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Geochemical signatures of polygenetic origin of a banded iron formation (BIF) of the Archaean Sandur greenstone belt (schist belt) Karnataka nucleus, India
Precambrian Research, 61 ( 1 9 9 3 ) 137-164
137
Elsevier Science Publishers B.V., A m s t e r d a m
Geochemical signatures of polygenetic origin of a banded iron formation (BIF) of the Archaean Sandur greenstone belt (schist belt) Karnataka nucleus, India C. Manikyamba, V. Balaram and S.M. Naqvi National Geophysical Research Institute, Hyderabad 500 007, India (Received April 25, 1991; accepted after revision June 30, 1992)
ABSTRACT Manikyamba, C., Balaram, V. and Naqvi, S.M., 1993. Geochemical signatures of polygenetic origin of a banded iron formation (BIF) of the Archaean Sandur greenstone belt (schist belt ) Karnataka nucleus, India. Precambrian Res., 61 : 137-164. Chert, ferruginous cherts, cherty banded iron formations (BIF), shaly BIF and shales are found interbedded at the eastern part of the Sandur greenstone belt. Cherts and shales are two end members in which deposition of Fe203 and A1203 in variable proportions has given rise to observed large-scale variation in major, trace and REE abundances. Iron, silica and REE have precipitated from the cold marine water which was enriched in these components by hydrothermal solutions from MOR vents. However, some minor amounts of Fe, Si and REE appear to have been brought in as dissolved load by rivers. Scatter in K, Mg, Ti, Cr, Ni, Zr, Hf, V, Sc, Y, Rb, Sr, Nb and Ta abundances and ratios indicate that the clastic components have both volcanoclastic and terrigenous sources. Large-scale variation in sumREE, moderate to pronounced La enrichment, positive Eu and negative Ce anomalies are characteristic of the majority of the samples of CBIF and SBIF. REE pattern shapes of CBIF are similar to hydrothermal solutions from EPR or Rs vents. Mixing of the FeO + SiO2 rich hydrothermal solutions with ambient sea water and clastic input has resulted in an observed large variation in LREE/ HREE ratio. However, total population of the samples shows the simultaneous increase of LREE and HREE from cherts to shales through CBIF and SBIF, and obliteration of hydrothermal signature in La content and the magnitude of the positive Eu anomalies. Negative Ce anomalies appear to be the result of a reaction with photosynthetic oxygen produced by bacteria which have developed stromatolites in the underlying formation. Hydrothermal and fluvial solutions provided Si, Fe and trace elements including REE to the cold marine waters of the seas. These hydrothermal solutions were emplaced in the relatively deeper part of the basin having a reducing and neutral to alkaline environment; and due to the thermal and chemical gradient convected towards the shoreline where photosynthesis was producing 02. Here, on the stable shelf region below the wavebase and photic zone FeO and Ce 3+ of these solutions were oxidized and mixed with clastic material of divergent origin. The Ce 4÷ was precipitated in the varves like SBIF and shales. Near the shoreline the environment was intermittently oxidizing but acidic most of the time and thus the precipitation of silica took place continuously, whereas precipitation of iron occurred intermittently due to the intermittent availability of oxygen or FeO or both. Our observations suggest that the BIFs of the Sandur belt are a product of hybridity between the hydrothermal emplacement of Si and Fe and divergent clastic sediments to ambient cold ocean water. The precipitation of Fe203 was biogenically mediated; a model combining these processes explains most of the feature of BIFs of the Sandur schist belt.
Introduction Banded iron formations (BIFs) are important but enigmatic rocks of the early and midCorrespondence to: S.M. Naqvi, National Geophysical Research Institute, H y d e r a b a d 500 007, India.
dle Precambrian. Their virtual confinement (James, 1983) to the Precambrian and absence of modern analogues have made these rocks extremely useful for the understanding of the ancient exogenic processes. It cannot be denied that if the early atmosphere was anoxic (Walker et al., 1983; Kasting, 1987)the FeO,
0 3 0 1 - 9 2 6 8 / 9 3 / $ 0 6 . 0 0 © 1993 Elsevier Science Publishers B.V. All rights reserved.
138 S i O 2 and some minor and trace element content of the BIF had been accumulating in the early Precambrian ocean reservoir for more than a billion years and the hydrothermal input seen as a source for BIF deposition may not be contemporaneous to the deposition. However, faster completion of the Wilson cycle in the Archaean as proposed by Hargraves (1986) suggests that the major part of the FeO and SiO2 content of BIF of a particular basin may be contemporaneous or it could have been carried over to another basin after the closure of an earlier one. Their origin remains highly controversial (Simonson, 1985). Although they have long attracted attention and a wealth of information is available on their associations, tectonic settings, petrology, mineralogy and economic potential, geochemical studies have been taken up only during past the two decades (Davy, 1983; Fryer, 1983; Dymek and Klein, 1988; Beukes and Klein, 1990; Derry and Jacobsen, 1990). Oxygen, carbon and sulphur isotopic data on the Archaean BIF and associated shales and carbonates are extremely scarce (Perry, 1983, Goodwin et al., 1985; Beukes et al., 1990; Schidlowski, 1990). Recent efforts in this field emphasise that a much larger and detailed geochemical and paleobiological data base should be created (Walter and Hofmann, 1983) for minimising the speculative part of the proposed genetic models. Sometimes "pure" samples made up only of chert and iron minerals are chosen for geochemical studies and the associated "clastics", regarded as "contaminants" are ignored (e.g. Derry and Jacobsen, 1990). The pure chemical fraction of BIF (i.e. the chert and iron minerals) has features which render it a poor recorder of minor element trends. Therefore, if an imprint of a genetic process is to be detected it will be preserved in the so-called "contaminants", which are an integral part of the BIF sequences. Although it is generally agreed that BIFs are chemogenic sediments, it is not certain what the source of Fe and Si was and how the ox-
c. MANIKYAMBAET AL.
ides, carbonates, sulfides and silica were precipitated. Other questions concern the mode of solution and transport of Fe and Si; the source of 02 needed for precipitation of Fe; and how microbands were formed. Are bacteria involved in producing these microbands (Nealson and Myers, 1990)? BIFs of several basins have been studied in detail, addressing these and some other questions, but the processes that form BIFs still remain somewhat elusive; genetic models range from exhalative (Gross, 1980; Simonson, 1982; 1985; Goodwin et al., 1985), evaporative (Eugster and Chou, 1973; Trendall, 1973; Garrels, 1987), biologically mediated precipitation (Cloud, 1973; La Berge, 1986; Nealson and Myers, 1990), and ocean upwelling (Holland, 1973; Drever, 1974 ). During the past two decades a number of summaries of the work carried out on BIF have been published (Trendall and Morris, 1983 ). In these summaries and interpretations most of the data used are from the BIFs of early Proterozoic basins although the occurrence of BIF in Archaean greenstone belts is a worldwide phenomenon (Condie, 1981). The Archaean BIF of greenstone belts of Canada (Goodwin, 1973 ), India (Radhakrishna et al., 1986), China (Zhai and Windley, 1990) and Africa (Anhaeusser, 1990; Wilson and Nutt, 1990) are relatively less studied than those of early Proterozoic age. The relevant information on the BIFs of India is extremely meagre, and especially the geochemical and genetic aspects have seldom been attempted. BIFs in India are confined only to Archaean greenstone belts. In spite of their wide development, Proterozoic sedimentary basins are generally devoid of BIFs (Radhakrishna et al., 1986). Attempts have begun recently to understand the geochemistry and genesis of the Archaean BIFs (Majumdar et al., 1982; Chakraborty and Majumdar, 1986; Devaraju and Laajoki, 1986; Manikyamba, 1992; Khan et al., 1992; Rao, 1992 ). In the Dharwar craton, deposition of BIFs has taken place at
BIF OF THE ARCHAEANSANDURGREENSTONEBELT
139
five stratigraphic horizons (Fig. 1 ) in strata formed between 3.5 (?) and 2.6 Ga ago (Naqvi et al., 1988 ). The most extensive and thick horizon is developed in a sedimentary sequence of stable shelf environment between 3.0 and 2.8 Ga ago (BIF3 in Fig. 1); generally known as the Bababudan Group which forms, the lower division of the Dharwar Supergroup (Swami Nath and Ramakrishnan, 1981 ). These BIFs are associated with QPC (quartz pebble conglomerate), current bedded-ripple marked quartzites and carbonaceous phyllites, phyllites and basic-acidic volcanics. They exhibit classical meso- and microbanding and thin in-
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terbedding with several types of extremely fine grained volcaniclastic and terrigenous sediments. The rock suite forming the Sandur schist belt are geologically correlated with Bababudan Group (Radhakrishna, 1983 ). It contains very large amounts of BIF and BMF (banded manganese formations) which are economically exploited on a large scale. These BIFs and BMFs were deposited under stable shelf conditions and are very closely associated with stromatolites and microbiota bearing cherts (Murthy and Reddy, 1984; Naqvi et al., 1987; Venkatachala et al., 1990, 1991 ). In this re-
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Fig. 1. Model stratigraphic sequence of greenstone belts (schist belts) in the Karnataka nucleus,India showing five successive cycles of BIF deposition in older and younger schist belts. The Sandur BIFs represent BIF3 which are found in stable shelf association along with stromatolites. (Modified after Naqvi et al 1988. )
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BIF OF THE ARCHAEAN SANDUR GREENSTONE BELT
gion a transition from stromatolitic carbonates to GIF (granular iron formation) to BIF is well exposed in the western part of the boatshaped Sandur schist belt (Fig. 2). The basal strata are not repeated on the eastern part where maximum development of BIF is seen. We present the results of our geochemical studies on the BIFs of the eastern part of the Sandur schist belt to elucidate their origin and to infer the physico-chemical conditions which prevailed during the deposition.
Summary of the geology of the Karnataka nucleus The geology of the Karnataka nucleus has been summarised by several workers (Swami Nath and Ramakrishnan, 1981; Radhakrishna and Naqvi, 1986; see also Naqvi and Rogers, 1987 ). The sequences are classified into older and younger schist belts ranging in age from 3.5 (?) to 2.5 Ga. Older and younger schist belts are separated by a widespread crust forming event at or around 3.0 Ga represented by the final consolidation of the Peninsular gneiss (Swami Nath and Ramakrishnan, 1981). Older schist belts are largely made up of mafic and ultramafic (spinifex textured komatiites) volcanics, amphibolites, quartzites (meta cherts), metapelites, paragneisses, and finegrained argillaceous and volcanogenic sediments including thin bands of recrystallized garnet-bearing BIFs. These belts are exposed near Holenarasipur, Nuggihalli, Krishnarajpet, Hadnur and several other places and have been metamorphosed to amphibolite facies (Naqvi and Rogers, 1983, 1987). They generally exhibit a complex deformational history (Naha et al., 1986). Sequences in the younger schist belts start with an unconform-
141
ity represented by quartz pebble conglomerate (QPC), current-bedded and ripple marked quartzites interbedded and overlain by basic volcanics, phyllites of variable composition with or without carbon and rare carbonate. They are followed by a thick sequence of BIF (and BMF at some places) with interbedded shales/phyllites. These basal portions of the younger schist belts are designated as the Bababudan Group, which is followed by the Chitradurga Group, made up of oligomictic or polymictic conglomerates, arkoses, quartzites, stromatolitic carbonate rocks, graywackes, basic-acidic volcanics and BIFs (oxide, carbonate and sulphide facies). This group is intruded by 2.6 Ga old K-granites. The entire succession of the Dharwar schist belts is not exposed at any single locality and shows an interfingering relation with polyphase gneisses (Naqvi, 1981 ). This aspect has given rise to many controversies about the structure and stratigraphy of the schist belts. Several conflicting models of stratigraphy have been proposed and the correlation of rock suites of various belts is often subjective due to discontinuous outcrop and scarcity of radiometric data. Naqvi et al ( 1988 ) proposed (Fig. 1 ) that at least five horizons of BIF deposition may be recognized. The maximum and most extensive deposition of BIF took place in the Bababudan Group around 3.0-2.8 Ga. The BIF of this Group can be subdivided into two classes: (1) predominantly oxide facies BIFs of Bababudan, Kudremukh and Kushtgi schist belts, lacking stromatolitic carbonate rocks, Mn formations, widespread carbonaceous phyllites and sulfide facies BIFs; and (2) the BIFs of the Bababudan Group in the Chitradurga, Shimoga and Sandur schist belts which are associated with stromatolitic carbonates (Ven-
Fig. 2. Geological map of the Sandur schist belt including stratigraphic sequence. Inset (A) shows location of the Sandur schist belt with respect to Karnataka, India. (B) Detailed geological map of the Sandur schist belt. The transition from quartzite-stromatolite sequence to BIF is exposed at the eastern limb (simplified and modified after Roy and Biswas, 1983). (C) Location of the samples analysed from exploratory bore hole cores and railway and road cuttings. Location of stromatolites, microfossils and carbon phyllites is also shown. (D) Map of a strip giving details of transition zone.
