels and perovskites which are relatively coarse-grained compared to those occurring in other on-craton orangeites. Many also contain significant amounts ofxenocrystal ilmenite (Albertshop 25K125 , Markt 26/K15, Welgevonden 26/K67, Silvery Home 27/Kl1), chrome-poor garnet macrocrysts (Middelwater-Ol 26/K19, Wielpan 26/K52, Nietgedacht 261K65), Ti02-rich garnet macrocrysts (Middelwater-Ol 26/K19, Kafferskolk 26/K20, Melton Wold 27/K9), and Cr-poor clinopyroxene macrocrysts (Wielpan South 26/K60, Silvery Home 271K 11). Although the presence of these macrocrysts is atypical of orangeites, Skinner et al. (1994) nevertheless consider their host intrusions to be petrographically and isotopically "group 2 kimberlites," i.e., orangeites.

Skinner etat. (1994) noted that many of the rocks exhibit petrographic features similar to those observed in olivine lamproites. These include the common late crystallization of amphibole and sanidine, and the presence of madupitic or oikocrystal coarse-grained, orange-brown, strongly pleochroic ground mass phlogopite and glomeroporphyritic aggregates of olivine. Rocks exhibiting such features are termed "evolved group 2 kimberlite" by Skinner et at. (1994), who also note the similarity of these rocks to amphibole- and sanidine-bearing orangeites in the Barkly West and Kroonstad Districts. Note that the Prieska District rocks could be termed "lamproites" using the terminology of Tainton (1992) and Tainton and Browning (1991). Describing the rocks as "orangeites" eliminates the previous ambiguities in nomenclature (see also 1.5).

Rocks occurring in domain V contain unusually coarse-grained and abundant clinopyroxene and perovskite in addition to having unusual isotopic signatures (3.8.1). They are petrographically unlike typical on-craton orangeites, although they contain abundant groundmass mica. The presence of primary amphibole and possible sanidine is considered by Skinner et al. (1994) as sufficient grounds to justify their designation as varieties of "group 2 kimberlite" (orangeite).

Archetypal kimberlites occurring in domains II and IV are no different in their mineralogy, macrocryst suites, and isotopic character from kimberlites occurring elsewhere in southern Africa in on- and off-craton tectonic settings.

Skinner et at. (1994) and Smith et at. (1994) have shown that the orangeites of domains I and II range in age from 114 to 127 Ma and are thus apparently contemporaneous with the main body of orangeite magmatism in the central parts of the craton. Intrusions in domain V are significantly older (136-167 Ma), although geochronological data are less precise. The kimberlites are all younger than 114 Ma (Skinner et at. 1994). Although some (Witberg 251KI0, domain I) are slightly older than the main period of contemporaneous kimberlitic on-craton magmatism (90 Ma) in southern Africa, the majority appear to have relatively young «80 Ma) ages, similar to those of other off-craton kimberlites emplaced in mobile belts.

In this work, the orangeites emplaced in domains I, II, and III, together with the isolated intrusion (Mount Pierre 27/K15) in domain IV, are considered to compose a single


consanguineous field. The intrusions in this field were derived from magmas which were apparently more evolved than those in the Barkly West and Boshof Districts. Further it is suggested that intrusions occurring in domain V originated during a distinctly different and older magmatic event and were not derived from orangeite parental magmas. They may represent archetypal kimberlite magmas which have been contaminated by melts derived from ancient enriched lithospheric mantle.

1.8.9. Summary

It is concluded from the above descriptions of several fields of orangeite magmatism that

1. Orangeites, in common with kimberlites and melilitoids, form bonafide diatremes. They do not form the wide, shallow vents characteristic of lamproitic magmatism. Diatreme formation is preceded by the crystallization of hypabyssal rocks which are subsequently incorporated into the volaniclastic rocks of the diatreme. Magma intruded into the lower parts of the diatreme subsequently crystallizes as a second generation of late-stage hypabyssal rocks.

2. Orangeite magmas ascend to high crustal levels as dikes utilizing preexisting fracture systems. Diatremes and blows of hypabyssal orangeite are developed above dike systems.

3. Many orangeites occur as en-echelon dike swarms, a style of intrusion not typical of kimberlites but characteristic of many suites of lamprophyres (Currie and Ferguson 1970). Composite intrusions are common.

4. Orangeite magmas differentiate to residual magmas which crystallize amphibole, sanidine, and possibly leucite (see 1.10).

5. Individual fields of orangeite differ with respect to their mineralogy and degree of evolution. Thus, the Barkly West field is characterized by olivine macrocrystal orangeites, whereas other fields are relatively poor in olivine macrocrysts. The Kroonstad and Prieska fields are relatively evolved compared to the Barkly West, Boshof, and Winburg fields. The Swartruggens orangeites are atypical in containing quartz and many zirconium-based minerals (see 2.12).

6. Lavas and plutonic rocks belonging to the orangeite clan have not been recognized. Clasts consisting of pyroxene and phlogopite or phlogopite alone have been recognized in several orangeites (Skinner 1989, this work). Dawson and Smith (1977) noted that xenoliths consisting of mica, amphibole, and diopside occur at Newlands but do not appear to be MARID xenoliths (Skinner 1989). They may represent disrupted cumulates formed in the magma prior to intrusion.

1.9. TEXTURAL-GENETIC CLASSIFICATIONS OF PETROLOGICAL CLANS

The extensive mineralogical and textural variations exhibited by kimberlites have given rise to a plethora of textural classifications (see Wagner 1914, Kovalskii 1963, Rabkin et al. 1961, Artsybasheva et al. 1964, Frantsesson 1968, 1970, Kornilova et al. 1983, Clement and Skinner 1985, Vladimirov et al. 1981, 1990). These classifications

36 CHAPTER 1

VOLCANICLASTICS

Figure 1.13. Model of an idealized kimberlite magmatic system. illustrating the relationships between crater diatreme and hypabyssal facies rocks (not to scale). Hypabyssal rocks include sills. dikes. root zone. and "blow" (after Mitchell 1986).

are commonly mutually exclusive, making comparison of kimberlites derived from different provinces difficult or impossible. Russian petrologists in particular have taken the erection of independent classifications to an extreme, with some groups devising different classification schemes for the same kimberlites.

Recognizing these problems, Clement and Skinner (1979, 1985) and Clement (1982) attempted to bring order to the chaos by devising a textural-genetic classification scheme based upon the concept that different kimberlite facies may be recognized. Kimberlites are considered to be a clan of rocks which formed from a volatile-rich magma. This magma, in common with others, undergoes differentiation and crystallization in diverse temperature and pressure regimes, giving rise to a spectrum oftexturally and petrographically different rocks. Classification is based on the premise that macroscopic and microscopic characteristics observed in textural variants of kimberlites are mainly related to different near-surface emplacement and crystallization processes. Thus, Clement and


Skinner (1985) recognized three textural genetic groups ofkimberlite associated with a particular style of magmatic activity, namely crater, diatreme, and hypabyssal facies rocks.

In principle, the facies approach to classification of Cas and Wright (1987) and Clement and Skinner (1985) is applicable to any other magma type. Thus, Mitchell and Bergman (1991) have devised a similar textural-genetic classification for the lamproite clan which recognizes lava, crater and pyroclastic, hypabyssal, and plutonic facies. Analogous schemes could be devised for the members of the melilitoid clan, ultrapotassic rocks of the kamafugite series and some minettes.

The revised textural-genetic classification of kimberlites presented below is a development of the Clement and Skinner (1985) model that is also applicable to orangeites and other diatreme-forming alkaline ultrabasic igneous rocks.

1.9.1. Kimberlites

Figure 1.13 illustrates an idealized model of a kimberlitic magmatic system (Mitchell 1986) based on observations of southern African kimberlites (Mannard 1962, Dawson 1971, Hawthorne 1975, Clement 1982), showing the spatial relationships between the different facies.

1.9.1.1. Crater Facies

Crater facies rocks are divided into lavas, pyroclastic rocks, and resedimented volcaniclastic rocks (Figure 1.14; Mitchell 1986). The latter have previously been termed "epiclastic kimberlites" (see below).

1.9.1.1.a. Lavas. Kimberlite lavas have not yet been conclusively identified. The only potential candidates are a series of small lava flows comprising the Igwisi Hills, Tanzania (Dawson 1994). Lavas may be described by standard terminology.

VOLCANICLASTIC ROCKS

METACHRONOUS VOLCANOGENIC

SEDIMENTARY DEPOSITS

KIMBERLITE LAVA

PYROCLASTIC KIMBERLITE

RESEOIMENTED VOLCANICLASTIC

KIMBERLITE

STANDARD TEXTURAL DESCRIPTIONS

STANDARD PARTICLE SIZE AND TEXTURAL TERMINOLOGY

STANDARD PARTICLE SIZE AND TEXTURAL

TERMINOLOGY

PSEUDO-CRATER FACIES ROCKS CONSISTING OF TERRUGINOUS COMPONENTS AND REWORKED VOLCANICLASTIC MATERIAL FILLING EROSIONAL DEPRESSIDNS AND DEPOSITED UPON ANCIENT VOLCANIC DIATREMES OR THE HYPABYSSAL ROOT ZONES OF OIATREMES ARE NOT CRATER FACIES DEPOSITS.

Figure 1.14. Textural genetic classification of crater facies kimberlites.

38 CHAPTERl

1.9.1.I.b. Pyroclastic Rocks. In situ examples of pyroclastic kimberlites are rare because the craters of most kimberlitic volcanoes have been destroyed by erosion. Unfortunately, there are no published modern descriptions of the pyroclastic kimberlites known to occur in Tanzania and Botswana (Mannard 1962, Hawthorne 1975). It is to be expected that subaerial and subaqueous tephra deposits, together with resedimented volcaniclastic deposits (see below), will eventually be described from Botswana crater facies kimberlites.

Standard pyroclastic terminology (Fisher and Schminke 1984, Cas and Wright 1987) may be applied to pyroclastic kimberlites. Following Fisher and Schmincke (1984), pyroclastic fragments are considered to be produced by processes associated with volcanic eruptions. They represent particles expelled through volcanic vents without reference to the causes of eruption or origin of the particles. Pyroclastic deposits may be interpreted in terms of standard volcanic facies models (Cas and Wright 1987).