142
katachala et al., 1991 ), Mn formations and extensive carbonaceous phyllites. Earlier, Srinivasan and Sreenivas ( 1972 ) designated it as a separate Dodguni Group. The BIF of different regions mentioned above show relative abundances of different facies (Manikyamba, 1992; Rao, 1992; Khan et al., 1992). The BIF of Chitradurga for example are made up of oxide, carbonate, sulfide and mixed carbonate sulfide facies, whereas BIF of Bababudan and Kudremukh are predominantly of oxide facies and mixed oxide silicate facies.
Geology of the Sandur schist belt and depositional environment
Distribution of rock types, stratigraphy, structure and metamorphism of the Sandur schist belt has been worked out in detail by Roy and Biswas ( 1979, 1983 ). Geological map and stratigraphic succession after Roy and Biswas (1983) is given in Fig. 2B. The stratigraphic sequence in the belt is divided into 4 formations: Yeshwanthanagar, Deogiri, Donimalai and Nandihalli. The Yeshwanthanagar Formation consists of amphibolites, metapyroxenites, metagabbro, quartzite and quartz-mica schist. The Deogiri Formation is believed to conformably overlie the Yeshwanthanagar Formation (Roy and Biswas, 1983). It is essentially a sedimentary sequence of manganiferous graywacke-argillite with bands of quartzite and arkoses with siliceous dolomite (stromatolitic) and banded ferruginous chert (BIF). The Donimalai Formation is made up of banded magnetite/haematite, chert/jasper and metavolcanics ranging from basic to acidic, metagabbros, graywackes, fuchsite-quartzite, carbonaceous schists, metapelites, and sulfide facies BIF. The youngest formation, Nandihalli, is characterized by the absence of BIF and is made up of metabasalts and metagabbros, acid volcanics and a graywacke-argillite sequence (Roy and Biswas, 1983 ).
C. MANIKYAMBA ET AL.
Transition from quartzite to BIF The transition from quartzites (ripple marked Mn-bearing sandstones) to BIF through stromatolitic carbonate rocks, Mn formations and GIF is found at the interface of the Deogiri and Donimalai Formations. Stable shelf strata of the Deogiri Formation gradually merge into relatively deeper water shelf microbanded strata of the Donimalai Formation in which volcanic rocks are relatively more abundant (Naqvi et al., 1992). Gradation and development of microbanded BIF and GIF at 5 km stone on Yeshwanthanagar to Sandur road, and deposition of graywackes indicates a relatively deeper environment of deposition than the strata having current bedded and ripple marked quartzites and stromatolites of small dimensions exposed about 5 km west of this outcrop (Fig. 2). Stromatolites are generally not found below the photic zone and ferythmites are not formed above the wave base. Graded bedding in GIF is indicating fluctuations of the wave base. These structures are clear indications of variation in the depth of deposition. The strata with its facies changes is believed to be folded (Fig. 2 ) into a synform (Roy and Biswas, 1983 ). On the western limb of this synform predominance of stable zone strata is seen whereas on the eastern limb BIF, GIF, interbedded with shales/phyllites graywackes and metavolcanics (all part of a relatively deeper shelf) dominate. Mn formations, quartzites, carbonates are not repeated on the eastern part (Fig. 2 ). It is in this eastern part, that both oxide and carbonate facies BIF are abundant along with a zone of carbonaceous phyllite. At a few places along this part the zones of sulfide (pyrite) facies and mixed sulfide (pyrite), oxide (jasper, haematite) facies or mixed sulfide (pyrite), carbonate (siderite) facies BIF are found. All along the eastern part, the BIF, both the cherty and shaly type, are interbedded with pillowed metavolcanic rocks. The Donimalai Formation appears to be a relatively deeper facies variant of
143
BIF OF THE ARCHAEAN SANDUR GREENSTONE BELT
the Deogiri Formation. The transition from quartzite to BIF occurs within a sequence that from the base upward grades from quartzites to stromatolitic dolomites and finally to BIF through Mn dolomite, Mn shales, Mn sandstones and GIF. Contacts between quartzites and stromatolitic/Mn dolomites, between dolomite and shales and sandstones are gradational, and between shales, sandstones and BIF vary from gradational to sharp. Within the first layer of BIF, which conformably rests over sandstone and shales, GIF and BIF are found exhibiting both vertical and lateral transition into each other (Naqvi et al., 1992 ). Grading of iron minerals in the chert is seen in thin sections. Gradual settling of the fine dust of haematite-magnetite and sometimes siderite in the fine grained quartzite/chert matrix is exhibited until a layer of coarse iron minerals is formed with sharp contact with another graded layer. Such sequences are repeated several times. Stromatolites of the Sandur schist belt are oval to elliptical in shape with convex laminae (see Venkatachala et al., 1991 ). The larger axis of the ellipsoids at the bedding plane varies from 2 to 20 cm and the length of the columns in sections perpendicular to the bedding plane varies from 5 to 15 cm. The carbonates of stromatolites are Fe-dolomite, siderite and Mnsiderite interbedded with chert. The sedimentary structure in the underlying quartzite indicates that these and the carbonates in IF ( iron formation) were deposited under passive shelf conditions and that some IF, especially of the eastern part of the Sandur belt, were deposited at a deeper shelf below the wave base. An exposed cratonic area, allowing for clastic input in the form of fine grained mature arenites, graywackes and highly aluminous shales, indicating deep and intense weathering, probably existed west of the Sandur belt. Similar lateral and vertical transition from current-bedded and ripple marked quartzites to oxide and carbonate facies BMF and BIF, through stromatolitic dolomite and Mn shales is seen north of
the Srinkatte dome and west of Bhimasamudra in the Chitradurga schist belt. This transition from quartzites to BMF and BIF, through dolomitic stromatolites, Mn shales and carbonaceous phyllites (shales) is also exposed in the Shimoga schist belt near the Kumsi manganese mines and Kalche (Naqvi et al., 1992 ). A stable shelf environment of deposition is clearly evident from the morphology of stromatolites and other sedimentary structures in these localities (see Srinivasan et al., 1989, 1990). Furthermore, the discoveries of stromatolites in the basal sections, grading into BMF and BIF, indicate extensive Archaean biogenic activity in Karnataka nucleus and therefore biogenic mediation for the precipitation of iron and Mn in the overlying formations. Earlier such a transition between stromatolitic carbonates and BIF has been described from the early Proterozoic Transvaal basin of South Africa, namely the Cambellrand-Kurman transition zone. From a sequence correlated with the Transvaal basin in Botsvana/Zimbabwe also Mn bearing stromatolitic carbonates have been described from Mn deposits (Litherland and Malan, 1973; Beukes 1983). Transitions between GIF (Granular iron formation ) and BIF are a common phenomenon in the early Proterozoic stratigraphic record. However, so far they have not been reported from Archaean greenstone belts (Gole and Klein, 1981 ).
Age metamorphism and deformation The Sandur schist belt is intruded by Closepet Granite, which is locally known as Torangal Granite. Closepet Granite is dated at 2.42.6 Ga and Torangal Granite has yielded a RbSr isochron of 2.4 Ga. The basic and acid volcanic suites of the Sandur belt have yielded a thermal resetting Rb-SR isochron age of 2.4 Ga (Bhaskar Rao et al., 1992). The basic volcanics of the Kudremukh belt (which are stratigraphically equated with the basic volcanics of the Sandur belt) have yielded a Sm-Nd age
144
of ~2.9 Ga (Drury et al., 1983). The gneissose basement on which this belt is resting is dated elsewhere at 3.1 Ga (Taylor et al., 1984). In view of the available radiometric age data, the age of the Sandur schist belt may be somewhere between 3.1 and 2.6 Ga. However, the exact radiometric age for the rock formations of this belt is not yet known. Efforts in this direction are being continued at NGRI. The Sandur schist belt has been deformed twice, represented by two distinct phases of (folding) deformation, designated as SD 1 and SD2, of which the earlier one SDI is quite pronounced producing the NW-SE trending regional synformal structure of the schist belt (Fig. 2 ). SD 1 has resulted in the development of macroscopic folds, tight to isoclinal mesoscopic folds, schistosity, mineral lineation, fold mullion and pebble lineation. Second generation S D 2 deformation gave rise to broad warps, tight mesoscopic folds and crenulation cleavage (Roy and Biswas, 1983). This interpretation has been recently disputed by Mukhopadhyay and Matin (1990). This belt is predominantly characterized by greenschist facies of regional metamorphism, which increases to amphibolite facies towards the outer margins of the schist belt.
Characteristics of the Sandur belt BIF Iron formations of the Sandur belt consist of cherts (C), ferruginous cherts (FC), cherty BIF (CBIF), shaly BIF (SBIF), ferruginous shales (FSH) and shales/phyllites (SH). In fact there are two end members namely cherts (C), and shales/phyllites (SH), in which a variation in the proportion of SIO2, Fe203 and A1203 bearing minerals has produced the observed compositional scatter, as has been earlier found by Beukes (1980) for the BIF of the Transvaal Basin. The Donimalai BIF from the eastern part of the Sandur belt (Fig. 2 ) is made up of meso- and microbands of chert and iron minerals. Iron oxides (haematite/magnetite) dominate over carbonates (siderite/anker-
c. MANIKYAMBAET AL.
ite). Sulfide (pyrite) and silicate (cummingtonite/grunerite) bands are also found but they are not as abundant as the iron oxide bands. Two types of banding are noticed, one in which along with the chert a great concentration of iron minerals is found. In this type the Fe and SiO2 rich layers are microlaminated. The other type is one in which laminated chert and carbonates with sporadic development of magnetite/haematite is found. In such cases a gradual change from an iron-rich layer to a chertrich layer is found. The individual thickness of the iron-rich laminae with abrupt change to chert varies from a few mm to 2-4 cm. The thickness of the laminae showing a gradual decrease of iron minerals into chert laminae are about 5-10 ram. Fine dust of iron oxide is seen gradually diminishing into pure chert. Irregular, wavy bedding and domal structures identical to stromatolites are very common. "Neither stromatolites nor convincing microfossils are known from Archaean IF" (Walter and Hofmann, 1983). However, Hofmann et al. ( 1991 ) have found unequivocal stromatolites from the carbonate facies BIF of the Michipicoten formation, Ontario, Canada. SBIF are generally microbanded with chert, iron mineral and shales. At several places due to interbedding of chert and iron minerals with shales, high order fissility has developed.