Crater facies may also contain hydroclastic rocks, such as base surge deposits, resulting from phreatomagmatic activity. These are most likely to be found outside the crater, as base surges within craters are primarily erosional rather than depositional. Hydroclastic is used, following Fisher and Schminke (1984), to describe the products of explosions due to steam derived from water of any kind. Hydroclastic crater facies rocks are different in character and origin to hydroclastic diatreme facies rocks (see below).

1.9.1.I.c. Resedimented Volcaniclastic (Epiclastic) Rocks. Current usage of the term "epiclastic" in kimberlite petrology (and elsewhere) is unsatisfactory. The term has traditionally been used by sedimentologists to describe igneous material weathered and eroded from older source rocks. Thus, epiclastic rocks are commonly considered to be sedimentary units with clasts of lithified pyroclastic or volcanic rock. Recently, volcanologists have extended the term to include material derived from contemporaneous and unconsolidated sources with or without lithic (terrigenous) components. Kimberlite petrologists have typically used the term "epiclastic" in this context rather that in the older sedimentological sense. Unfortunately, there is no agreement among volcanologists as to a definition of "epiclastic" (see Fisher and Schminke 1984, p. 89, Cas and Wright 1987, pp. 8-9, Cas 1991, Allen 1991, McPhie et al. 1993). A significant further problem with the existing terminology is that it gives no indication as to when a particular "epiclastic" deposit was formed relative to the age of the juvenile clasts.

Recently, McPhie et al. (1993) have proposed the elimination of the term as a consensus has not yet been agreed upon regarding a redefinition. Thus, "epiclastic rocks" derived from contemporaneous pyroclastic rocks are renamed "resedimented volcaniclastic deposits." The resedimentation of juvenile pyroclastics (and effusives) may take place by mass flow, traction, or suspension.

Discussion of the merits of the nomenclature of McPhie et al. (1993) versus the traditional epiclastic nomenclature are beyond the scope of this work, but it is suggested here that epiclastic kimberlites may be better described using the terminology of McPhie et al. (1993), as this gives a clearer indication of their origin. Hence, in this work crater facies "epiclastic" rocks are referred to as resedimented volcaniclastic kimberlites (sensu McPhie et al. 1993). However, many readers of this monograph may prefer to continue using the older terminology until volcanologists agree upon the terminology ofthese types


of rocks. Consequently, the term "epiclastic" is also given in parentheses wherever the new terminology is used, to provide continuity with existing literature.

Resedimented volcaniclastic deposits are named by McPhie et at. (1993) as resedimented ash-rich mudstone, resedimented ash-rich sandstone, etc., following the standard particle size terminology employed in volcanology (Fisher and Schminke 1984). In the case of kimberlites they may be termed "resedimented kimberlitic ash-rich sandstone," etc. However, the particle size terminology was adopted from that devised for clastic sediments (Wentworth 1922). The use of a name which has both grain size and compositional implications causes many volcanologists to be resistant to terming a resedimented pyroclastic deposit which does not contain significant quantities of detrital lithic material a "sandstone." Conversely, sedimentologists are loathe to term such deposits exhibiting sedimentary structures a "tuff." Unfortunately, there is as yet no agreed upon solution to this problem (see McPhie et al. 1993, Fritz 1993, Zuffa 1991).

Resedimented volcaniclastic rocks currently described as "epiclastic kimberlites" are commonly preserved in the slightly eroded kimberlites of Tanzania and Botswana (Mannard 1962, Hawthorne 1975). Xenoliths of these rocks are also found as downrafted blocks in the diatremes (Clement 1982). Despite their relatively common occurrence, their character and genesis are poorly understood as they have not been investigated by modem petrological techniques and are extensively altered. The majority of the crater facies resedimented volcaniclastic (epiclastic) kimberlitic rocks which have been studied to date contain juvenile and extraneous lithic clasts in addition to authigenic components. They have been usually interpreted to have formed in crater lakes formed above diatremes (Mannard 1962, Hawthorne 1975, Mitchell 1986).

Two principal varieties of resedimented volcaniclastic (epiclastic) deposit may be recognized: syn-eruptive and post-eruptive. The former are produced during periods of quiescence during the lifetime of the vent and may be interbedded with pyroclastic rocks (see below). Post -eruptive resedimented volcaniclastics (epiclastics) are formed immediately after eruptions have ceased and represent the products of destruction of the volcanic edifice by contemporaneous erosion. They may occur within the crater as crater lakes become filled with resedimented pyroclastic material. Infilling may occur soon after termination of activity at a time when the tuff ring and crater walls still exist. Such resedimented volcaniclastics (epiclastics) may contain a high proportion of juvenile clasts. Other post-eruptive deposits may form contemporaneously outside the crater (see below).

When volcanism is episodic, resedimented volcaniclastic (epiclastic) rocks may be interbedded with subaerial pyroclastic rocks. Fragmentation of these rocks by later volcanic events will result in the recycling of lithic and juvenile debris into newly formed pyroclastic rocks (Houghton and Smith 1993). Reworking of the latter will generate new resedimented volcaniclastic (epiclastic) rocks. The proportions of resedimented volcaniclastic (epiclastic) to pyroclastic rocks will vary as the vent (or vents) evolves. Identification of a given unit as pyroclastic or resedimented volcaniclastic (epiclastic) becomes extremely difficult when components have been recycled several times. Note that determination of the origin of crater facies units from drill-core samples is particularly difficult as the lithological relationships between units cannot be observed.

40 CHAPTER!

Most discussions of crater facies rocks, in the kimberlitic environment have been restricted to those occurring within the crater. This is a direct consequence of the absence of modern kimberlitic volcanism and the generally extensive erosion of kimberlite intrusions. Base surge deposits, pyroclastic rocks, and resedimented volcaniclastic (epiclastic) rocks, such as debris flows, lahars, and lacustrine deposits, formed outside craters have not been previously considered as they are not typically preserved. However, such extracraier rocks may comprise portions of kimberlite intrusions at Fort a la Corne, Saskatchewan (Lehnert-Thiel et al. 1992), and Bakwanga, Zaire (Bardet 1974). Unfortunately detailed modern mineralogical and petrological studies of these rocks are not available.

Lacustrine and other extra-crater resedimented volcaniclastic (epiclastic) rocks can be classified and described according to standard volcano-sedimentological terms and facies analyses as described by Cas and Wright (1987).

1.9.1.1.d. Volcanogenic Sedimentary Deposits. In some cases, sedimentary rocks of similar appearance to syn- and post-eruptive resedimented volcaniclastic (epiclastic) rocks may represent material washed into depressions formed above vents or diatremes, long after volcanism has ceased. In these examples the original volcanic edifice has been removed by erosion prior to the formation of the sedimentary unit, and thus a crater lake is not present. The depressions may be formed either by subsidence of the material in the vent or differential erosion. Examples oflakes occupying such depressions are the "pans" and post-glacial lakes formed above kimberlites in southern Africa and Canada, respectively.

Material washed into the depressions may be devoid of juvenile components, being entirely derived from the terrain outside the depression. Such deposits should not be termed either "crater facies rocks" or "resedimented volcaniclastic (epiclastic) rocks," as they are merely normal lacustrine sediments.

In other cases, rocks containing a juvenile component may occupy the depressions. These may be formed by the reworking of preexisting igneous rocks coupled with input of extraneous lithic material. The time of deposition of these materials may be millions of years after the formation of the vents. Erosion of the vent to hypabyssal or lower rootzone facies may have occurred and the sediments deposited directly upon these rocks. The presence in the rocks of igneous clasts qualifies them as an "epiclastic" deposit in the original sedimentological sense of the term, in that the juvenile material is derived from lithified rocks. However, they cannot be regarded as crater facies rocks as they did not accumulate in, or adjacent to, an active volcanic crater and are not directly related to the volcanic activity. In this work such pseudo crater facies rocks are termed metachronous volcanogenic sediments (epiclastics). Metachronous (Greek meta = after + khronos = time) is used to indicate that they formed a long time after the magmatic event. The rocks are termed "volcanogenic" rather than "resedimented volcaniclastic" to indicate that the juvenile component was derived from ancient lithified igneous rock (McPhie et al. 1993).

It is very important to distinguish these rocks from bona fide synchronous (syn- and post-eruptive) crater facies rocks of similar appearance. Failure to do this can lead to incorrect interpretations of the evolution of kimberlite intrusions (see 1.9.1.4).


Pseudo crater facies volcanogenic sedimentary (epiclastic) rocks are known from the Kelsey Lake kimberlite, Colorado (Coopersmith 1993), several Yakutian kimberlites (see 1.9.1.4), and diatremes in the Arkhangelsk diamond province. In the Kelsey Lake occurrence, Recent metachronous volcanogenic sedimentary (epiclastic) rocks form at the current erosional surface above Devonian kimberlites which contain downrafted bona fide resedimented volcaniclastic (epiclastic) xenoliths. This is an excellent example of the deficiencies of the older terminology, as the same name is applied to genetically and temporally different rocks.

Topographic depressions in the upper levels of diatremes of the Arkhangelsk Province (Kudrjavtseva et al. 1991, Kharkiv 1992, Sinitsyn et al. 1992) are filled with lacustrine breccias, the components of which are commonly derived entirely from the Vendian country rocks. Petrographic studies reveal a few examples containing a very minor juvenile igneous component consisting of altered phlogopite and olivine macrocrysts or very thin layers of tuff. Russian petrologists term these metachronous volcanogenic sedimentary (epiclastic) rocks "compensational sediments" and interpret them as crater facies rocks, even though they recognize that they were formed subsequent to the cessation of volcanic activity.

Brief general descriptions of crater facies kimberlites have been provided by Mannard (1962), Edwards and Howkins (1966), Rolf (1973), Hawthorne (1975), Clement (1982), and Clement and Skinner 1985). A review of the geology of crater facies kimberlites has been provided by Mitchell (1986).

1.9.1.2. Diatreme Facies To reflect the association of diatremes with a variety of rocks, including kimberlitic,

minette, and melilititic volcanism, Mitchell (1986, p. 74) defined diatremes as cone shaped. downward tapering. inclined or vertical structural units. composed wholly or partly of angular. or rounded. clasts of cognate or xenolithic origin. with or without a matrix. Xenolithic clasts may be derived from the walls or roof of the body. They are commonly well-mixed, and some xenoliths have apparently sunk within the diatreme. Diatremes are volcanic features associated with volatile-rich magmatism commonly of an ultrabasic composition.