Sampling and analytical techniques Most of the samples studied come from exploratory bore hole cores of the National Mineral Development Corporation or from the recent railway or road cutting in the mining area. The location of the samples is shown in Fig. 2C. Special care has been taken to collect and study only fresh unaltered samples. One hundred and two samples of the oxide facies BIF and associated cherts, ferruginous cherts, shaly BIF and shales were chosen. Ore microscopy for opaques and thin section study for silicate minerals were carried out for all samples. Composition of the constituent minerals have
145
BIF OF THE ARCHAEAN SANDUR GREENSTONE BELT
been determined by an electron microprobe analyser. Major elements were determined on fused pellets by a Phillips 1400 XRF. Trace elements and REE (for 102 samples) were estimated on ICP-MS and AAS. Fe203 in all samples was estimated on AAS and by colorimetric method. Among trace elements Zr and Y were determined on XRF. For details of the analytical techniques used and precision of the estimates refer to Govil (1985) and Balaram et al. (1991).
Petrology and mineralogy BIFs of the Donimalai region are characterised by simple mineralogy. CBIF consists of microcrystalline quartz/chert, haematite, specularite, with minor cummingtonite and grunerite. Magnetite is also found in some microbands. Hornblende is restricted to mixed silicate-oxide BIFs. Haematite and jasper are often present as ultrafine spheroids resembling coccoids (10 # m and less). Haematite and magnetite are also found as tetrahedral and cubic crystal clusters in the layers. Needle shaped haematite is seen in a few sections. In some samples veins of pyrite and chalcopyrite are present. Carbonates such as dolomite, ferroan dolomite, siderite, magnesian siderite, magnesian dolomite are found in a few sections in minor quantity. SBIFs are composed of chert and shaly layers with iron minerals. In addition to chert, haematite _ magnetite and jasper, they contain kaolin, sericite, and chlorite in varying proportions. In these formations minor carbonate is present. Ferruginous shales are made up of mainly kaolin _+ sericite, chlorite and iron oxides. Extremely fine grained and finely laminated varved types of shales/phyllites chiefly composed of kaolin are found interbedded with both types of BIFs. Some shale samples are thinly laminated and rhythmically bedded with chert having very little or no iron oxide, carbonate or sulphide minerals. In the iron ore zone of Donimalai where supergene enrichment has occurred and mining is in
progress, haematite, magnetite, specularite, martite, goethite and lepidocrocite are found. Syngenetic iron-rich bands are mainly made up of haematite and often magnetite. Haematite and chert appear to be primary minerals. Siderite occurs as primary and diagenetic varieties. Diagenetic siderite is interstitial and is often found as well developed rhombohedras. Magnetite appears to be of metamorphic origin. Details of mineralogical work form a separate paper (Manikyamba, 1992). Specularite, martite, goethite, limonite and lepidocrocite, etc. are secondary minerals. Samples containing secondary minerals were not included for petrological and geochemical studies.
Geochemistry Behaviour of major elements Mean and standard deviations of major, trace, and rare earth element contents of BIF and associated chert, shale/phyllites are given in Table 1. There is a large scale variation in the SiO2/Fe203 (T) molar composition. On the basis of SiO2, A1203 and Fe203 content the samples analysed are grouped as cherts (C), ferruginous cherts (FC), cherty BIF (CBIF), shaly BIF (SBIF) and phyllite/shales (SH). Samples containing less than 15% Fe203 and 80-75% SiO2 are classified as cherts. Ferruginous cherts contain 5-14.9% Fe203 and rest SiO2; whereas cherty BIF contains SiO2 and more than 15% Fe203; shaly BIF contains more than 15% Fe203 and more than 2% A1203; Shales (phyllites) contain less than 5% Fe203 and more than 5% A1203. A clear distinction between CBIF and SBIF can be made on the basis of A1203 content (Dimroth, 1986). Almost all major element constituents exhibit a large scale variation in their abundances in SBIF, whereas CBIF for a major part is made up of SiO2 and Fe203(T). The alkali content of both the varieties, with the exception of a few samples of SBIF, is lower than 1.0%.
146
C. MANIKYAMBAET AL.
TABLE 1 Average chemical composition of BIF from the Sandur schist belt Cherts (7) Mean
FC (8) SD
Mean
CBIF (51) SD
SBIF (27)
Mean
SD
Mean
Shales ( 9 ) SD
Mean
SD
Major elements (wt%) SiO2 TiO2 A1203 Fe203 MnO MgO CaO Na20 K20
P205
96.71 0.01 0.02 2.50 0.04 0.27 0.14 0.02 0.11
1.31 0.01 1.52 0.05 0.43 0.31 0.01 0.09
86.36 0.01 0.05 11.01 0.10 0.16 0.12 0.05 0.01 0.11
4.07 0.05 4.52 0.11 0.13 0.26 0.06 0.02 0.11
51.80 0.03 0.23 44.30 0.09 0.50 0.09 0.19 0.07 0.08
20.25 0.03 0.29 21.30 0.14 0.70 0.26 0.23 0.35 0.10
41.84 1.77 17.47 29.14 0.06 2.03 1.11 0.28 1.08 0.14
11.33 1.74 7.35 13.93 0.07 4.00 1.61 0.44 1.70 0.19
62.18 0.88 17.63 8.09 0.08 1.46 1.31 0.75 2.16 0.06
9.75 0.76 8.40 3.90 0.13 2.06 2.50 1.00 1.87 0.04
Trace elements (ppm) Sc V Cr Co Ni Cu Zn Rb Sr Y Zr Nb Ba Hf
0.16 7.53 11.40 98.56 7.53 23.75 3.76 0.07 1.61 0.46 0.71 0.15 1.08 0.09
0.11 11.86 12.09 15.99 3.61 41.78 3.96 0.04 1.63 0.37 0.47 0.06 0.72 0.10
0.38 2.36 5.48 73.01 10.73 35.54 3.55 0.18 0.54 0.67 0.92 0.22 3.61 0.12
0.36 4.52 5.10 33.02 5.30 81.96 2.25 0.12 0.43 0.39 0.70 0.13 3.79 0.08
1.30 11.28 55.74 33.72 10.23 12.51 11.21 2.66 3.78 1.78 4.01 0.29 14.55 0.18
1.40 16.62 82.00 28.54 5.70 16.12 10.47 11.92 5.41 1.72 9.05 0.25 26.32 0.28
20.58 210.20 276.01 16.79 82.91 76.13 59.89 40.89 43.09 29.21 181.16 6.27 172.20 3.64
8.83 145.34 187.84 19.58 106.90 66.92 60.96 77.12 49.52 26.99 164.58 5.03 268.99 3.40
23.82 307.79 354.94 24.70 96.02 149.66 98.82 86.52 101.53 19.50 258.73 9.73 249.73 3.75
13.37 220.20 368.35 16.84 81.41 156.12 88.94 96.01 202.78 14.58 262.65 8.05 221.22 2.17
La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu ~REE
3.57 1.14 0.66 5.79 0.49 0.16 0.16 0.03 0.53 0.10 0.29 0.03 0.15 0.02 13.32
4.80 1.13 1.41 9.13 0.89 0.23 0.07 0.02 0.48 0.08 0.21 0.01 0.07 0.02 16.18
3.66 0.97 0.10 0.24 0.14 0.09 0.08 0.04 0.19 0.04 0.15 0.02 0.14 0.02 6.11
9.65 1.47 0.15 0.17 0.13 0.06 0.04 0.06 0.14 0.03 0.10 0.02 0.10 0.02 12.21
5.97 2.10 0.26 1.30 0.34 0.18 0.26 0.14 0.37 0.07 0.25 0.03 0.19 0.03 11.42
9.41 1.92 0.26 1.36 0.32 0.13 0.26 0.16 0.28 0.05 0.35 0.02 0.14 0.02 10.70
20.29 39.11 3.78 20.39 4.13 1.51 3.77 0.64 4.07 0.89 2.73 0.38 2.68 0.58 96.78
22.39 40.88 3.82 22.32 4.04 1.55 3.06 0.71 4.60 0.94 2.72 0.37 2.95 0.65 107.37
14.49 38.33 2.73 9.35 2.20 0.73 2.22 0.40 2.22 0.47 1.49 0.21 1.48 0.24 74.82
9.68 30.09 1.87 6.12 1.55 0.48 1.39 0.24 1.43 0.27 0.77 0.10 0.81 0.17 49.80
( 1 ) Number of samples for each type analysed is given in parentheses. (2) Data for the samples will be made available on request.
However, MgO and CaO are fairly abundant and show great scatter due to the presence of carbonates (Table 1 ). A1203 appears to be one of the most diagnostic constituents o f the dif-
ferent varieties of IF in the belt. T i O 2 also is found to distinguish between the two main varieties of the BIFs. MnO and P205 are of very low abundance and exhibit an overlap be-
BIF OF THE ARCHAEAN
SANDUR
GREENSTONE
147
BELT
elevated to more than 0.4% and 0.2% respectively (Fig. 3A). Even at elevated concentrations of TiO2 and A1203, as in SBIF, the scatter in TiO2 and A1203 concentrations shows that the clastic input for different layers and at different levels has been inconsistent and may probably represent different types of source rocks. Thus the A1203 and TiO2 abundances in Sandur BIFs indicate that clastic sedimentation from variable compositional sources has also occurred simultaneously with chemical precipitation of SiO2 and Fe203, and the quantity and quality of this clastic sedimentation has been responsible for the observed compositional variation in both types of BIF. The
tween the two varieties (Table 1 ). Elevated A1203 and TiOe reflects a clastic component (Ewers and Morris, 1981 ). Most of our CBIF samples have 0.01 to 0.1% TiO2 with A1203 values ranging between 0.01% and 1.35%. In SBIF both the TiO~ and A1203 values are elevated from 0.2 to 8.5% and from 3.4% to 33% respectively. CBIF and SBIF form seperate groups, though the TiO~/A1203 ratio for both types of BIF is always less than one (Fig. 3A). Linear relationships between A1203 and TiO~ of S macrobands has been observed by several workers (Ewers and Morris. 1981). In our samples we observe crude linearity, but only after the concentration of TiO2 and A1203 are 100
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148
scatter in the TiO2/AI203 ratios (Fig. 3A) and the composition of the thinly laminated varvelike shales interbedded with oxide facies BIF at Donimalai, containing more than 15% A1203 due to the presence of kaolin, clearly indicate that terrigenous sedimentation continued with chemical precipitation. In thick layers of shales interbedded with CBIF several sedimentary cup-like structures, are present indicating deposition of shales on shallow shelves as such cupsuate structures are formed by fluctuations of the wave base (Naqvi et al., 1992). The TiO2/Zr of the shales and SBIF are similar to sedimentary clays and slates. Higher TiO2/ A1203 ratios in some of our samples also indicate that in addition to terrigenous sedimentation mafic volcanic activity also contributed to the shaly part of the BIF. The MgO content of the SBIF has a bimodal distribution. Twenty-two samples have a MgO content between 0.01 to 0.86% whereas five samples have from 2.9 to 13.4% MgO. A considerable amount of MgO is in the form of ferromagnesian minerals like chlorite and amphiboles contributed from basic volcanism, probably volcanic ash. Generally samples of SBIF having very high A1203 contents have low MgO abundances and vice versa (Table 1 ).
D&tribution and behaviour of trace elements Data for fourteen trace elements are given in Table 1 indicating a large scatter in their values as reflected by their standard deviations. The chromium content of both types of BIF and associated shales/phyllites generally varies from 4 to 400 ppm and their Ni abundances range from less than 2 ppm to 88 ppm. In one sample of shale Cr is found at 976 ppm level. A wide scatter of Ni and Cr contents (r=0.63) of these rocks is illustrated in Fig. 3B, where it can be observed that on the basis of Ni and Cr abundances and their ratios, no distinction can be made between CBIFs, SBIFs and shales. Even cherts and ferruginous chert, exhibit fairly high Ni and Cr abundances, es-
C. MANIKYAMBA ET AL.
pecially chromium. Ferruginous sediments, formed by chemical precipitation from normal seawater do not have such high concentrations of these elements. Therefore, higher abundances of these elements indicate that enrichment in Cr and Ni contents is probably due to the contribution of volcaniclastic input. The presence of 88 ppm Ni in cherts and 300 ppm in shales, 976 ppm Cr in shales and 300 ppm in CBIF probably cannot be explained satisfactorily unless the fumerolic and explosive volcanic activity producing ash is assumed to occur in this basin. Ni abundances vary between 2 and 250 ppm and Zr varies from 0.1 ppm to 716 ppm in cherts to SBIF through CBIF. A gradual and simultaneous increase ( r = 0.28 ) of these two elements is seen from CBIF and cherts to SBIF and shales (Fig. 3C). Zr abundances of SBIF and SH are distinctly higher than those of CH, FH and CBIF and thus the N i / Z r ratios and behaviour demonstrate the clastic contribution to these rocks. Maximum concentration of Zr found in SBIF is 280 ppm and that of SH is 700 ppm. In CBIF it varies between 0.1 and 20 ppm. Evidently Zr is contributed from terrigenous sediments. In the Isua BIF also Zr values are highest in the aluminous variety (Dymek and Klein, 1988). The Zr and Cr relationship (r=0.52) brings out this aspect more clearly (Fig. 3D). In general both Zr and Cr enrichment has taken place simultaneously in CBIF and SBIF. SBIFs relative to CBIF are enriched in both Cr and Zr suggesting that hydrothermal, volcaniclastic and terrigenous sources have contributed to the observed composition of BIF. This is further substantiated from the Zr and Cr data of shales, where shales enriched in MgO have higher Cr and shales relatively depleted in MgO and enriched in A1203 are more enriched in Zr (Fig. 3D). Similarly very low concentrations ofZr ( < 3 ppm ) in the majority of C, FC, and CBIF samples (Fig. 3D) are suggestive that contribution of solutions derived from chemical leaching of the land outside the basin to the Fe and Si of the
BIF OF THE ARCHAEAN SANDUR GREENSTONE
149
BELT
BIF may not be significant but it cannot be ruled out as Zr would have been left in the leached rock suites and this will come to the basin only as clastic component. Wherever terrigenous contribution has increased the Zr value has increased too. This is further strengthened by the Zr and Y behaviour (Fig. 3E) and Zr and YREE patterns (Fig. 3F). A gradual increase of both Zr and Y (r=0.73) is observed (Fig. 3E) and also no compositional gap between the two is seen. Similarly a sympathetic positive relationship (r=0.66) between Zr and YREE (Fig. 3F) with no gap in between suggests that mixing of chemical precipitates and clastics has been the main reason for the observed compositional variations. Zr and Hf both increase simultaneously ( r = 0.77) in our samples and show an excellent linear relationship (Fig. 4A) between their values from CH, FC, CBIF, SBIF, and SH. There is no gap between cherts to shales through CBIF or SBIF. This suggests that a continuous supply of a varying quantity of the divergent types of terrigenous and volcaniclastic debris has been maintained to the basin and the mixing of hydrothermal solutions with clastics derived from two divergent sources has taken place. Zr, Y and Hf are provided by felsic rocks and if felsic volcanic ash would have provided these elements, appreciably high concentrations of K20 and NazO of the order of their abundances in felsic volcanics should be preserved in SBIFs and SH, as the settling of volcanic ash being a rapid processes without weathering and fluvial abrasive transportation will not allow removal of alkalies from feldspars. Contrary to this, only in four samples of SBIF and one sample of SH relatively higher concentrations of alkalies are found (Table 1 ). Abnormally high AI203 contents from 8 to 33% in SBIF and shales and low alkali contents in the corresponding samples, suggests that probably a considerable part of the clay of the BIFs has been derived from a deeply weathered continental source and not from within the basin felsic volcanic source.
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150
XRD data have shown that most of it is kaolin and hence indicates to a feldspar-rich source. Rb and Sr in all geochemical process generally accompany alkalies. In general CBIF and SBIF appear as separate groups on the Rb-Sr plot (Fig. 4B ). Most of the CBIF contain Rb below one ppm. A few samples have Rb more than 13 ppm. Whereas the Rb content of SBIFs generally remains below 100 ppm but goes up to 257 ppm in two samples. Similarly the maximum Sr content of CBIF is found to be only 19 ppm and in SBIF it goes up to 66 ppm. In general CBIF have lower abundances of Sr, whereas SBIF and shales have high Sr content which reaches up to 638 ppm in sample No. SM 17. The large and random scatter of Rb and Sr (r=0.12) in both varieties of BIF, shows a multicomponent contribution for the constituents of BIF (Fig. 4B). Vanadium and scandium show sympathetic linearity (r=0.87) and a noticeable distinction between CBIF and SBIF. Maximum enrichment of V and Sc has occurred in SBIF and a few samples of chert and ferruginous chert and CBIF are extremely depleted in both V and Sc. A complete gradation of CBIF into SHBIF is seen on the V-Sc plot (Fig. 4C ). Similar distinctive behaviour is seen in the Ta and Nb distribution (Table 1 ). Cu and Zn concentrations and sulfide veining in source samples indicate a late stage hydrothermal syngenetic or postgenetic emplacement of these elements.
REE geochemistry REE are estimated in 102 samples but sixty seven chondrite normalised patterns are presented in Figs. 5 and 6. Based on REE pattern shape and La, Ce and Eu anomalies, these samples are divided into the following twelve groups. Since both La and Ce enrichment and/ or depletion has taken place, N d N / Y b N ratios are used, instead of La and Lu ratios to get an idea about REE fractionation. Groups 1 to 8 are of the samples belonging to chert, FC, and
C. MANIKYAMBAET AL.
CBIF. Groups 9 and 10 are of SBIF and Groups 11 and 12 represent shales. (1) Group 1 (Fig. 5A) consists of eight samples with extremely depleted EREE, loworder La enrichment, positive Eu anomalies, and NdN/YbN= 1 and LREE/HREE > l, LREE from La to Sm show fractionation, HREE are flat. (2) Group 2 (Fig. 5B) consists of 4 samples with extremely depleted YREE, low-order La enrichments, negative Ce anomalies, negative Eu anomalies. NdN/YbN 1. Sample No. SM24 is a pure chert exhibiting negative Ce and Eu anomalies. (3) Group 3 consists of four samples (Fig. 5C) with extremely depleted ~REE, depletion in La, positive Ce and Eu anomalies, NdN/YbN ~<1, and LREE/HREE slightly < 1. (4) Group 4 consists of seven samples (Fig. 5D) in which all REEs, except La and Eu, are extremely depleted. La is highly enriched, Ce is depleted and shows strong negative anomalies. Eu exhibits weak to strong positive anomalies. NdN/YbN>~ 1, including La, HREE/ L R E E < I , excluding La, L R E E / H R E E < I . One sample, No. CM40, shows negative Eu anomaly. From Nd to Lu the pattern shapes are fiat. NdN/YbN and HREE/LREE, excluding La, are close to one. This group of samples is very important because of its high-order La enrichment. (5) Group 5 consists of six samples (Fig. 5E) having moderate to very high ZREE abundances. This group is characterized by very high order La enrichment and negative Ce anomalies, positive Eu anomalies. NdN/ Y b N > I , excluding La and Ce, LREE/ HREE > 1. Sample No. MM24 shows very high N d N / Y b N and LREE/HREE ratios. Sample No. CM36 shows a very strong negative Ce anomaly. In fact negative Ce anomalies are the most characteristic and important feature of this group. (6) Group six (Fig. 5F) consists of six samples which are characterized by moderate ~REE abundances with positive Eu anomaly,
BIF OF THE
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151
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Fig. 5. (A) Extremely depleted REE pattern in ferruginous cherts and cherty BIF showing slight La enrichment and strong positive Eu anomalies. (B) Extremely depleted REE patterns in cherts and cherty BIF showing negative Ce and Eu anomaly. (C) REE patterns with slight La depletion strong Eu enrichment and weak positive Ce anomalies. (D) Extremely depleted REE patterns showing high order La enrichment, slight negative Ce anomalies, strong positive Eu anomalies, with almost flat pattern shape between Nd to Yb. (E) REE patterns moderately enriched in •REE with very strong negative Ce anomalies, showing fractionated pattern from Pr to Lu and positive Eu anomalies. Four samples show Gd depletion. Sample No. MM24 shows extremely fractionated pattern. Extrapolation from Pr to La in each case shows enrichment in La. (F) REE patterns of cherty BIF with moderate ~ REE abundances, showing fractionation from La to Lu with positive Eu anomalies. No La enrichment or Ce depletion is recorded in these samples. For legend see also Fig. 3.
NdN/YbN> 1 and LREE/HREE> 1. Sloping REE patterns due to slight depletion in HREE and positive Eu anomaly are characteristic features of this group. No La enrichment and Ce depletion are recorded in this group. (7) Group seven (Fig. 6A) consists of six samples which are characterized by moderate ZREE abundances, fractionated pattern from
La to Sm, positive Eu anomalies, flat HREE except in one sample (SM 19), NdN/YbN> 1 and LREE/HREE > 1 fractionation from La to Sm is the most characteristic feature of this group. (8) Group eight (Fig. 6B) consists of five samples with moderate •REE abundances, moderate negative Ce anomalies and moder-
152
C. MANIKYAMBAET AL. I000
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Fig. 6. (A) REE patterns with moderately enriched YREE, slight La enrichment, positive Eu anomalies, sloping LREE and fiat HREE. (B) Moderate total REE enriched patterns with weak negative Ce anomalies, positive Eu anomalies, three samples exhibit flat LREE from Pr to Sm. HREE are sloping. (C) Highly enriched REE patterns in SBIF with very weak negative Eu anomaly, with almost average fiat patterns. (D) Extremely enriched Y'REE patterns in SBIF showing no or very weak positive or negative Eu anomaly. (E) Enriched REE patterns in shales (phyllites) with negative and positive Eu anomalies and moderate to strong positive Ce anomalies. (F) Moderately enriched REE patterns of shales resembling NASC. For legend see also Fig. 3.
ate positive Eu anomalies. Three samples show fiat pattern shape from Ce to Sin, whereas two other samples show fractionated pattern shape from Pr to Lu, NdN/YbN and LREE/ HREE> 1. A general fractionation pattern shape is the significant characteristic of this group. (9) Group nine consists of seven samples of SBIF (Fig. 6C) this group is characterized by enrichment in ~REE compared to CBIF, no La or Ce anomalies, weak positive or no Eu
anomalies and NdN/YbN and LREE/ HREE> 1. (10) Group ten (Fig. 6D) consists of eight samples of SBIF having enrichment in Y.REE with fiat to slightly sloping pattern shapes and weak positive and negative Eu anomalies. Enrichment in total REE is the main distinguishing feature of this group. NdN/YbN and HREE/LREE< 1. In this group 2 samples which are not plotted in this illustration also exhibit positive cerium anomalies.
BIF OF THE ARCHAEAN SANDUR GREENSTONE BELT
( 11 ) Group eleven (Fig. 6E) consists of 3 samples of shales showing enriched ZREE with strong positive Ce and weak negative Eu anomalies. N d y / Y b y and H R E E / L R E E < 1. The appearance of positive cerium anomalies is one of the most outstanding aspects of this group. (12) Group twelve (Fig. 6F) consists of three samples of shales/phyllites with enrichment in EREE relative to CBIF. They show both very weak positive and negative Eu anomalies. NdN/YbN and LREE/HREE are slightly greater than one. These patterns of shales resemble those of NASC.