Note that the term "diatreme" is not synonymous with "volcanic vent." Kimberlite diatremes are long carrot-shaped bodies underlying crater facies rocks

(Figure 1.13). They have near-vertical axes and steeply dipping margins (80-85°). Their elliptical or roughly circular cross-sectional area decreases regularly with depth until they terminate in a region known as the root zone. This zone (Figure 1.13) is marked by the expansion, contraction, or splitting up of the diatreme into an irregularly shaped multiphase intrusion of hypabyssal kimberlite. Detailed descriptions of kimberlite diatremes are provided by Hawthorne (1975), Clement (1982), and Mitchell (1986). Note that kimberlitic diatremes in the Yakutian kimberlite province might not conform in structure to the South African model (see 1.9.1.4) as diatreme morphology may be strongly influenced by the fluid content, type, and rigidity of the country rock.

Diatreme facies kimberlites are currently classified as "tuffisitic kimberlite" and "tuffisitic kimberlite breccia" (Clement and Skinner 1985, Clement 1982). The latter differ from the former only in that they contain more than 15 vol % clasts greater than 4 mm in maximum dimension. The rocks consist of lithic fragments and discrete mineral

42 CHAPTER 1

grains of diverse origin cemented by extremely fine-grained minerals which are interpreted by Clement (1982) as quench products derived from the condensates of vapor phases in short-lived, vapor-liquid-solid fluidized systems. Clement (1982) thus regards the rocks as intrusive tuffs or tuffisites. Use of the term "tuffisite" (Cloos 1941) has definite genetic connotations and implies formation of a diatreme by gas-solid fluidization processes, i.e., tuffisitization. Clement and Skinner's (1985) terminology is correctly applied if all diatreme facies rocks form by such a process. However, there is no agreement on how diatreme facies rocks are formed. Currently, opinions cover the entire spectrum between hydrovolcanic and fluidization origins (see reviews by Clement 1982, Mitchell 1986, Clement and Reid 1989). Thus, while the use of genetic-specific terms is appropriate for a textural-genetic classification, they should not be applied indiscriminantly. A specific genetic term should only be used when the origin of a particular diatreme facies rock is understood.

As the nature of the volcanic processes giving rise to most diatreme facies rocks is not known, they are better termed volcaniclastic kimberlites. The general nongenetic term volcaniclastic was introduced by Fisher (1961) to include alI clastic volcanic materials formed by any process of fragmentation, dispersed in any transporting agent, and deposited in any kind of environment.

Thus, it is suggested that Clement and Skinner's (1985) genetic term "tuffisitic" be abandoned as a general term for diatreme facies kimberlites and they be described as volcaniclastic kimber/ites and volcaniclastic kimberlite breccias. Only if the origins of the clasts can be determined should a second level of classification be employed, e.g., tuffisitic kimberlite, hydroclastic kimberlite. Figure 1.15 presents the recommended nomenclature of diatreme facies kimberlites. Note that this nomenclature is applicable to orangeites and melilitoids but not to lamproites. Other textural terms which have been inadequately defined in previous works include pelIetal lapilli, autoliths, and nucleated autoliths.

DIATREME FACIES

VOLCANICLASTIC KIMBERLITE BRECCIA (15 %; )4 mm Clasts ) 15 %; )4 mm Clasts

) 1 0 % PeHetal Lapilli

No

Ve.

PELLETAL-TEXTURED

Figure 1.15. Textural-genetic classification of diatreme facies kimberlites_


Pelletal lapilli (Clement 1973, 1982, Clement and Skinner 1985) are particularly characteristic of diatreme rocks and may be regarded as the "hallmark" for the recognition of diatreme environments. Especially important is that pelletallapilli are not confined to kimberlites and may be found in melilitite (Cloos 1941, Rust 1937, Reed and Sinclair 1991) and orangeite diatremes, e.g., Finsch (Clement 1982). Pelletal lapilli are thus characteristic products of a particular process operating during the formation of diatremes but are not indicators of a particular magma type. However, all rocks containing pelletal lapilli appear to be derived from magmas that are C02"fich relative to common magma types.

Pelletal lapilli are discrete spherical-to-elliptical lapilli-sized (2-64 mm) clasts consisting of fine-grained primary igneous material. Lapilli commonly contain, at their centers, a single, relatively large euhedral crystal or crystal fragment (Figure 1.16). These cores, or kernels (Clement 1973), consist typically of olivine (fresh or pseudomorphed) and, less commonly, phlogopite or other macrocrysts. Country rock clasts very rarely form the cores oflapilli, and juxtaposed country rock clasts are typically devoid of igneous mantles.

The mantles consist of very fine-grained microphenocrystal material. Minerals found within the mantle are primary groundmass and mesostasis phases characteristic of the parental magma. Thus, in kimberlites the mantle consists of microphenocrysts of euhedral olivine and/or laths of phlogopite set in a groundmass of perovskite, spinel, calcite, and serpentine. Prismatic minerals are commonly oriented around the kernel of the lapilli and a poorly-to-well-developed concentric structure may be discernible (Figure 1.16).

Pelletallapilli within a given sample exhibit a wide range of textural charaCteristics and degree of crystallinity. They have the textural characteristics of having been formed by the rapid crystallization of volatile-poor magma containing phenocrysts and may be described as microphenocrystal hypabyssal kimberlite. They are not, on the basis of their mineralogy and texture, examples of accretionary lapilli.

The origins of pelletallapilli have been reviewed by Clement (1982), Lorenz (1979), and Mitchell (1986). Currently, they are believed to represent magma droplets formed by the explosive fragmentation or frothing of magma due to either rapid expulsion of dissolved volatiles (Clement 1973, 1982, Dawson 1980) or interaction with groundwater (Lorenz 1979, Mitchell 1986). In the latter mechanism, known as a fluid-coolant (magma-water) interaction (Sheridan and Wohletz 1983), groundwater is believed to flash instantaneously into steam, leading to the explosive ~isruption of the magma into an "aerosol-like" cloud of rotating droplets of silicate liquid supported by superheated steam.

Regardless of the mechanism of formation, the presence of olivine nuclei indicates that, prior to the formation of the lapilli, either some crystallization of the magma had occurred or that it contained abundant olivine macrocrysts. Subsequent to their generation, surface tension effects must have resulted in magma adhering to these earlier-formed crystals. The shapes and textures of the lapilli undoubtedly result from surface tension effects combined with rotation of the drops during transport. The magma which formed the mantles was initially fluid, and their textures demonstrate that crystallization continued after formation of the lapilli. Rapid quenching and the formation of glass are not a

44 CHAPTER!

Figure 1.16. (a) Pelletal lapillus. Postmasburg orangeite. Oriented laths of mica surrounding a kernel of altered olivine. (Field of view (FOY)) = 1 nun; (b) pelletallapilli. Finsch orangeite (FOY = 4 nun); (c) pelletallapilli. Chomur kimberlite. Russia (FOY= 4 nun); (d) pelletal lapillus. Urach melilitite (FOY = 2.25 nun).


Figure 1.16. (continued)

46 CHAPTER 1

part of the lapilli-forming process in kimberlites, although glassy matrices might have been originally present in melilititic pelletallapilli.

Autoliths are defined here as angular-to-subrounded, lapilli- to ash-sized clasts formed by the fragmentation of preexisting solidified material. The term is used in the original sense of referring to fragments of solidified kimberlite found within· a younger kimberlite (Rabkin et al. 1961). Autoliths can be derived from any facies and therefore have any texture. However, autoliths of hypabyssal material appear to be the most common variety.

The textures of hypabyssal autoliths are identical to those of hypabyssal intrusive rocks and they may be classified and described by terminology appropriate to the latter (1.9.1.3). No preferred orientation of minerals is evident within the clasts. Commonly, fractured microphenocrysts and/or macrocrysts located at the margins of the clasts provide indisputable evidence that they originated from the fragmentation of solid material. Autoliths of differing mineralogy and textural type may coexist, and they may or may not occur with pelletal lapilli (Figure 1.17). Subrounded autoliths may be impossible to distinguish from pelletallapilli with eccentric kernels that exhibit only weak concentric structures.

Considerable confusion exists in kimberlite literature concerning the meaning of the term "autolithic kimberlite." In many Russian publications autolithic kimberlites are synonymous with pelletal-textured kimberlites, and no distinction is made between pelletal lapilli and autoliths, even though both may occur in the same sample. In other cases, diatreme facies rocks are said to contain "nucleated autoliths," an imprecisely defined term that may refer to either bona fide pelletal lapilli or large (up to 8 cm) spheroidal masses of hypabyssal rock with cores of country or basement rocks. The latter have been described by Ferguson et al. (1973) and Danchin et al. (1975) as "nucleated autoliths." They are not regarded by Clement (1982) or Mitchell (1986) as varieties of pelletallapilli (see below).

In this work it is suggested that

• Pelletal-textured volcaniclastic diatreme facies rocks are those which contain significant amounts (> 15 vol %) of pelletallapilli as defined above, e.g., pelletal-textured volcaniclastic kimberlite. Pelletallapilli are primary magmatic constituents and not xenolithic clasts.

• Autolithic volcaniclastic diatreme facies rocks are those which contain significant amounts (> 15 vol %) of autoliths as defined above, e.g., autolithic volcaniclastic kimberlite breccia.

• The term "nucleated autolith" (sensu Danchin et al. 1975) be abandoned (see below).

A given diatreme facies rock may contain significant amounts of both pelletallapilli and autoliths and should be termed a pelletal-textured autolithic volcaniclastic kimberlite breccia. Note that a rock containing only pelletallapilli is not in itself a breccia, although pelletal-textured rocks which contain significant amounts (> 15 vol %) of country rock clasts may be termed pelletal-textured lithic volcaniclastic kimberlite breccias. Rocks containing both autoliths and country rock clasts may be described as heterolithic volcaniclastic kimberlite breccias. Representative idealized examples of these diverse


Pelletal Lapilli

~'QJ.1: Autoliths .: ... ' ..

~ Macrocrystal "-2...:::) Olivi ne

47

Figure 1.17. Idealized representations of the textures of diatreme facies volcaniclastic kimberlite and volcaniclastic kimberlite breccias (VKB): (A) pelletal-textured volcaniclastic kimberlite. Note this rock is not a breccia; (B) autolithic VKB; (C) pelletal-textured autolithic VKB; (D) pelletal-textured heterolithic VKB; (E) pelletal-textured lithic VKB; (F) heterolithic VKB.

textural varieties of diatreme facies rocks are depicted in Figure 1.17. Note that although olivine (and other) macrocrysts are common in diatreme facies volcaniclastic rocks, their presence is usually not considered in the terminology of diatreme facies rocks.