Enrichment of XREE: consequence of clastic contribution to BIF The absolute REE content of the Archaean Sandur schist belt BIF shows great variation in abundance, from very low to very high (Table 1 ). Such patterns (Figs. 5 and 6) represent the end product of a complex series of events that record the properties of the solutions that precipitated them with clastic sediments. A contemporaneous input of clastic material appears to have greatly effected the absolute abundance of REE in these rocks. The overall shape of the REE patterns of the oxide facies BIF samples comprising only chert and iron minerals is generally similar to patterns of modern seawater, with the exception of La, Ce and Eu which show anomalous behaviour. Derry and Jacobsen (1990) have also found a similarity in the overall shape of REE patterns of Archaean oxide facies BIF and modern seawater, with the exception of Ce and Eu. Despite a large scatter in the LREE/HREE ratio (Fig. 7A) ( < 1 to > 4 0 ) , LREE and HREE increase simultaneously when the entire population ofCH, FC, CBIF, SBIF and SH is considered together. The N d / Y b ratio shows the same scatter and a simultaneous increase of Nd and Yb from pure CBIF to SBIF. The lowest Nd and Yb values are found in chert and pure CBIF (irrespective of the mutual varia-
15 3
tion of SiO2 and Fe203 content, both these oxide comprise almost 100% of the sample) whereas the highest values of Nd and Yb occur in shales and SBIF (Fig 7B). A complete gradation between these values is seen in the LREE/HREE and N d / Y b plots. A change in ZREE from < 1.00 ppm to > 421 ppm from CBIF to SBIF is apparent in Figs. 5 and 6. These data clearly demonstrate that the varied amount of clastic input into the chemogenic sediments has increased the YREE, and the variation in LREE/HREE and Nd/Yb. Fractionation from La to Lu, a characteristic of felsic rocks is not observed and the REE patterns and shapes with La, Ce and Eu anomalies appear to be coming from from three different sources. The modern oceans receive most of their REEs from three sources: (1) the dissolved load of rivers (Goldstein and Jacobsen, 1988a, b), (2) hydrothermal alteration of the oceanic crust (Michard and Albarede, 1986), and ( 3 ) solutions derived from sediments undergoing diagenesis. The magnitude and pattern of the diagenetic REE contribution is uncertain, but on global scale it is probably small relative to erosional and hydrothermal fluxes (Sholkovitz et al., 1989). REE data from metalliferous sediments from the East Pacific Rise (EPR) show evidence of scavenging of REE by Fe oxyhydroxides and a clear pattern of gradual decrease of LREE depletion with distance from the ridge axis (Ruhlin and Owen, 1986). According to Derry and Jacobsen (1990) these relations imply that metal oxy-hydroxides precipitating from the hydrothermal plumes are preferentially enriched in LREEs such that by the time the REEs in the plume are far removed from the ridge axis, the residual water mass is LREE depleted. The similarities between the REE patterns of modern metalliferous sediments and the oxide facies BIFs suggest that scavenging by Fe hydroxides was the mechanism responsible for the incorporation of the REEs into BIF. Derry and Jacobsen ( 1990 ) further conclude that the overall trend
154
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Fig. 7. (A) Mutual relationship between LREE and HREE indicating simultaneous increase of LREE and HREE from ferruginous cherts to SBIF. (B) Nd-Yb relationship showing the scatter in the ratio from 1 to 40 and simultaneous increase of Nd and Yb from CBIF to SBIF. (C) Eu-Eu* relatioship indicating the change in the nature and magnitude of Eu anomalies from CBIF to SBIF and SH. (D) (ZREE-La) are plotted against La abundances to show the highly La enriched characteristics of some CBIF samples, (E) La-Fe203 relationship shows that between 40 and 60% of Fe203 a large number of samples of CBIFs have high order La content similar to those of SBIF in which anomalous behaviour of La is not recorded. (F) Ce-Fe203 relationship shows that at 40-60% level of Fe203, maximum Ce depletion has taken place. For legend see also Fig. 3. o f REEs in BIFs are compatible with a mechanism o f oxide facies BIF formation in which Fe hydroxides precipitated from seawater with a REE pattern similar to the modern ocean (except for Ce and Eu ). The variability o f REE patterns in the IF samples probably results from differences in scavenging efficiency, with the most depleted patterns resulting from the most efficient REE removal. Our data, however, indicate that the variability o f REE patterns in BIF samples results from the mixing o f the clastic c o m p o n e n t s with that o f the sea-
water in which hydrothermal solutions were emplaced at an Archaean mid-oceanic ridge (AMOR).
Eu, La and Ce anomalies." possible signatures of AMOR hydrothermal activity and biogenic oxidation Eu anomalies Except 5 samples o f C, CBIF and 7 samples o f SBIF all samples o f BIF, chert, and ferruginous chert from the Sandur belt exhibit posi-
BIF OF THE ARCHAEAN SANDUR GREENSTONE BELT
tive Eu anomalies with variable magnitude (Fig. 7C). Such positive Eu anomalies are observed in BIFs irrespective of their age or locality and can only result from the input of Eu enriched hydrothermal fluids into the water column of the oceans (Graf, 1978; Fryer et al., 1979; Barret et al., 1988). Kerrich and Fryer (1979) and Fryer et al. (1979) suggested that the Eu enriched nature of Archaean and early Proterozoic seawater was due to a major contribution of strongly reducing hydrothermal fluid discharging on to the sea floor during these times. Experimental work of Flynn and Burnham ( 1978 ) supports this view. Modern seawater does not typically display Eu anomaly except near hydrothermal vents. REEs in present day ocean water occur in exceedingly low concentrations with patterns that show depletion in Ce and Eu and LREE < HREE. These characteristics are interpreted to indicate a control by continental input on the abundance of Eu, while the relative depletion of the LREE, especially Ce, is due to selective scavenging of seawater REE during the formation of Mn nodules. The source of anomalous Eu concentration in the BIFs requires that the hydrothermal REE flux comprised a much larger fraction of the total REE flux than today (Derry and Jacobsen, 1990). Present day hydrothermally active regions in EPR and M P R (MidPacific Rise) are characterized by low EREE abundances, extreme positive Eu anomalies and LREE enrichment, because the REEs of the fluids are controlled by high temperature rockwater interactions in the volcanic sequence at depths (Michard et al., 1983; Campbell et al., 1988). Such conditions favour the stability of Eu 2+ in aqueous solution and the development of positive Eu anomalies (Sverjensky, 1984). The signature of hydrothermal REE appears to die out rapidly by mixing with seawater and the fact that Pacific sea water near EPR is relatively enriched in Eu compared to North Atlantic water may represent the effect of a hydrothermal component (Klinkhammer et al., 1983). Dymek and Klein (1988) have
15 5
computed that a 1:100 mix of North Atlantic seawater and EPR hydrothermal fluid is able to retain positive Eu anomaly. This 1:100 mix of EPR and NASW (North Atlantic Sea Water) yielded patterns similar to uncontaminated BIFs of Isua Formation. The deep (2500 m) hydrothermal plume waters directly above the EPR at 19°S (Klinkhamer et al., 1983) have an REE pattern completely different from that of the vent waters, yet still distinct from normal seawater (De Baar et al., 1985). Although vent driven water in this plume is strongly diluted by ambient sea waters, a hydrothermal component is recognized on the basis of tracers such as Mn and He 3, (Weiss, 1977; Lupton and Craig, 1981 ) for distance of hundreds meters above and hundreds of kilometers west of EPR. By contrast, the hydrothermal Eu anomaly is lost immediately following the discharge ofhydrothermal solutions into seawater. The Eu anomaly is also absent in metalliferous sediments from various ridge related and DSDP sites in the eastern Pacific (Ruhlin and Owen, 1986; Barrett and Jarvis, 1988). Clearly this anomaly is not transmitted in modern oxidized open oceans, not even to plumes only a few hundred meters above the ridge axis. This is apparent because of rapid oxidation of hydrothermal Eu 2÷ to E u 3 + in the water column; any precipitates formed also absorb sea water REE that lack a positive Eu anomaly. However, ifArchaean oceans were relatively reducing, transmittance of positive Eu anomaly away from its hydrothermal source seems plausible (Barrett et al., 1988). The total range of the REE content of the BIFs of the Sandur schist belt is < 1.00 p p m to >421 p p m (Table 1) and as the )~REE increase from CBIF to SBIF, the amplitude of the positive Eu anomaly gradually diminishes and even becomes slightly negative (Fig. 7C ). This shows that the REE patterns and Eu anomalies of BIFs of the Sandur schist belt are greatly affected as a result of mixing, not only with the seawater but also with the clastic component.
156
La enrichment and Ce depletion Apart from the positive Eu anomalies, which are characteristic of BIF all over the world (Fryer, 1983; Derry and Jacobsen, 1990), the most important and significant aspect of the REE data from the Sandur belt BIF are La enrichment and negative Ce anomalies. Thirtysix samples of CBIF, chert and ferrugenous chert exhibit La enrichment of low to very high order. Since La is behaving abnormally, its concentration is deducted from YREE and thus a plot of ( ~ R E E - L a ) against La is made to illustrate the behaviour of La in the total population of samples analysed. Among the samples of SBIF where YREE is very high, La and ( E R E E - La) show a well defined sympathetic linear relationship whereas most of the CBIF, chert and FC samples fall below that line of linearity showing La enrichment (Fig. 7D). In CBIF a maximum variation in the La content (from 0.025-47 ppm) is found at 40-60% Fe203 content (Fig. 7E). Ce is also anomalous and at 40-60% Fe203 many samples show strong Ce depletion. The maximum depletion in the Ce (0.08 ppm) content is recorded in sample No. CM36 having about 85% Fe203 (Fig. 7F). Ce depletion is excellently projected in Fig. 8A in which ~ R E E - ( L a + C e ) is plotted against Ce. Between 0.3 to 1.0 ppm Ce depletion level the 3~REE- ( L a + C e ) varies from 0.56 to 7.0 ppm. SBIFs depict a very clear cut sympathetic linear relationship between Ce and ~ R E E - ( L a + C e ) (Fig. 8A). Nd is not showing any abnormal behaviour in our samples. Therefore, ratios between Nd, Ce and La in a ternary plot illustrate the La enrichment and Ce depletion (Fig. 8B). The other locality from where La enrichment consequent of hydrothermal spiking has been recorded is the Archaean Greenstone belt of Canada. Barrett et al. (1988) have interpreted that this La enrichment is caused by the La spiking from hydrothermal solutions. In support of their inferences, Barrett et al. have cited an example of Red Sea hydrothermal sediments where all chemical precipitates show
C. MANIKYAMBA ET AL.
distinct La enrichment. They postulate that IF with La enrichment and strong positive Eu anomalies have been formed from reduced bottom waters carrying a significant hydrothermal signal. Accordingly, IF with no La enrichment and smaller Eu anomalies may have been precipitated from solutions that were diluted strongly by ambient ocean water during periods of lessened discharge of hydrothermal solutions or from solutions far removed from vent sources. Barrett et al. (1988) proposed that the ambient seawater, although probably characterized by positive Eu anomalies, had an otherwise near fiat pattern with no La enrichment. Archaean seawater in general is expected to be more reducing, which should be favourable for the retention of hydrothermal signatures, but the precipitation of iron oxides in jaspery chert and haematite/magnetite indicate that away from the locus of hydrothermal input an oxygenated environment existed and that the FeO + SiO2 were transported from the reduced environment and were precipitated in the oxygenated environment. During this transport, mixing with the ambient seawater and clastic sediments took place resulting in the obliteration of La spiking in many samples. Whenever mixing of ocean water and clastic sediments was incomplete and ineffective La and Eu enrichment is pronounced. On the other hand, samples with less pronounced or nil La and/or Eu enrichment indicate that ocean water and clastic mixing with the hydrothermal solutions has obliterated and masked the hydrothermal signature. In some samples of chert and cherty BIF (Fig. 5B)extremely depleted ~REE abundances are associated with negative Ce and negative Eu anomalies. These samples could have been deposited from oxygenated ocean waters near the shoreline having their Fe and REEs brought by rivers as dissolved load from continental source to a zone where oxygen was produced more rapidly by stromatolites developing photosynthetic bacteria. Seven samples of FC and CBIF also show La depletion and Ce enrichment. These data
BIF OF THE ARCHAEAN SANDUR GREENSTONE BELT
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Fig. 8. (A) Relationship between ~REE- (Ce+La) and Ce. The linear relationship is shown by the straight line from SBIF to ferruginous chert, most of the CBIF samples are falling above the line indicating Ce depletion in CBIFs and ferruginous cherts suggesting oxidation of the hydrothermal solutions. (B) shows La:Ce:Nd ratio distribution in the total population of the samples analysed. Since La and Ce both are showing abnormal behaviour with respect to Nd this ratio is extremely useful for illustrating La enrichment and Ce depletion with respect to Nd. (C) Relationship between EREE and AI203 illustrates that variation and scatter in REE is consequent of mixing with the divergent clastic source. (D) Relationship between EREE and sum of Ni, Cu and Co. Fields of hydrothermal deposits and hydrogenous sediments are after Klein and Beukes (1989). For legend see also Fig. 3. indicate that precipitation o f CBIF has taken place in a weak and intermittently oxygenated environment. The inference that both types namely reducing and oxidizing environments prevailed in Archaean oceans is further substantiated by the size o f the negative Eu anomalies found in some cherts and cherty BIF samples. Based on limited data available at that time Fryer ( 1983 ) inferred that significantly anomalous Ce abundances are not known from Archaean IF, whereas in Proterozoic BIFs Ce is definitely anomalous with examples o f both enrichment and depletion being observed. Thus, Fryer (1983) suggested that Ce was being separated from the rest of the REE by
oxidation to the Ce 4+ state, at least back to the early Proterozoic, which would require that areas of moderately to strongly oxidizing conditions must have existed in the Proterozoic seas. REE patterns of IF of the early Proterozoic Transvaal Supergroup are characterized by strong depletion in LREE relative to H R E E , strong negative Ce anomalies and marked positive Eu anomalies (Beukes and Klein, 1990). Also, Ce depletion in the modern seawaters is attributed to the removal of Ce to Mn nodules by oxidation of Ce 3+ to Ce 4+ (Addy, 1979) for which a typical marine p H of ~ 8 and Eh values o f 0.15-0.45 V are required. The early Proterozoic IFs, which exhibit anomalous Ce, are also enriched in Eu compared to their clas-
158
tic sediments. Interpreting these data Fryer (1983) suggests that there was a significant hydrothermal input to the Proterozoic seas from which these deposits are formed. Our data (Figs. 5 and 6) profusely demonstrate the presence of modest to strong negative Ce anomalies in many samples of CBIF along with positive Eu anomalies and La enrichment. Two SBIF samples from the same locality exhibit positive Ce anomalies similar to some samples of shales interbedded with CBIF and SBIF, (Fig. 6E). A few samples of CBIF (Fig. 5 ) also show weak to moderate positive Ce anomalies. Discussion
The relationship between ZREE and A1203 is very convincing in demonstrating that increases in ZREE are a direct consequence of simultaneous clastic deposition and chemogenic precipitation (Fig. 8C). A mixing of clastic and chemogenic components is illustrated in Figs. 3 and 4. The clastic component was probably contributed in in a volcaniclastic and a terrigenous form. This aspect is demonstrated by the Ni, Cr, Zr, MgO, Ti, V, Hf and other elements (Figs. 3 and 4). High Cr and Zr contents for example in highly aluminous SBIF can be contributed to the two divergent clastic sources. Since the end product is an outcome of three divergent processes absolute linearity between any two distinguishing parameters will not develop, because the third will interfere with the other two. Positive Eu anomalies, La enrichment and depleted patterns of REE indicate that hydrothermal solutions brought these REEs and supplied FeO and SiO2 from deeper sources (see Morey, 1983). The inference that the REE, FeO and SiO2 of the BIF was mainly supplied by hydrothermal solutions emplaced at AMOR is strongly substantiated by Fig. 8D. Here, almost all data points of CBIF fall within the field of hydrothermal deposits, whereas those of SBIF and SH fall outside and away from this field. These hydrothermal solutions were em-
C. MANIKYAMBA ET AL.
placed into the ambient seawater which was largely reducing at that time. There is convincing evidence of production of photosynthetic oxygen through stromatolites in the shallower part of the Sandur basin. In other regions the evidence of stromatolites push the record of life and photosynthesis back as far as 3.4 or 3.5 Ga (Lowe, 1980; Walker et al., 1983). Biogenically produced oxygen reacted with solutions containing REE, FeO and SiO2 and precipitated iron as Fe hydroxides and oxidized Ce to its tetravalent state resulting in negative Ce anomalies as in modern oceans. The Ce thus removed from BIF was accummulated in some shales (Fig. 6E ). This reaction appears to have taken place at the shallow shelf region at the wave base interaction level not very much below the photic zone, where the precipitation of ferrous iron to ferric iron was mediated by biogenic oxidation; simultaneous volcanic activity and terrigenous sedimentation resulted in a ficterogenous product. Thus the observed compositional characters in EREE, La enrichment, Ce depletion, positive Eu anomalies, variation in A1203, TiO2, MgO, Rb, Sr, Cr, Ni, Zr, Hf, Sc, V, Y and other elements appear to be controlled by divergent clastic input to the ambient ocean, its mixing with the hydrothermal solutions and oxidation by photosynthetic oxygen. Combination of these four processes has diluted the effect of each one of them and a hybrid sequence of chert to shales through CBIF and SBIF is developed in a shelf sequence. The hydrothermal activity appears to have taken place in a relatively deeper portion of the basin and was accompanied by intermittent volcanism. Hydrothermal solutions were enriched in dissolved ferrous iron. Due to thermal and chemical gradient and upwelling, these solutions were mixed with ocean water which was reducing except in places where photosynthetic bacteria were producing oxygen. As a result of ocean circulation and thermo-chemical gradients these solutions moved towards the shoreline (see Borahert, 1960) and met oxygen below the wave base or
BIFOFTHEARCHAEANSANDURGREENSTONEBELT
159
TABLE 2 Flow chart of the Sandur belt BIFs RAMORHydrothermal solutions ] Exposed continental nuclei educing andalkalinetoneutralenvironment (depleted S~REE,I IAnaerobic environmental, intense chemical and physical a, Eu enrichment, Higher concentration of SiO2 and FeO) J ~veathering I
dTransport towards shoreline due to thermal and chemical gra~ Dissolved load in river waters suspended load of terrigenous / material (negative europium anomalies, high YREE, A1203, Zr ient and ocean circulation 'and Hf ) Volcanism within the basin, I volcaniclastic input (high Cr and Ni ! e a r stable shelf Biogenically mediated oxida- I ntermittently oxidizing and alternating acidic tion. (negative Ce anomalies,[ nd alkaline environment I presence of stromatolites mi-] [crobiota organic carbon) ] I Si02 precipitation and rhythmic intermittent Fe/03 ontinuous below wave base and photic zone. Fluctuations in [precipitation Iwave base
~
at a fluctuating wave base level, got oxidized and precipitated as Fe203, deposited below the wave base and the photic zone. However, fluctuations in the wave base and probably the photic zone also produced some chert bands in which microfossils could be preserved (Naqvi et al., 1987; Venkatachala et al., 1990). The precipitation of silica in such an enormous quantity, in the absence of any known Archaean biogenic activity, capable of silica deposition remains an enigma, although LaBerge (1986) believes that some form of silica precipitating microorganism existed in that period. Some workers believe that changes in Eh and pH values have caused silica precipitationu (Morey, 1983), whereas others suggest that evaporitic conditions and seasonal changes have caused precipitation of iron and silica (Trendall, 1983; Garrels 1987 ). The clay composition of the shaly beds in the Sandur, Chitradurga and Shimoga belts is mostly kaolinitic, indicating its derivation from a highly weathered source. Iron present i n the continental rocks which gave rise to kaolin must have been brought in solution to the basin and precipitated there with the iron from hydrothermal solutions. Some of the Fe would also
have been provided by leaching of the contemporaneous basic volcanic rocks. This may also explain the variation in the Eu anomalies from strongly positive to even negative in both CBIF and SBIF. We propose that the Sandur belt BIFs have been generated by a combination of factors represented in a flow chart given in Table 2. In this model, sedimentation of BIF occurred on the shallow stable shelf while intermittent volcanism and hydrothermal activity was in progress in the eastern part of the basin probably on an Archaean mid-oceanic ridge.
Conclusions From the above discussion and presentation of geochemical data on the oxide facies BIFs the following conclusions can be drawn: ( l ) Cherty and shaly BIFs are interbedded in the Sandur schist belt. Cherts and shales are two end members in which mixing of minerals containing A1203 and Fe203 in various proportions have given rise the observed compositional variations from cherts to ferruginous cherts, cherty BIF, shaly BIF to shales. These variations caused by mixing of volcaniclastic and terrigenous sediments with chemogenic
160 precipitates is illustrated by Ni, Cr, Zr, Y, V, Sc, Ti, FREE and NdN/YbN ratios. (2) The main source for the iron and silica were hydrothermal solutions generated at AMOR. The hydrothermal sources of iron and REE are reflected by La and Eu enrichment and extremely depleted EREE in CBIF. These patterns are similar to the REE patterns of EPR and Red Sea hydrothermal solutions. Fluvial contribution of FeO and REEs in the form of dissolved load from the land is not completely ruled out. (3) Signatures of the volcaniclastic debris are preserved in the high Cr and Ni content of both the CBIF and SBIF, and terrigenous sedimentation is reflected in Zr, Y, V, Sc, Hf, Ta and Nb abundances. (4) A large number of samples show negative Ce anomalies and indicate oxidation of hydrothermal solutions on stable shelf zones. The magnitude of La enrichment is variable suggesting the mixing of hydrothermal solutions with ambient seawater. ( 5 ) Strong positive Eu anomalies are found in CBIF and weak positive to negative Eu anomalies are found in SBIF indicating the diminishing nature of the hydrothermal signature from CBIF to SBIF. (6) Based on the geochemical data and field evidence it is suggested that the Sandur schist belt BIFs were deposited by a combination of four processes. Hydrothermal solutions and fluvial solutions provided the SiO2 and FeO and REE of the deposits. Hydrothermal solutions were emplaced in a reducing and neutral to alkaline environment near vents at AMOR. These solutions due to the thermal and chemical gradient convected towards the shoreline and got mixed with ambient ocean water and clastic input of volcanic and terrigenous origin. Existence of stromatolites in the underlying carbonates indicates photosynthetic activity which provided oxygen for the oxidation of ferrous iron to ferric iron and C e 3+ to C e 4+ and thus biogenically mediated precipitation of iron took place along with terrigenous and vol-
c. MANIKYAMBAETAL. caniclastic varved type beds. Near the shoreline the environment was acidic and alkaline with alternating precipitation of SiO2 and Fe203 in response to changes in Eh and pH. The banding of the BIF represents the break in precipitation of iron due to the nonavailability of photosynthetic 02 or hydrothermal FeO or both. Therefore, four processes, namely (1) hydrothermal activity in a deeper reducing alkaline environment and its movements towards the shoreline, (2) terrigenous sedimentation, (3) volcanic sedimentation, and (4) oxidation from biogenically produced oxygen, have produced the Sandur schist belt BIFs.
Acknowledgements We express our profound gratitude to Prof. D. Gupta Sarma FNA, Director NGRI, for his encouragement, keen involvement, support and permission for publication of this work. The work has been supported under the DST Grant No. SP 12/PC2/86 and NGRI project on Lithos 5.1. Mr. T. Gnaneshwar Rao, Mr. S.L. Ramesh and Mr. K.V. Anjaiah have provided help in the analysis of the samples. Mrs. Nancy Rajan has prepared the manuscript. We are deeply indebted to Dr. B.P. Radhakrishna, Dr. M. Ramakrishnan, Dr. R. Srinivasan, Dr. V. Divakara Rao, Dr. B.L. Narayana for a critical review of the manuscript and their suggestions for the improvement. We express our gratitude to Dr. G.L. Reddy, Mr. K.K. Reddy, M. M.H. Shareef, Mr. Masood Rizvi of NMDC, Donimalai iron ore mines who have greatly helped in sampling and provided core samples from the core library of NMDC. Dr. A.K. Chaterjee has extended all help and facilities at the Donimalai mine area. An earlier version of the manuscript has been greatly improved by extremely valuable review, comments and suggestions from Prof. H.L. James and A.M. Goodwin. We are extremely grateful to both of them.