Pelletal lapilli, autoliths and country rock clasts are set in a matrix that may be described as having a uniform or segregationary texture (Clement and Skinner 1985). The terms refer to whether or not the matrix minerals constitute a uniform aggregate or have crystallized into discrete patches of differing mineralogy. The interclast matrices are very fine-grained and consist predominantly, when fresh, of micro- to cryptocrystalline diopside and serpophitic serpentine. Such matrices are particularly prone to subsolidus hydrothermal alteration and/or weathering, and are commonly replaced by an optically unresolvable mixture of serpentine, chlorite, and clay mineral. Interclast matrices do not contain primary groundmass typomorphic mineral assemblages. Thus, in kimberlitic

48 CHAPTER 1

diatreme facies rocks the interclast matrices are devoid of second-generation olivine, spinel, perovskite, monticellite, and calcite.

Textural features of the matrix are commensurate with rapid nonequilibrium deposition from a hydrothermal fluid (Mitchell 1986) or the condensate derived from a vapor-solid fluidized system (Clement 1982). Variations in the temperature, composition, and rate of crystallization of such fluids can account for gradations between the uniform and segregationary groundmass textures. In some instances, crystallization of the fluid commences with the nucleation of diopside upon substrates provided by the clasts and macrocrysts, the residue crystallizing subsequently as serpentine. Associated pelletal lapilli and autoliths do not contain diopside.

Mitchell (1986) has explained the abundance of diopside and the absence of monticellite in the interclast matrices as a consequence of contamination of the matrix -forming fluids with silica derived from country rock clasts. The process is considered to be analogous to the contamination of hypabyssal kimberlites by country rock xenoliths. The latter are commonly surrounded by fringes of diopside which forms as a result of reaction between the silica-rich xenolith and the magma. Assimilation of the xenolith is believed to raise the silica activity of the magma to levels which support the crystallization of diopside in preference to monticellite.

1.9.1.3. Hypabyssal Facies

Hypabyssal kimberlites comprise the root zones of diatremes and occur as dikes and sills (Figure 1.13, Mitchell 1986, Clement and Skinner 1985, Clement 1982, Dawson 1980). The recommended textural-genetic terminology is given in Figure 1.18. Depending upon the presence or absence of clasts, rocks may be described as hypabyssal kimberlite or hypabyssal kimberlite breccia. The latter are defined as containing more

HYPABYSSAL FACIES

APHANITIC

MACROCRYSTAL

PORPHYRITIC

SEGREGATIONARY

Figure 1.18. Textural-genetic classification of hypabyssal facies kimberlites.


than 15 vol % of clasts greater than 4 mm in maximum dimension (Clement and Skinner 1985). The clasts may comprise country rocks and/or kimberlitic autoliths. The latter are regarded as fragments of previously consolidated earlier generations of hypabyssal kimberlites. Thus, lithic, autolithic, and heterolithic hypabyssal kimberlite breccias may be recognized. Note that autoliths in the diatreme and hypabyssal environment may be mineralogically similar but texturally different, with the latter being typically coarser grained and less altered. Pelletallapilli do not occur in hypabyssal facies rocks, although spherical masses of kimberlite, known as "globular segregations," are relatively common (see below).

Kimberlites and kimberlite breccias may be further described as aphanitic or macrocrystal in character, with porphyritic kimberlites being extremely rare (see 1.2). Previous textural-genetic classifications have not incorporated any information as to the nature of the macrocrysts. Although olivine is by far the most abundant macrocryst, there do exist kimberlites (De Beers, Tunraq, Zagodachnaya, Koidu) which contain significant amounts of macrocrystal phlogopite, ilmenite, or garnet. It is recommended that the term "macrocrystal" when used without prefix refers to a macrocrystal assemblage dominated by olivine. Prefixes may be added to recognize the presence of significant modal amounts (>15 vol %) of other macrocrystal phases, e.g., a phlogopite macrocrystal kimberlite would contain macrocrysts of both phlogopite and olivine. The recognition ofphlogopite macrocrystal kimberlites is very important as these rocks are not petrologically synonymous with, and cannot be termed, "phlogopite kimberlites," as this term is reserved for describing kimberlites rich in primary groundmass phlogopite.

The groundmass may be described as uniform or segregationary, depending upon whether the primary groundmass phases and the mesostasis have crystallized together or separately. Rocks with a uniform groundmass may be described by standard terminology and named using the mineralogical-genetic principles outlined in Section 1.4.4., e.g., hypabyssal macrocrystal apatite monticellite calcite kimberlite heterolithic breccia.

Segregationary-textured kimberlites have a nonuniform distribution of groundmass minerals and mesostasis. The groundmass is commonly identical in texture and character to that of uniformly textured hypabyssal kimberlite. The segregations are amoeboid-tospherical discrete regions consisting of relatively coarse-grained phlogopite-kinoshitalite solid solutions, apatite, calcite, and primary serpentine. Segregations consisting of calcite and serpentine are particularly common. Although the mineralogy of the mesostasis of the uniformly textured portions of the rock and the segregations may be similar, the latter typically lack perovskite and spinel. Segregations may be described as being bounded or gradational, depending upon whether their contact with the uniformly textured groundmass is sharp « 50 Jlm) or gradational (20-500 Jlm). Commonly, euhedral crystals of apatite, calcite, and phlogopite project into the serpentine-filled centers of segregations.

The segregations result from the separation of late-crystallizing components from the silicate oxide groundmass into discrete masses. In some instances, e.g., the Benfontein and Skinner's Sills, segregations appear to have migrated through their partially crystallized parental liquids prior to complete consolidation of the magma. The origin and mobilization of segregations has been discussed at length by Dawson and Hawthorne (1973), Donaldson and Reid (1982), Mitchell (1984a, 1986), and Clement (1982). A consensus regarding their genesis has not been reached and segregations are variously

50 CHAPTER 1

regarded as gas condensates in vesicles, filled breached vesicles, immiscible liquids, or low-temperature residual fluids. In the latter case the segregations form as a result of surface tension effects between the water-rich segregation and more viscous crystal-rich silicate oxide groundmass (Mitchell 1986).

Globular segregations are spherical masses of hypabyssal material which may range up to 100 mm in diameter. When present in large numbers, they confer a pseudoconglomeritic appearance to the rock. In kimberlites they consist of relatively fine-grained hypabyssal kimberlite and may be found locally in coarser grained, but otherwise similar, uniformly textured hypabyssal kimberlite. Clement (1982) and Mitchell (1986) have suggested that the segregations are generated by surface tension effects in boiling magmas in near-surface hypabyssal environments.

There are also known spherical masses of kimberlite which contain distinct cores of xenolithic material. The xenoliths appear to be predominately basement gneisses, and relatively few cores consist of high-level country rocks. These objects are typically found irregularly distributed within diatreme facies volcaniclastic rocks. Commonly, they are strongly concentrically zoned, the zonation resulting from the presence of many thin concentric shells of fine-grained hypabyssal kimberlite of differing thickness and mode. The size of the kimberlitic mantle bears no systematic relation to the size of the xenolithic core. Examples are known from several Lesothan (Ferguson et al. 1973) and Yakutian (Legkaya, Zarnitsa) kimberlites.

The objects have been termed "nucleated autoliths" by Danchin et al. (1975) and Ferguson et al. (1973). This term is inappropriate as they are not fragments of preexisting rocks. They have not crystallized in their current hosts and must represent a transported assemblage. In contrast, the globular segregations described from Finsch and Dutoitspan (Clement 1982) or Mukorob (Mitchell 1986) appear to have crystallized in situ in the hypabyssal environment.

Whether the two varieties of spherical "segregations" have the same origin, and thus should have the same name, is debatable, as they have not yet been sufficiently studied. Until further information regarding their character and genesis becomes available, it is recommended that the term "nucleated autolith" (sensu Danchin et al. 1975) be abandoned and these objects be referred to by the nongenetic term nucleated globules. In contrast, spherical objects which appear to have been generated in situ by local devolatization, or boiling of magma in closed systems in near-surface hypabyssal environments (Mitchell 1986, Clement 1982) should retain the appellation globular segregation.

Note that globular segregations are not confined to kimberlites and are known from lamproite dikes (Pilot Butte, Wyoming, this work) and orangeites (Finsch, Clement 1982), suggesting that a common mechanism is responsible for their formation in these diverse volatile-rich magmas. Thus, the terminology, while genetic in character, is not magma specific.

Globular segregations have petrographic similarities to some pelletallapilli (Clement 1982); however, they are typically coarser grained and commonly lack a macrocrystal nucleous. Clement (1982) has suggested that some coarse-grained pelletallapilli grew as small globular segregations and were subsequently mixed with fine-grained "spray" or "aerosol" pelletal lapilli that were generated during the diatreme-forming process. This hypothesis assumes that pelletallapilli were also formed primarily by volatile degassing.


However, lapilli may have hydroclastic origins (Mitchell 1986), thus ruling out the existence of a genetic continuum between pelletallapilli and globular segregations. Of course, petrographically similar, but genetically different, lapilli may coexist as a hybrid assemblage. Clearly, much further work on these problems is required.

1.9.1.4. Spatial Relationships between Diatreme and Hypabyssal Facies Kimberlites

Current ideas regarding the spatial relationships between diatreme and hypabyssal facies rocks are based upon data obtained from the deep mining of several southern African kimberlites. According to this model (Figure 1.13), hypabyssal rocks occur only in the root rones of diatremes and as sills and dikes. The high levels of diatremes lack intrusive hypabyssal rocks, and in this regime they are present only as clasts of disaggregated preexisting rocks. There are no obvious physicochemical reasons precluding the formation of hypabyssal rocks at high levels in diatremes. Such rocks are common in melilitite diatremes and typical of lamproitic vents.

Recently, information has become available regarding the structure of some Yakutian diatremes which seems to be at variance with the southern African model. Thus, diatremes in the Daldyn-Alakit field differ from those in southern Africa in that they are commonly multiple intrusions in which massive macrocrystal hypabyssal kimberlite occurs at higher levels in the diatreme.