BIF OF THE ARCHAEAN SANDUR GREENSTONE BELT
References Addy, S.K., 1979. Rare earth element patterns in manganese nodules and micronodules from northwest Atlantic. Geochim. Cosmochim. Acta, 43:1105-1115. Anhaeusser, C.R., 1990. Precambrian crustal evolution and metallogeny of South Africa. In: S.M. Naqvi (Editor), Precambrian Continental Crust and its Economic Resources. Elsevier, Amsterdam, pp. 123-156. Balaram, V., Manikyamba, C., Ramesh, S.L. and Anjaiah, K.V., 1991. REE and trace elements analysis of iron formation refrance samples by inductively coupled plasma mass spectrometry. Proc. 5th Natl. Symp. on Mass Spectrometry, Ahmedabad EPS, pp. 1-3. Barrett, T.J. and Jarvis, I., 1988. Rare-earth elements geochemistry of metalliferous sediments from DSDP Reg 92: the East Pacific Rise transect. Chem. Geol., 67: 243-259. Barrett, T.J., Fralick, P.W. and Jarvis. I., 1988. Rare-earth element geochemistry of some Archaean iron formalions north of Lake Superior, Ontario. Can. J. Earth Sci., 25: 570-580. Beukes, N.J., 1980. Suggestion towards and classification of and nomenclature for iron formation. Trans. Geol. Soc. S. Aft., 83: 285-290. Beukes, N.J., 1983. Paleoenvironmental setting of iron formations in the depositional basin of the Transvaal Supergroup, South Africa. In: A.F. Trendall and R.C. Morris (Editors), Iron Formations: Facts and Problems. Elsevier, Amsterdam, pp. 131-209. Beukes, N.J. and Klein, C., 1990. Geochemistry and sedimentology of a facies transition from microbanded to granular iron formation in the Early Proterozoic Transvaal Supergroup, South Africa. Precambrian Res., 47: 99-139. Beukes, N.J., Klein, C., Kaufman, A.J. and Hayes, J.M., 1990. Carbonate petrography, kerogen distribution and carbon and oxygen isotope variations in an early Proterozoic Transition from limestone to iron formation deposition, Transvaal Supergroup, South Africa. Econ. Geol., 85: 663-690. Bhaskar Rao, Y.J., Sivaraman, T.V., Pantulu, G.V.C., Gopalan, K. and Naqvi, S.M., 1992. Rb-Sr ages of Late Archaean metavolcanics and granites, Dharwar craton, evidence for Early Proterozoic thermotectonic event (s). Precambrian Res., submitted. Borahert, H., 1960. Genesis of marine sedimentary iron ore. Trans. Inst. Min. Metall., 69: 261-277. Campbell, A.C. and 10 co-authors, 1988. Chemistry of hot springs on the Mid-Atlantic Ridge. Nature, 335:514519. Chakraborty, K.L. and Majumdar, T., 1986. Geological aspects of the banded iron formation of Bihar and Orissa. J. Geol. Soc. India, 28: 71-91. Cloud, P.E., 1973. Palaeobiological significance of the banded iron formation. Econ. Geol., 68:1135-1143.
161
Condie, K.C., 1981. Archaean Greenstone Belts. Amsterdam, Elsevier, 434 pp. Davy, R., 1983. A contribution on the chemical composition of Precambrian iron formations. In: A.F. Trendall and R.C. Morris (Editors), Iron Formations: Facts and Problems. Elsevier, Amsterdam, pp. 325-343. De Baar, HA.W, Bacon, M.P. and Brewer, P.G., 1985. Rare earth elements in the Pacific and Atlantic Oceans. Geochim. Cosmochim. Acta, 49: 1943-1959. Derry, L.A. and Jacobsen, S.B., 1990. The chemical evolution of Precambrian sea water: Evidence from REEs in banded iron formations. Geochim. Cosmochim. Acta, 54: 2965-2977. Devaraju, T.C. and Laajoki, K., 1986. Mineralogy and mineral chemistry of the manganese poor and manganiferous iron formations from the high grade metamorphic terrain of southern Karnataka, India. J. Geol. Soc. India, 28: 134-164. Drever, J.I., 1974. Geochemical model for the origin of Precambrian banded iron formations. Geol. Soc. Am. Bull., 85:1099-1106. Dimroth, E., 1986. Depositional environments and tectonic settings of cherty iron formations of the Canadian Shield. J. Geol. Soc. India, 28: 239-250. Dymek, R.F. and Klein, C., 1988. Chemistry, petrology and origin of banded iron formation lithologies from the 3800 Ma Isua supracrustal belt, West Greenland. Precambrian Res., 39: 247-302. Drury, S.A., Holt, R.N., Vanclastern, P.C. and Beckinsale, R.D., 1983. Sm-Nd and Rb-Sr ages of Archaean rocks in western Karnataka, South India. J. Geol. Soc. India, 24: 454-459. Eugster, H.P. and Chou, I.M., 1973. Depositional environments of Precambrian banded iron formations. Econ. Geol., 68:1144-1168. Ewers, W.E. and Morris, R.C., 1981. Studies on the Dales Gorge Member of the Brockman Iron Formation, Western Australia. Econ. Geol., 77:1929-1953. Flynn, R.T. and Burnham, C.W., 1978. An experimental determination of rare earth partitioning coefficients between a chloride containing vapour phase and silicate melts. Geochim. Cosmochim. Acta, 42: 685-701. Fryer, B.J., Fyfe, W.S. and Kerrich, R., 1979. Archaean volcanogenic oceans. Chem. Geol., 24: 25-33. Fryer, B.J., 1983. Rare earth elements in iron formations. In: A.F. Trendall and R.C. Morris (Editors), Iron Formation: Facts and Problems. Elsevier, Amsterdam, pp. 345-358. Garrels, R.M., 1987. A mode for the deposition of the microbanded Precambrian iron formations. Am. J. Sci., 287: 81-106. Gole, M.J. and Klein, C., 1981. Banded iron formation through much of time. J. Geol., 89:169-183. Goodwin, A.M., 1973. Archaean iron-formations and tectonic basins of Canadian Shield. Econ. Geol., 68:915933.
162 Goodwin, A.M., Thode, H.G., Chou, C.L. and Karkhansis, S.N., 1985. Chemostratigraphy and origin of the late Archaean siderite-pyrite rich Helene Iron Formation, Michipicoten belt, Canada. Can. J. Earth Sci., 22: 72-84. Goldstein, S.J. and Jacobsen, S.B., 1988a. Nd and Sr isotopic systematics of river water suspended material: implications for crustal evolution. Earth Planet. Sci. Lett., 87: 249-265. Goldstein, S.J. and Jacobsen, S.B., 1988b. REE in the great Whale river Estuary, NW Quebec. Earth Planet. Sci. Lett., 88: 241-252. Govil, P.K., 1985. X-ray fluorescence analysis of major, minor and selected trace elements in new IWG refrance rock samples. J. Geol. Soc. India, 26: 38-42. Graf, J.L., 1978. Rare earth elements, iron formations and sea water. Geochim. Cosmochim. Acta, 42:1845-1850. Gross, G.A., 1980. A classification of iron formation based on depositional environment. Can. Mineral., 18:215222. Hargraves, R.B., 1986. The metallogenetic significance of iron formation and related stratifer rocks. J. Geol. Soc. India, 28: 92-108. Hofmann, H.J., Saxe, R.P. and Herdusco, E.N., 1991. Archaean stromatolites in Michipictoen group siderite ore at Wawa, Ontario. Econ. Geol., 86: 1023-1030. Holland, D.H., 1973. The oceans: a possible source of iron in iron formations. Econ. Geol., 68:1169-1172. James, H.L., 1983. Distribution of banded iron-formation in space and time. In: A.F. Trendall and R.C. Morris (Editors), Iron Formation: Facts and Problems. Elsevier, Amsterdam, pp. 471-490. Kasting, J.F., 1987. Theoretical constraints on oxygen and carbondioxide concentrations in the Precambrian atmosphere. Precambrian Res., 34: 205-229. Kerrich, R. and Fryer, B.J., 1979. Archaean precious metal hydrothermal systems Dome mine, Abitibi greenstone belt: II REE and oxygen isotope relations. Can. J. Earth Sci., 16: 440-458. Khan, R.M.K., Govil, P.K. and Naqvi, S.M., 1992. Geochemistry and genesis of BIF from Kudremukh schist belt, Karnataka Nucleus, India. J. Geol. Soc. India, submitted. Klein, C. and Beukes, C., 1989. Geochemistry and sedimentology of facies transition from limestone to iron formation deposition in the early Proterozoic Transvaal Supergroup, South Africa. Econ. Geol., 84:17331773. Klinkhammer, G., Elderfield, H. and Hudson, A., 1983. Rare earth elements in seawater near hydrothermal vents. Nature, 305: 185-188. La Berge, G.L., 1986. A model for the biological precipitation of Precambrian iron formation. LPI Conference on Greenstone belts (Abstract) LPI Rep., Houston, TX, pp. 71-75. Litherland, M. and Malan, S.P., 1973. Manganiferous
c. MANIKYAMBAET AL. stromatolites from the Precambrian of Botswana. J. Geol. Soc. London, 129: 543-544. Lupton, J.E. and Craig, H., 1981. A major Helium-3 source at 15 °S on East Pacific Rise. Science, 214:13-18. Lowe, D.R., 1980. Stromatolites, 3400 Myr old from the Archaean of Western Australia. Nature, 284: 441-443, Majumdar, T., Chakraborty, K.L. and Bhattacharyya, 1982. Geochemistry of banded iron formation of Orissa, India. Miner. Deposit, 17:107-118. Manikyamba, C., 1992. Geochemistry of iron and manganese formations of Sandur schist belt, Dharwar craton, India. Ph.D. thesis, Osmania Univ., 327 pp. Michard, A. and Albarede, F., 1986. The REE content of some hydrothermal fluids. Chem. Geol., 55:51-60. Michard, A., Albarede, F., Michard, G., Minster, J.F. and Charlan, J.L., 1983. Rare earth elements and uranium in high temperature solutions from the East Pacific Rise hydrothermal vent fields. Nature, 314: 325-330. Morey, G.B., 1983. Animikie basin, Lake Superior region, USA. In: A.F. Trendall and R.C. Morris (Editors), Iron Formation: Facts and Problems. Elsevier, Amsterdam, pp. 13-68. Mukhopadhyay, D. and Matin, A., 1990. A new interpretation of the structural pattern in the Sandur schist belt in the Dharwar craton, Karnataka (Abstract). Seminar on Evolution of Precambrian crust in India, Calcutta, p. 33. Murthy, P.S.N. and Reddy, K.K., 1984. 2900 million year old stromatolite from Sandur greenstone belt of Karnataka Craton, India. J. Geol. Soc. India, 25: 263-266. Naha, K., Srinivasan, R. and Naqvi, S.M., 1986. Structural unity in the Early Precambrian Dharwar tectonic province, Peninsular India. Q. J. Geol. Min. Metal. Soc. India, 58: 219-243. Naqvi, S.M., 1981. The oldest supracrustal of Dharwar craton. J. Geol. Soc. India, 22: 458-469. Naqvi, S.M. and Rogers, J.J.W., 1983. Precambrian of South India. Mem. Geol. Soc. India, 4: 1-575. Naqvi, S.M. and Rogers, J.J.W., 1987. Precambrian Geology of India. Oxford Univ. Press, New York, 223 pp. Naqvi, S.M., Venkatachala, B.S., Manoj Shukla, Kumar, B., Natarajan, R. and Mukund Sharma, 1987. Silicifled cyanobacteria from the cherts ofArchaean Sandur schist belt, Karnataka, India. J. Geol. Soc. India, 29: 535-539. Naqvi, S.M., Sawkar, R.H., Subba Rao, D.V., Govil, P.K. and Gnaneshwar Rao, T., 1988. Geology and geochemistry and tectonic setting of Archaean greywackes from Karnataka nucleus, India. Precambrian Res., 39: 193-216. Naqvi, S.M., Manikyamba, C. and Gnaneshwar Rao, T., 1992. Geology of transition from stromatolitic carbonates to banded iron and manganese formations in the stable shelf zone of the Archaean Greenstone belts of Karnataka Nucleus, India. submitted. Nealson, K.H. and Myers, C.R., 1990. Iron reduction by
163
BIF OF THE ARCHAEAN SANDUR GREENSTONE BELT