Figures 1.19 to 1.21 illustrate the structures of the Sitikanskaya, Krasnopresnenskaya, and Udachnaya intrusions of the Daldyn-Alakit field. These are all double intrusions which coalesce into a single complex body at high structural levels. On the basis of these data, Russian geologists believe that many geographically closely related

-.=:-.=---=---~-=-:::-::: -~-=-::: -=-:::-:::-=-:::-:::-:::-:::-::=-':--~-:-==--p:' T :::-:::-=-:::--:::---_-_-_-_ -:... __ - - - - - - - - - - - - - - - - - -- -- -- - - -- -- -- - - -~-$;::;;:;:;;:;:;:;:;:~

- 230

Figure 1.19. Cross section of the Sitikanskaya diatreme, Alakit field, Yakutia. This double pipe consists of an earlier intrusion of volcaniclastic kimberlite breccia and a later intrusion ofmacrocrystal hypabyssal kimberlite. Note the downrafted block of Silurian sediment (S 1), the Carboniferous-Permian metachronous volcanogenic (epiclastic) sedimentary rocks (C-P) and the capping of Permo-Triassic (P-T) basaltic lavas (after data provided by the Amakinsk Geological Expedition, Aikhal, Yakutia).

52 CHAPTER I

Figure 1.20. Isometric diagram illustrating the fonn of the Krasnopresnenskaya kimberlite, Alakit field, Yakutia. This Paleozoic age double pipe (stippled) has been intruded by a basaltic sill (dashed) of Penno-Triassic age. Bowl-shaped depressions in the upper parts of the pipes are filled with metachronous volcanogenic (epiclastic) sedimentary rocks and covered with younger basaltic volcanic rocks (see Figure 1.23 for details) (after Kriuchkov et al. 1994).

intrusions represent the eroded roots of such double-, or even multiple-, pipe systems. This approach is illustrated in Figure 1.22 which shows that four intrusions, exposed at the current level of erosion in the Ukukitskoye field, are believed to have formed a large coalescing vent at the time of intrusion.

Cross sections of the Krasnopresnenskaya (Figure 1.23) and Yubileinaya (Figure 1.24) intrusions illustrate further significant differences with respect to southern African pipes. The upper part of the Krasnopresnenskaya intrusion (Kriuchkov et at. 1994) flares rapidly into a wide bowl-shaped body above the cylindrical, smooth-sided main feeder vent. The sides of the bowl-shaped unit dip at 50-60° adjacent to the feeder and at 15-25° at the margins. The bowl is filled with Lower Carboniferous metachronous volcanogenic (epidastic) lacustrine sediments which may be up to 50 m thick and which cover most of

WEST EAST

Figure 1.21. Cross section of the Udachnya kimberlite. Each pipe is filled with different varieties of volcaniclastic and macrocrystal kimberlites as indicated by the different ornamentation (after Milashev 1984).

BALATINSKAYA

FESTIVALNAYA VASILYEOSTROVSKAYA

PETROGRADSKAYA Figure 1.22. Postulated structure of a group of kimberlite diatremes in the Ukukitskoye field. Yakutia, prior to erosion. At the current level of exposure there occur four apparently unconnected intrusions (after V. Kornilova. pers. comm.).

S3

54 CHAPTER 1

Figure 1.23. Cross section of the Krasnopresnenskaya diatreme. Alakit field. Yakutia. Kl and K2 are volcaniclastic kimberlite breccias. Note that K2 contains abundant xenoliths of wall rock Silurian limestone together with xenoliths of Ordovician marl which have been transported upward in the diatreme. The bowl-shaped depression above K2 contains Carboniferous metachronous volcanogenic (epiclastic) sedimentary rocks (KCF). The intrusion is capped by a variety of Permo-Triassic basaltic rocks and cut by a basaltic sill of the same age (after Kriuchkov et al. 1994).

the underlying diatreme. Late Carboniferous sedimentary rocks unconformably overlie the volcanogenic (epiclastic) sedimentary rocks and cover the entire diatreme. The kimberlites and the overlying sedimentary rocks have been intruded, and covered, by Permo-Triassic basaltic dikes, sills, and "tuffaceous units."

KIMBERLITES AND ORANGElTES

o

400

600

800

1000 m

w

•VKB

+ +

+

+

+

ss

E

+ +

+ +

+ + +

+. + VKB + +

+ + + +

Figure 1.24. Cross section of the Yubileinaya kimberlite, Alakit field. Yakutia. The intrusion consists of an early volcaniclastic kimberlite breccia (VKB) which has been intruded by a macrocrystal hypabyssal kimberlite breccia (MKB). Note the "flaring" of the latter unit at the top of the pipe. The bowl-shaped depression above the MKB is filled with metachronous volcanogenic (epiclastic) sedimentary rocks (ME), and the whole intrusion is capped by Perrno-Carboniferous (P-C) sediments (after Kharkiv 1990).

The vent is filled with two types of kimberlites. The upper phase is termed a "kimberlite tuff breccia," the lower phase a "autolithic tuff breccia." Samples of these rocks are not yet available for study, and consequently it is not possible to describe them in terms of current textural-genetic classifications.

Figure 1.24 shows that the upper portions of the Yubileinaya intrusion (Kharkiv 1990) also flare out rapidly into a bowl-shaped body. The intrusion consists of at least two kimberlite types with the youngest of these forming the flared body. Russian petrologists consider that each intrusion consists of "autolithic kimberlite breccia" and "kimberlite with a massive cement" (sensu Komilova et at. 1983). The former is equivalent to autolithic kimberlite breccia, and the latter to macrocrystal hypabyssal kimberlite as used in this work. Note that hypabyssal kimberlites are apparently present at high structural levels in this intrusion as the erosion level in this field is not great enough to expose bonafide root zones and feeder dike systems.

The upper levels of the intrusion are covered by metachronous volcanogenic (epiclastic) sedimentary rocks that are considered by Kharkiv (1990) to be crater facies rocks (see below).

The Sitikanskaya (Figure 1.19) pipe shows similarities to the Yubileinaya intrusion. Here a complex multiphase unit of pelletal-textured and volcaniclastic diatreme facies kimberlite breccias is intruded by a late-stage unit of macrocrystal hypabyssal kimberlite

56 CHAPTER 1

• • • • • • • • • • • • • • • • • • • • • • • • • • • • • DOLERITE P2- T3 • • • • • • • • • • • • •

LIMESTONE

Figure 1.25. Cross section of the Odintsov pipe. Alakit field. Yakutia (after Kriuchkov er al. 1994). K-I: porphyritic kimberlite; K-2: autolithic kimberlite breccia; K-3: supra-pipe breccia-limestone + kimberlite.

and autolithic kimberlite breccia. This latter unit again flares out into a shallow-dipping body in the upper parts of the vent. Metachronous volcangenic (epiclastic) sedimentary rocks are preserved in a bowl-shaped depression above the northeastern intrusion.

The Odintsov pipe (Kriuchkov et al. 1994) appears to be a proto- or blind diatreme (Clement 1982, Mitchell 1986), although this intrusion has no exact counterpart in the southern African kimberlite province. Russian geologists consider the body to be a cryptovolcanic structure. The "vent" consists of a complex breccia unit termed the "carbonate cap," which is underlain by a pipe of kimberlite (Figure 1.25). The carbonate breccia consists of local country rock limestones which have been fragmented by explosive volcanism. Clasts in this breccia range in size from 5 to 10 cm together with large (several meters) xenoliths. The clasts are cemented together by kimberlite-derived material and the cap rock is extensively veined in the lower regions by massive kimberlite. The cap rock clearly represents the fragmentation of the country rock above an advancing pulse of kimberlite magma. The body is described as a "semi blind kimberlite" by Russian


geologists as the kimberlite-derived fluids are believed to have penetrated to the surface, while the kimberlite did not actually erupt.

Intrusive kimberlites are present below the cap rock. These consist of an early generation of "porphyritic kimberlite" which has been intruded by a later "autolithic kimberlite" (sensu Kriuchkov et al. 1994). The former probably corresponds to macrocrystal hypabyssal kimberlite, although the exact nature of the latter is uncertain. If it is a volcaniclastic diatreme facies unit, the sequence of intrusion would be quite unlike that seen in southern Africa.

Multiple pipes are not, with the exception of the Mir-Sputnik pair, characteristic of the Malo-Botuobinsk field, and diatremes (International, 23CPC) in these fields are similar to southern African diatremes. The kimberlite fields in the northern parts of the Yakutian province, e.g., the Kuoiskoye, Molodinskoye, Toluopskoye fields, are more deeply eroded, and only root zones and the feeder dike systems, consisting of hypabyssal kimberlites, are preserved. Many of the small hypabyssal intrusions, e.g., Obnazhennaya, Russlovaya, Festival, Anomaliya 23, in the Kuoiskoye field, are similar to "blows" developed along dike systems in the southern African fields, e.g., pipe 200 (Kresten and Dempster 1973). Russian geologists explain these differences from the Daldyn-Alakit and Chomur fields on the basis of the relative degree of erosion of the fields. Thus, multiple and flared pipes are not present in the above fields due to the greater extent of erosion.

In summary, the character of high-level intrusions in the Daldyn-Alakit fields suggests that

• Macrocrystal hypabyssal facies kimberlite may exist in the upper levels of diatremes.

• Volcaniclastic kimberlites may intrude hypabysssal kimberlites. • Multiple coalescing intrusions are common. • The upper parts of intrusions commonly flare out into wide bowl-shaped units

which subsequently form erosional depressions which may be filled with pseudo crater facies, metachronous volcanogenic (epiclastic) sedimentary rocks.

With respect to the last conclusion, which contradicts Kharkiv (1990), it is apparent that volcanogenic (epiclastic) sedimentary rocks in the Daldyn-Alakit field are not true crater facies rocks. This may be demonstrated at Sitikanskaya where these rocks are found to rest directly upon a downrafted block of Silurian limestone enclosed within volcaniclastic diatreme facies rocks (Figure 1.19). Thus, the volcanogenic sedimentary units cannot be either in situ or downrafted crater facies material. They merely occupy an erosional depression formed in the pipe subsequent to erosion of the actual crater and uppermost parts of the diatreme (see 1.9.1.1). They and other pseudo crater facies rocks found at Aikhal, Krasnopresnenskaya, and Yubileinaya formed on a Lower Carboniferous paleoerosion surface, and have been preserved only because of their subsequent burial beneath a protective cover of Permo-Triassic basalts. Consequently, although high structural levels of diatremes are found in the Daldyn-Alakit region, true crater facies rocks are not preserved in the Yakutian province.

The differences in the intrusive style of the Yakutian diatremes relative to the southern African diatremes are significant, but as yet unexplained. Paired or triple diatremes which

58 CHAPTERl

coalesce at higher structural levels are known from southern Africa, e.g., Venetia, Jwaneng. Clement (1982) has suggested that the Koffiefontein-Ebenhaezer and Dutoitspan-Bultfontein diatremes may also have coalesced at higher levels. However, such occurrences are not typical of the province as a whole.

As the petrographic character of the kimberlites in the two provinces is similar, it is apparent that the differences in intrusive style must reflect differences in the nature and water contents of the intruded rocks rather than differences in the parental magmas.

1.9.2. Orangeites

Orangeites occur as diatremes, dikes, and sills (l.8). These intrusions, on the basis of our current knowledge, apparently do not differ in style from those of archetypal kimberlites. The textural-genetic classification developed for kimberlites in this work (Figures 1.14, 1.15, and 1.18) is therefore considered to be entirely applicable to orangeites. Thus, the diatreme facies Fl unit at Finsch (Clement 1982) may be described as a pelletal-textured heterolithic volcaniclastic orangeite breccia. Individual hypabyssal facies orangeites may be described according to the mineralogical-genetic classifications outlined in Section 1.4.4 and described in Section 1.1 O.

The majority of orangeite intrusions appear to belong to the hypabyssal and root-zone facies. Large diatremes are represented only by the Finsch pipe, and crater facies rocks have not been described.

1.9.3. Melilitite Clan

Melilitite magmas form diatremes, dikes, and sills in addition to occurring as lavas and plutonic rocks. The textural-genetic classification developed for kimberlites is applicable to rocks belonging to this clan with the addition of a plutonic facies. Most melilitoids can be described by standard igneous petrological terms.

Melilititic diatremes are common, e.g., Swabia (Cloos 1941, Lorenz 1979), Missouri (Singewald and Milton 1930), Montana (Hearn 1968), Sutherland, South Africa (McIver and Ferguson 1979), and James Bay Lowlands, Ontario (Janse et al. 1989). These display the full range of textures exhibited by diatreme facies kimberlites (and orangeites). Thus pelletal-textured autolithic volcaniclastic melilitite breccias (Figure 1.16) may be found in such diatremes. Particular care must be taken in exploration programs in distinguishing these rocks from superficially similar volcaniclastic kimberlites. This cannot be achieved using macroscopic observations alone, and correct assessment of petrological clan affinity requires detailed mineralogical studies using a combination of optical and electron microbeam methods (Mitchell 1995).

Although detailed discussion of the nature of melilititic diatremes is beyond the scope of this work, note that they differ from most kimberlite diatremes in exhibiting well-developed internal structures consisting of bedded volcaniclastic units and central conduits of hypabyssal material (Hearn 1968, Lorenz 1975, 1984). The occurrence of high-level magmatic rocks is similar to the presence of high-level hypabyssal facies material in some of the Yakutian diatremes, although the latter lack well-defined internal structures.


Figure 1.26. Macrocrystal orangeite. Star dike. M = phlogopite; 0 = olivine (TL. FOY = 4 mm).

Figure 1.27. Macrocrystal orangeite. Sover. M = phlogopite; 0 = olivine (TL. FOV = 4 mm).

60 CHAPTER 1

1.10. PETROGRAPmC CHARACTERISTICS OF ORANGEITE

Orangeites may be described using mineralogical-genetic classifications, as described in 1.4.4. The majority of orangeites do not differ greatly from one another in their petrographic character within and between intrusions. The pri!1cipal differences are with respect to the amount of macrocrystal olivine and the ratio of macrocrystal to microphenocrystal phlogopite. Only evolved orangeites contain significant modal amounts of diopside and/or sanidine.

Although diatreme and hypabyssal facies orangeites are known, it is only from the latter that samples suitable for undertaking detailed petrographic and mineralogical studies may be obtained. In common with kimberlites, diatreme facies rocks are typically altered and usually do not contain typomorphic minerals because of their rapid crystallization (Scott Smith 1992). Diatreme facies rocks are thus best identified as orangeites by their consanguineous association with hypabyssal facies rocks.

The principal varieties of hypabyssal orangeite are the following:

Figure 1.28. Macrocrystal orangeite, Bellsbank. Note the region consisting entirely of mica and apatite (white) at lower right. Backscattered electron image (BSE-image). 0 = olivine.


Figure 1.29. Macrocrystal orangeite. Sover. Note the areas rich in apatite (white) in the fine-grained phlogopite-rich matrix. BSE image. 0 = olivine.

Orangeite, a rock consisting principally of microphenocrystal phlogopite set in a fine-grained groundmass consisting essentially of phlogopite-tetraferriphlogopite and minor apatite, chromite, Mn-ilmenite, and perovskite, with a mesostasis of calcite and/or dolomite together with serpentine (Figures 1.26-1.45). The ratio of microphenocrysts to groundmass mica and/or mesostasis varies widely. Some rocks consist predominantly (85-90 vol %) of closely packed tablets of microphenocrystal phlogopite (Figure 1.33). Many of the microphenocrysts and groundmass micas are strongly-zoned from pale yellow phlogopite cores to bright-red tetraferriphlogopite margins (Figure 1.39).

Phlogopite macrocrysts are common (1-10 vol %) and typically deformed and broken (Figures 1.26, 1.27, 1.31). Typically, their composition is broadly similar to that of microphenocrystal phlogopite. Minor reverse zoning and mantling may be present. Cryptogenic macrocrysts of green or brown pleochroic biotite are rarely «<1 vol %)

present. These are typically mantled by colorless phlogopite (Figure 1.32). Relative to kimberlites the groundmass of orangeites is fine-grained, poor in spinels

and perovskites (see 1.11), and rich in apatite. The latter may occur as euhedral prisms

62 CHAPTER 1

Figure 1.30. Macrocrystal orangeite. Boshof. Note that the matrix consists primarily of microphenocrystal phlogopite (gray) and euhedral apatite (white). BSE image. 0 = olivine.

(Figures 1.37, 1.41) or poikilitic plates. Rare relatively coarse-grained patches (?segregations) in the ground mass consist predominantl y of phlogopite and apatite set in a calcite mesostasis (Figure 1.41).

Small euhedral-to-subhedral primary olivines may be present but are not typically abundant and are commonly completely serpentinized.

Macrocrystal orangeite is a rock identical to the above but characterized by a pronounced inequigranular texture due to the presence of olivine macrocrysts (Figures 1.26-1.30). The latter may be fresh, or partially or completely serpentinized. As a consequence of the extremely inhomogeneous distribution of macrocrysts within a given intrusion, there is a complete gradation from macrocrystal orangeite to rocks entirely lacking olivine macrocrysts. The majority of the macrocrysts, on the basis of their habit, undulose extinction, and occurrence as multiple-grain aggregates, are considered to be xenocrysts.

Diopside orangeite and macrocrystal diopside orangeite (Figures 1.46, 1.49) are similar to orangeite and macrocrystal orangeite as described above, but differ in contain-

Figure 1.31. Orangeite, New Elands. Macrocrysts of phlogopite (M) set in a very fine-grained groundmass consisting of euhedral spinels (black) and fine-grained phlogopite (TL, FOY = 4 mm).

Figure 1.32. Orangeite, Star. Macrocrysts of phlogopite (PHL) and dark-colored aluminous biotite (B) together with euhedral, mantled microphenocrystal phlogopite (M), and opaque spinel (TL, FOY '" 2.25 mm).

63

Figure 1.33. Orangeite, Swartruggens. Closely packed aggregate of microphenocrystal phlogopite (TL, FOY =

1 mm).

Figure 1.34. Orangeite, New Elands. Microphenocrysts of phlogopite and subhedral opaque spinels set in carbonate-rich mesostasis (TL, FOY = 1 mm).

64

Figure 1.35. Orangeite, Finsch. Plates of groundmass phlogopite enclosing opaque subhedral spinels (TL, FOY = 2.25 mm).

Figure 1.36. Orangeite, Makganyene. Flow-aligned microphenocrysts of phlogopite (TL, FOY = 2.25 mm).

65

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dral

apa

tite

(w

hite

).

~ gg ::c 5 ~ ~ t::

I o ~ t"l ~ ~

12

0

urn

Fig

ure

1.41

. O

rang

eite

, S

war

trug

gens

. E

uhed

ral

mic

roph

enoc

ryst

s o

f zo

ned

and

Fig

ure

1.42

. O

rang

eite

, N

ew E

land

s. C

hlor

itiz

ed,

dist

orte

d m

icro

phen

ocry

sts

of

man

tled

phl

ogop

ite

(M)

and

euhe

dral

cry

stal

s of

apa

tite

(A

) se

t in

a se

rpen

tine-

<:al

-ph

logo

pite

(M

) se

t in

a m

atri

x o

f ap

atit

e (w

hite

) an

d ca

lcit

e (C

). B

SE

imag

e.

cite

mes

osta

sis

(C).

BS

E im

age.

~

n = ~ i:I:

:I ...

15

0

urn

Fig

ure

1.43

. O

rang

eite

, L

ace.

Mic

roph

enoc

ryst

s o

f ph

logo

pite

(M

) se

t in

a

mat

rix

of s

ubhe

dral

apa

tite

(w

hite

) an

d ca

lcit

e (C

). B

SE

imag

e.

12

0

urn

Fig

ure

1.44

. O

rang

eite

, Bel

lsba

nk. C

hlor

itiz

ed (

dark

ban

ds),

mic

roph

enoc

ryst

s o

f ph

logo

pite

(M

) m

antl

ed b

y te

traf

erri

phlo

gopi

te (

ligh

t gr

ay)

and

subh

edra

l gr

ound

mas

s ap

atit

e (A

) se

t in

a m

esos

tasi

s o

f car

bona

te (

mot

tled

) an

d se

rpen

tine

(d

ark

gray

). B

SE

imag

e.

~ = ~ til > ~ o i ~ ~

Fig

ure

1.45

. O

rang

eite

, S

war

trug

gens

. E

uhed

ral

grou

ndm

ass

apat

ite

(whi

te).

In

clus

ions

(da

rk)

in t

he a

pati

te a

re s

erpe

ntin

e ps

eudo

mor

phs

afte

r an

unkn

own

min

eral

. C =

cal

cite

. M

= m

icro

phen

ocry

stal

phl

ogop

ite.

BSE

imag

e.

Fig

ure

1.46

. D

iops

ide

oran

geit

e, P

ostm

asbu

rg. M

antl

ed m

icro

phen

ocry

sts

of p

hlo

gopi

te (

M)

set

in a

mat

rix

of

diop

side

(d)

, ap

atit

e (w

hite

), c

alci

te (

dark

gra

y),

and

pota

ssiu

m f

elds

par

plus

ser

pent

ine

(bla

ck).

BS

E im

age.

.... <:>

("') ~ ....

Fig

ure

1.47

. D

iops

ide

oran

geit

e, P

ostm

asbu

rg.

Zon

ed a

nd m

antl

ed m

icro

phe

nocr

ysts

of p

hlog

opit

e-te

traf

erri

phlo

gopi

te (m

) an

d pr

ism

atic

dio

psid

e (d

) se

t in

a m

atri

x o

f apa

tite

(w

hite

), c

alci

te (

gray

), a

nd s

erpe

ntin

e (b

lack

). B

SE

imag

e.

Fig

ure

1.48

. D

iops

ide

sani

dine

ora

ngei

te.

Pos

tmas

burg

. E

uhed

raI

crys

tals

of

grou

ndm

ass

diop

side

(D)

wit

h m

antl

es o

f tit

an ia

n ae

giri

ne (l

ight

gra

y ri

ms)

. Als

o pr

esen

t is

sub

hedr

al a

pati

te (

whi

te).

Dar

k gr

ay m

esos

tasi

s is

san

idin

e (S

). B

SE

im

age.

~ t:I:I ~ "-l > ~ o ~ ~ "-l

-..I

-

72 CHAPTER 1

Figure 1.49. Diopside-bearing orangeite, Swartruggens. Typical appearance of a diopside microphenocryst (d) in an unevolved orangeite. M = phlogopite microphenocryst (TL, FOY = 1 mm).

ing widely ranging amounts of microphenocrystal diopside (1-20 vol %). The diopside phenocrysts are typically not zoned and are commonly resorbed. Groundmass potassium feldspar is absent. Diopside is common in many, but not all, dikes in the Winburg, Boshof, and Kroonstad areas. Diopside-bearing orangeites seem to be typically absent from the Sover-Doomkloof, Main Bellsbank, Bobbejaan, and Newlands dikes (Wagner 1914, Dawson et al. 1977, Tainton 1992).

Sanidine diopside orangeite is a rock similar to orangeite as described above but containing less phlogopite and characterized by the presence of abundant euhedral-tosubhedral prisms of diopside (20-50 vol %). The latter may be zoned from colorless diopside toward greenish titanian aegirine and are commonly set in a matrix of potassium feldspar (Figure 1.48). Examples are known from the Voorspoed, Postmasburg, Makganyene, and Prieska region orangeites.

Richterite orangeite is known from Pniel (Tainton 1992) and consists of macrocrystal olivine with reaction coronas of phlogopite which, together with diopside and chromite, are set in a groundmass of phlogopi te-tetraferriphlogopite. In this rock, richterite occurs as anhedral tablets which may subpoikilitically enclose phlogopite and diopside (Figure 1.52). Sanidine is absent from the Pniel samples examined by Tainton (1992) and Erlank (1973). Its absence, and the similarity of the rock to amphibole-free "group 2 kimberIites" in the Barkly West region, led Tainton (1992) to describe the rock as an "amphibole-bearing phlogopite kimberlite."


Figure 1.50. Macrocrystal sanidine orangeite, Besterskraal. Olivine macrocrysts (0) and microphenocrystal phlogopite (M) set in a matrix of apatite (white) and sanidine (black).

Very similar rocks to the above occur at Besterskraal, Sover North, Lace, and Makganyene. In these, richterite also crystallizes after tetraferriphlogopite as a late-stage groundmass mineral. The rocks differ in that the richterite is intergrown with colorless potassium feldspar (Figure 1.53). The latter is commonly replaced by clay minerals and serpentine-like material. Tainton (1992) classified these rocks as "macrocrystal Krichterite phlogopite lamproites," regardless of their association with typical "group 2 kimberlites." In this work, and Mitchell (1994), these richterite sanidine orangeites (see 1.5, 1.12) are regarded as the products of differentiation of orangeite magma and termed evolved orangeites.

Richterite sanidine orangeites from Sover North and Besterskraal commonly contain poikilitic plates of groundmass phlogopite (Figure 1.54). These oikocrysts contain chadacrysts of chromite, diopside, apatite, and subspherical inclusions now consisting of potassium feldspar and/or serpentine-like material. Such poikilitic micas have been found only in the most evolved orangeites which are rich in K-Ba titanates, sanidine, and richterite. The spherical inclusions, on the basis of their morphology, are considered by

74 CHAPTER!

Figure 1.51. Macrocrystal diopside sanidine orangeite, Sover North. Macrocrysts of olivine (0) with phlogopite reaction rims, subhedral diopside (D), and apatite (white) set in a matrix of groundmass phlogopite and sanidine (black).

Tainton (1992) to have been originally leucite. The texture of the rocks is reminiscent of that ofmadupitic olivine lamproites (Mitchell and Bergman 1991, p. 34), and it is in these rocks that the mineralogical convergence toward lamproitic mineral assemblages is best developed (see 1.12).

Photomicrographs illustrating the petrographic characteristics of orangeites are presented in Figures 1.26-1.54. These should be compared and contrasted with the illustrations of kimberlites given in Figures 1.55-1.76.

1.11. PETROGRAPHIC DIFFERENCES WITH RESPECT TO KIMBERLITES

Hypabyssal kimberlites are typically characterized by the presence of macrocrysts and subhedral microphenocrysts of olivine set in a groundmass consisting of spinel, perovskite, monticellite, phlogopite-kinoshitalite, apatite, serpentine, and calcite (Figures 1.55-1.76). Modes vary widely as a consequence of differentiation and alteration. Kimberlites are described according to the abundances of the primary groundmass

Fig

ure

1.52

. R

icht

erit

e or

ange

ite,

Pni

el. G

roun

dmas

s ph

logo

pite

(M

) w

ith

diop

side

and

apa

tite

incl

usio

ns s

et in

a m

atri

x o

f pot

assi

um r

icht

erit

e (R

) (T

L,

FO

V =

I m

m).

~ ~ ~ l'"J

til ~ 1:;1 o ! til -.

.J

VI

76 CHAPTER 1

Figure 1.53. Richterite sanidine orangeite, Besterskraal. Subhedral richterite (R) and diopside (D) are set in a matrix of phlogopite (M) and altered sanidine (S) (TL, FOV = 2.25 mm).

minerals, as described in Section 1.4.4. using the principles established by Skinner and

Clement (1979). Microphenocrystal mica is not a characteristic mineral ofkimberlites, although many

contain minor amounts «1 vol %) of macro crystal phlogopite (Figure 1.56). When mica is present, it occurs primarily as late-stage colorless poikilitic plates and laths of phlogopite-kinoshitalite (Figures 1.66, 1.67, 1.74, 1.75) and less commonly as small groundmass tablets of phlogopite. Although such rocks may be described as phlogopite kimberlites, it is very important to realize that this term is not used in the original sense

of Skinner and Clement (l979), as they did not discriminate between orangeites and

archetypal kimberlites rich in macrocrystal phlogopite. Hence, in their terminology both varieties were termed "phlogopite kimberlite."

Kimberlites differ substantially from orangeites with respect to their overall texture and the type and abundance of primary minerals, as described in Section 1.5 and Table

1.1.

Many kimberlites are characterized by segregation-textured groundmasses (1.9.1.3;

Figures 1.69, 1.71-1.73). Such calcite-serpentine segregations appear to be absent from orangeites. Spinels and perovskites are abundant in kimberlite and coarse-grained relative to those in orangeites (1.5; Skinner 1989). Sheafs of quench-textured apatite are common in kimberlites but not typical of orangeites. In the latter, apatites occur as euhedral prisms


Figure 1.54. Evolved orangeite, Sover North. Groundmass poikilitic phlogopite with spherical pseudomorphs of serpentine which are considered by Tainton (1992) to be formed after leucite. BSE image.

or poikilitic groundmass plates. Atoll-textured spinels (Figure 1.76; Mitchell 1986, pp. 217-219), although common in kimberlites, appear to be absent from orangeites.

The only kimberlites which are petrographically similar to unevolved macrocrystal orangeites are those which are modally enriched in macrocrystal mica, e.g., Tunraq, Koidu, Zagodochnaya, Ngopoetsu, Chicken Park, as a consequence of concentration of this mineral by differentiation processes. These kimberlites typically exhibit other textural and mineralogical features characteristic of kimberlite which serve to set them apart from orangeites, i.e., they are oxide-rich, contain atoll spinels, and have segregation textures while tetraferriphlogopite and diopside are absent.

Kimberlites do not contain richterite, sanidine, or primary diopside and consequently are easily distinguished from evolved orangeites containing these minerals.

Photomicrographs of hypabyssal kimberlites illustrating their characteristic texture and mineralogy are given in Figures 1.55-1.76 and should be compared and contrasted with those of orangeites in Figures 1.26-1.54.

Figure 1.55. Macrocrystal kimberlite, Wesselton Mine. Note the presence of euhedral-to-subhedral primary groundmass olivine. Unresolved dark matrix consists of spinel, perovskite, carbonate, and serpentine. (TL, FOV=4nun).

Figure 1.56. Macrocrystal kimberlite with a phlogopite macrocryst (M), De Beers Mine. Note the presence of abundant small primary olivines and opaque spinels in the groundmass (TL. FOV = 4 nun).

78


Figure 1.57. Macrocrystal kimberlite, Wesselton Mine. Typical appearance of the oxide-rich groundmass of an archetypal kimberlite consisting of subhedral primary olivine (0) and opaque minerals (spinel and perovskite) set in a uniform mesostasis of serpentine and calcite (TL. FOV = 2.25 rum).

1.12. PETROGRAPmC DIFFERENCES WITH RESPECT TO LAMPROITES

Orangeites and macrocrystal orangeites are petrographically so different from most lamproites that they are easily distinguished from the latter. However, evolved orangeites contain minerals (richterite, sanidine, diopside, K-Ba titanates) similar in type and composition to some of those found in lamproites. Thus, considered in isolation and without reference to their consanguineous antecedents, some orangeites might indeed be petrographically classified as lamproites, e.g., previous descriptions of rocks from Sover North (1.5, 1.11; Tainton 1992) or Pniel (Mitchell and Bergman 1991) as "lamproite."

Richterite sanidine orangeites differ texturally from all bona fide extrusive phlogopite lamproites (see Mitchell and Bergman 1991) in that the latter contain abundant microphenocrystal sanidine, diopside, leucite, and only minor olivine. Mineralogical and petrographic similarities are closest with respect to some hypabyssal sanidine richterite lamproites such as are found at Endlich Hill (Wyoming), Cancarix (Spain), and Mount North (Australia). These rocks are characterized by intergrown groundmass sanidine and richterite. However, they typically lack olivine, perovskite, spinel, and calcite and are relatively rich in leucite. Mineral compositions provide further discriminants (see 2.5, 2.8). Further, all evolved orangeites are poor in shcherbakovite and wadeite relative to such lamproites.

The presence of oikocrysts of Ti-phlogopite containing inclusions of (?)pseudoleucite in the Sover North and Besterskraal orangeites further emphasize the affinities of

Fig

ure

1.58

. M

acro

crys

tal k

imbe

rlit

e, W

esse

lton

Min

e. M

antl

ed m

acro

crys

ts o

f ol

ivin

e (0

) to

geth

er w

ith

smal

l su

bhed

ral

prim

ary

grou

ndm

ass

oliv

ines

lac

king

m

antle

s se

t in

an o

xide

-ric

h (w

hite

) se

rpen

tine

cal

cite

mat

rix.

BS

E im

age.

Fig

ure

1.S

? M

acro

crys

tal

kim

berl

ite,

Wes

selto

n M

ine.

Hig

her

mag

nifi

cati

on

BS

E im

age

of t

he g

roun

dmas

s o

f Fig

ure

1.58

sho

win

g su

bhed

ral p

rim

ary

oliv

ines

(P

) se

t in

a m

atri

x o

f spi

nel

and

pero

vski

te (

whi

te),

lat

hs o

f ph

logo

pite

-kin

oshi

ta

lite

(gr

ay),

cal

cite

(gr

ay),

and

ser

pent

ine

(bla

ck).

~

("J ~ ~ ....

Figure 1.60. Groundmass of a serpentine calcite kimberlite, Lipa Pipe, Alakit field, Russia, showing euhedral primary olivine and subhedral opaque spinels set in a mesostasis of serpentine (S) and calcite (C) (TL, FOV = 2.25 mm).

Figure 1.61. Groundmass of an oxide-rich serpentine calcite kimberlite, De Beers Mine, showing subhedral primary olivine (0) within a groundmass of subhedral opaque spinel and perovskite set in a uniform calcite-serpentine mesostasis.

81

Figure 1.62. Groundmass of an oxide-rich serpentine calcite kimberlite, Michigan (U.S.A.), showing primary olivine (0) with a mantle (M) containing rutile and spinel inclusions, together with abundant euhedral opaque spinel and perovskite set in a uniform mesostasis of calcite and serpentine (TL, FOV = 2.25 mm).

Figure 1.63. Groundmass of a perovskite-rich serpentine calcite kimberlite, Wesselton Mine. Perovskites are rounded crystals (P). Unresolved uniform matrix consists of monticellite, calcite, and serpentine. OM = olivine macrocryst; 0 = primary olivine (TL, FOV 2.25 mm).

82

Figure 1.64. Groundmass of a monticellite calcite kimberlite. Pipe 200 (Lesotho), showing small rounded-tosubhedral crystals of monticellite (high relief phase with fluid inclusions), rounded crystals of perovskite (P), and subhedral opaque spinels set in a mesostasis of calcite (TL, FOY = 2.25 rom).

Figure 1.65. Segregation-textured groundmass, Ham kimberlite, Somerset Island (Canada). Segregations consist of resorbed calcite rhombs (high relief) set in a serpophitic serpentine matrix (S). Note the very high concentration of opaque oxides in the silicate oxide portion of the groundmass and their absence in the segregations (TL, FOY = 2.25 rom).

83

Figure 1.66. Oxide-rich serpentine calcite kimberlite. Xi-Yu (China). with poikilitic laths of phlogopite-kinoshitalite (TL. FOY = 2.25 mm).

Figure 1.67. Oxide-rich serpentine calcite kimberlite. Iron Mountain (Wyoming. U.S.A.) with poikilitic laths of phlogopite-kinoshitalite (TL. FOY = I mm).

84

Fig

ure

1.68

. G

roun

dmas

s o

f uni

fonn

-tex

ture

d ki

mbe

rlit

e, W

esse

lton

Min

e.

Anh

edra

l m

acro

crys

tal

and

subh

edra

l sm

alle

r pr

imar

y ol

ivin

es (

dark

gra

y)

are

set

in a

n un

reso

lved

uni

fonn

mix

ture

of

spin

el a

nd p

erov

skit

e (w

hite

),

calc

ite,

and

ser

pent

ine.

BSE

imag

e.

Fig

ure

1.69

. G

roun

dmas

s o

f a

segr

egat

iona

ry t

extu

red

kim

berl

ite,

fro

m t

he

Wes

selt

on M

ine,

sho

wn

at t

he s

ame

scal

e as

Fig

ure

1.68

. T

he g

roun

dmas

s co

nsis

ts o

f ir

regu

larl

y sh

aped

cal

cite

seg

rega

tion

s (C

) w

ithi

n an

oxi

de-r

ich

(whi

te)

unif

onn

serp

enti

ne c

alci

te m

atri

x. A

nhed

ral m

acro

crys

tal a

nd s

ubhe

dral

pr

imar

y ol

ivin

es (

dark

gra

y) a

re a

lso

pres

ent.

BS

E im

age.

~ ~ 5 trJ ~

>

Z

t:l o ~ trJ ~ ~

Fig

ure

1.70

. M

onti

cell

ite

kim

berl

ite,

Wes

selto

n, M

ine.

Sub

hedr

al p

rim

ary

oli

vine

s (d

ark

gray

) ar

e se

t in

a m

atri

x co

nsis

ting

ofm

onti

cell

ite

(gra

y) a

nd s

pine

l (w

hite

). B

SE

imag

e.

Fig

ure

1.71

. S

egre

gati

on-t

extu

red

phlo

gopi

te k

imbe

rlit

e, I

ron

Mou

ntai

n (W

yo

min

g, U

.s.A

.).

Seg

rega

tion

s co

nsis

t of

lath

s o

f phl

ogop

ite-

kino

shit

alit

e (P

) se

t in

a c

alci

te m

atri

x (d

ark

gray

). S

imil

ar l

aths

cro

wde

d w

ith s

pine

l in

clus

ions

(w

hite

) oc

cur

in t

he o

xide

-sil

icat

e-ca

lcit

e un

ifor

m p

orti

ons

of th

e gr

ound

mas

s.

0=

pri

mar

y gr

ound

mas

s ol

ivin

e. B

SE

imag

e.

~

n == ~

Fig

ure

1.72

. S

egre

gati

on-t

extu

red

kim

berl

ite.

Wes

seIt

on M

ine.

Seg

rega

tion

s co

nsis

t of e

uhed

ral-

to-r

ound

ed (r

esor

bed)

cal

cite

cry

stal

s (C

) set

in a

ser

poph

itic

se

rpen

tine

mat

rix

(S).

The

uni

form

por

tions

of t

he g

roun

dmas

s co

nsis

t of o

paqu

e ox

ides

(w

hite

). p

hlog

opit

e-ki

nosh

ital

ite.

cal

cite

. and

ser

pent

ine.

BS

E im

age.

Fig

ure

1.73

. Se

greg

atio

n-te

xtur

ed k

imbe

rlite

. W

esse

lton

Min

e. T

he l

arge

seg

re

gati

on c

onsi

sts

of e

uhed

ral

calc

ite (

C)

and

lath

s o

f st

rong

ly w

ned

phl

ogop

ite-

ki

nosh

itaIi

te s

et i

n a

mat

rix

of

serp

ophi

tic s

erpe

ntin

e (S

). T

he r

emai

nder

of

the

grou

ndm

ass

cont

ains

am

oebo

id c

alci

te-r

ich

segr

egat

ions

. ato

ll-te

xtur

ed s

pine

ls. a

nd

lath

s o

f phl

ogop

ite-

kino

shit

alit

e se

t in

a ca

lcite

ser

pent

ine

mes

osta

sis.

BS

E im

age.

~ ~ == 5 to:!

Vl > ~ o ~ to:! ~ Vl

~

Fig

ure

1.74

. Po

ikil

itic

lat

h of

phl

ogop

ite-

kino

shit

alit

e (P

) en

clos

ing

atol

l-te

xtu

red

spin

els

(whi

te).

Iro

n M

ount

ain

kim

berl

ite (W

yom

ing.

U.S

.A).

BS

E im

age.

10

0

urn

Fig

ure

1.75

. Po

ikil

itic

lat

hs o

f ph

logo

pite

-kin

oshi

tali

te (

P) e

nclo

sing

spi

nels

an

d pe

rovk

site

s (w

hite

). M

esos

tasi

s is

cal

cite

(C

). B

SE

imag

e.

gg

("J ~ ...


Figure 1. 76. Atoll-textured spinels, Wesselton Mine kimberlite. This characteristic texture of many kimberlitic spinels consists of cores of magnesiochromite-magnesian ulvospinel solid solution separated from thin rims of Ti-magnetite by zones of serpentine and calcite. The texture is believed to be formed by the resorption of portions of an original complexly mantled spinel during the later stages of crystallization of the groundmass. BSE image.

some orangeites with lamproites rather than kimberlites. However, these orangeites differ from madupitic lamproites in that the latter are olivine lamproites which do not typically contain sanidine or richterite.

Some evolved orangeites contain olivines which exhibit complex parallel growth (or resorption) morphologies, e.g., Sover North, Bellsbank West Fissure (Tainton 1992). The habit is similar to that of olivines found in some olivine lamproites. However, this morphological feature cannot be used to assert that the rocks are lamproites because it has no genetic significance, i.e., similar olivines are found in rocks derived from other magmas types, e.g., melilitites (McIver 1981, Moore and Erlank 1979).

Evolved orangeites are taxonomically challenging rocks in that they cannot be classified without recourse to detailed mineralogical (and geochemical) investigations. They can only be unambiguously recognized as belonging to the orangeite clan by their association with less-evolved consanguineous rocks. They demonstrate the problems of

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identifying rocks by simple petrographic methods using only isolated specimens. Richterite sanidine orangeites illustrate the mineralogical convergence of the orangeite clan with rocks of the lamproite clan that results from differentiation of similar but genetically different potassic peralkaline magmas. The relationships of orangeites to lamproites are discussed further in Section 4.7.2.

Kimberlites, Orangeites, and Related Rocks || Kimberlites and Orangeites

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