bacteria: A potential role in the genesis of banded iron formations. Am. J. Sci., 290A: 35-45. Oehler, D.Z., Shopf, J.W. and Kvenvolden, K.A., 1972. Carbon isotope studies of organic matter in Precambrian rocks. Science, 175:1246-1248. Perry, E.C., Jr., 1983. Oxygen isotope geochemistry of iron formations. In: A.F. Trendall and R.C. Morris (Editors), Iron Formation: Facts and Problems, Elsevier, Amsterdam, pp. 359-371. Radhakrishna, B.P., 1983. Archaean granite greenstone terrain of South India. In: S.M. Naqvi and J.J.W. Rogers (Editors), Precambrian of South India. Mem. Geol. Soc. India, 4: 1-46. Radhakrishna, B.P. and Naqvi, S.M., 1986. Precambrian continental crust of India and its evolution. J. Geol., 94: 145-166. Radhakrishna, B.P., Devaraju, T.C. and Mahabaleshwar, 1986. Banded iron formations of India. J. Geol. Soc. India, 28: 1-17. Roy, A. and Biswas, S.K., 1979. Metamorphic history of the Sandur schist belt, Karnataka. J. Geol. Soc. India, 20: 179-187. Roy, A. and Biswas, S.K., 1983. Stratigraphy and structure of Sandur schist belt, Karnataka. J. Geol. Soc. India, 24: 19-28. Rao, T.G., 1992. Geology and geochemistry of BIFs of the central part of the Chitradurga schist belt, Karnataka, India. Unpublished Ph.D. thesis, OU, Hyderabad, 350 pp., unpubl. Ruhlin, D.E. and Owen, R.M., 1986. The rare earth element geochemistry ofhydrothermal sediments from the East Pacific Rise: examination of a seawater scavenging mechanism. Geochim. Cosmochim. Acta, 50: 393400. Schidlowski, M., 1990. Early evolution of life and economic mineral and hydrocarbon resources. In: S.M. Naqvi (Editors), Precambrian Continental Crust and its Economic Resources. Elsevier, Amsterdam, pp. 605-630. Sholkovitz, E.R., Piepgras, D.J. and Jacobsen, S.B., 1989. The pore water chemistry of rare earth elements in Buzzards Bay sediments. Geochim. Cosmochim. Acta, 53: 2847-2856. Simonson, B.M., 1982. Sedimentology of Precambrian iron formation with special reference to the Sokoman Formation and associated northeastern Canada. Ph.D dissert., John Hopkins Univ., Baltimore, MD, 359 pp. Simonson, B.M., 1985. Sedimentological constraints on the origins of Precambrian iron formations. Geol. Soc. Am. Bull., 96: 244-252. Srinivasan, R. and Sreenivas, B.L., 1972. Dharwar stratigraphy. J. Geol. Soc. India, 13: 72-83. Srinivasan, R., Manoj Shukla, Naqvi, S.M., Yadav, V.K., Venkatachala, B.S., Uday Raj, B. and Subba Rao, D.V., 1989. Archaean stromatolities from Chitradurga schist belt, Dharwar craton, South India. Precambrian Res., 43: 239-250.
Srinivasan, R., Naqvi, S.M. and Vasantha Kumar, B., 1990. Archaean shelf-facies and stromatolite proliferation in Dharwar Supergroup, North Kanara District, Karnataka, J. Geol. Soc. India, 35:203-212. Sverjensky, D.A., 1984. Europium redox equilibria in aqueous solution. Earth Planet. Sci. Lett. 67: 70-78. Swami Nath, J. and Ramakrishnan, M., 1981. Present classification and correlation. In: J. Swami Nath and M. Ramakrishnan (Editors), Early Precambrian Supracrustals of Southern Karnataka. Mem. Geol. Surv. India, 112: 23-38. Taylor, P.N., Moorbath, S., Chadwick, B., Ramakrishnan, M. and Viswanatha, M.N., 1984. Petrography, chemistry and isotopic ages of Peninsular gneiss, Dharwar acid volcanic rocks and the Chitradurga granite with special reference to the Late Archaean evolution of Karnataka craton, Southern India. Precambrian Res., 23: 349-375. Trendall, A.F., 1973. Iron formations of the Hamersley Group of Western Australia: type examples of varved Precambrian evaporites. In: Earth Sci. (UNESCO, Paris), 9: 257-268. Trendall, A.F., 1983. The Hamersley Basin. In: A.F. Trendall and R.C. Morris (Editors) Iron Formation: Facts and Problems, Elsevier, Amsterdam, pp. 69-130. Trendall, A.F. and Morris, R.C. (Editors), 1983. Iron Formation: Facts and Problems. Elsevier, Amsterdam, 558 pp. Venkatachala, B.S., Shukla, M., Sharma, M., Naqvi, S.M., Srinivasan, R. and Uday Raj, B., 1990. Archaean microbiota from the Donamalai Formation, Dharwar Supergroup, India. Precambrian Res., 47: 27-34. Venkatachala, B.S., Naqvi, S.M., Chadha, M.S. and 13 coauthors, 1991. Geology and geochemistry and paleobiology of Precambrian stromatolites and associated sedimentary rocks from the Dharwar craton: constraints on Archaean biogenic processes. Proc. IndoSoviet Symp. on Stromatolites and Stromatolitic Deposits. Wadia Inst. of Himalayan Geology, Dehradun, pp. 1-20. Walker, J.C.G., Klein, C., Schidlowski, M., Schopf, J.W., Stevenson, D.J. and Walter, M.R., 1983. Environmental evolution of the Archaean-Early Proterozoic Earth. In: J.W. Schopf (Editor), Earth's Earliest Biosphere Its Origin and Evolution. Princeton Univ. Press, Princeton, pp. 260-290. Walter, M.R. and Hofmann, H.J., 1983. The palaeobiology and palaeoecology of Precambrian iron formations. In: A.F. Trendall and R.C. Morris (Editors), Iron Formation: Facts and Problems. Elsevier, Amsterdam, pp. 373-400. Weiss, R.F., 1977. Hydrothermal manganese in the deep sea: scavenging residence time and M n / H e relationships. Earth Planet. Sci. Lett., 37: 257-262. Wilson, J.F. and Nutt, T.H.C., 1990. The nature and occurrence of mineralization in the Early Precambrian -
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164 Crust of Zimbabwe. In: S.M. Naqvi (Editor), Precambrian Continental Crust and its Economic Resources. Elsevier, Amsterdam, pp. 555-592. Zhai, M. and Windley, B.F., 1990. The Archaean and early
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Proterozoic banded iron formations of North China: their characteristics, geotectonic relations, chemistry and implications for crustal growth. Precambrian Res., 48: 267-286.
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Geochemical signatures of polygenetic origin of a banded iron formation (BIF) of the Archaean Sandur greenstone belt (schist belt) Karnataka nucleus, India C. Manikyamba, V. Balaram and S.M. Naqvi National Geophysical Research Institute, Hyderabad 500 007, India (Received April 25, 1991; accepted after revision June 30, 1992)
ABSTRACT Manikyamba, C., Balaram, V. and Naqvi, S.M., 1993. Geochemical signatures of polygenetic origin of a banded iron formation (BIF) of the Archaean Sandur greenstone belt (schist belt ) Karnataka nucleus, India. Precambrian Res., 61 : 137-164. Chert, ferruginous cherts, cherty banded iron formations (BIF), shaly BIF and shales are found interbedded at the eastern part of the Sandur greenstone belt. Cherts and shales are two end members in which deposition of Fe203 and A1203 in variable proportions has given rise to observed large-scale variation in major, trace and REE abundances. Iron, silica and REE have precipitated from the cold marine water which was enriched in these components by hydrothermal solutions from MOR vents. However, some minor amounts of Fe, Si and REE appear to have been brought in as dissolved load by rivers. Scatter in K, Mg, Ti, Cr, Ni, Zr, Hf, V, Sc, Y, Rb, Sr, Nb and Ta abundances and ratios indicate that the clastic components have both volcanoclastic and terrigenous sources. Large-scale variation in sumREE, moderate to pronounced La enrichment, positive Eu and negative Ce anomalies are characteristic of the majority of the samples of CBIF and SBIF. REE pattern shapes of CBIF are similar to hydrothermal solutions from EPR or Rs vents. Mixing of the FeO + SiO2 rich hydrothermal solutions with ambient sea water and clastic input has resulted in an observed large variation in LREE/ HREE ratio. However, total population of the samples shows the simultaneous increase of LREE and HREE from cherts to shales through CBIF and SBIF, and obliteration of hydrothermal signature in La content and the magnitude of the positive Eu anomalies. Negative Ce anomalies appear to be the result of a reaction with photosynthetic oxygen produced by bacteria which have developed stromatolites in the underlying formation. Hydrothermal and fluvial solutions provided Si, Fe and trace elements including REE to the cold marine waters of the seas. These hydrothermal solutions were emplaced in the relatively deeper part of the basin having a reducing and neutral to alkaline environment; and due to the thermal and chemical gradient convected towards the shoreline where photosynthesis was producing 02. Here, on the stable shelf region below the wavebase and photic zone FeO and Ce 3+ of these solutions were oxidized and mixed with clastic material of divergent origin. The Ce 4÷ was precipitated in the varves like SBIF and shales. Near the shoreline the environment was intermittently oxidizing but acidic most of the time and thus the precipitation of silica took place continuously, whereas precipitation of iron occurred intermittently due to the intermittent availability of oxygen or FeO or both. Our observations suggest that the BIFs of the Sandur belt are a product of hybridity between the hydrothermal emplacement of Si and Fe and divergent clastic sediments to ambient cold ocean water. The precipitation of Fe203 was biogenically mediated; a model combining these processes explains most of the feature of BIFs of the Sandur schist belt.
Introduction Banded iron formations (BIFs) are important but enigmatic rocks of the early and midCorrespondence to: S.M. Naqvi, National Geophysical Research Institute, H y d e r a b a d 500 007, India.
dle Precambrian. Their virtual confinement (James, 1983) to the Precambrian and absence of modern analogues have made these rocks extremely useful for the understanding of the ancient exogenic processes. It cannot be denied that if the early atmosphere was anoxic (Walker et al., 1983; Kasting, 1987)the FeO,
0 3 0 1 - 9 2 6 8 / 9 3 / $ 0 6 . 0 0 © 1993 Elsevier Science Publishers B.V. All rights reserved.
138 S i O 2 and some minor and trace element content of the BIF had been accumulating in the early Precambrian ocean reservoir for more than a billion years and the hydrothermal input seen as a source for BIF deposition may not be contemporaneous to the deposition. However, faster completion of the Wilson cycle in the Archaean as proposed by Hargraves (1986) suggests that the major part of the FeO and SiO2 content of BIF of a particular basin may be contemporaneous or it could have been carried over to another basin after the closure of an earlier one. Their origin remains highly controversial (Simonson, 1985). Although they have long attracted attention and a wealth of information is available on their associations, tectonic settings, petrology, mineralogy and economic potential, geochemical studies have been taken up only during past the two decades (Davy, 1983; Fryer, 1983; Dymek and Klein, 1988; Beukes and Klein, 1990; Derry and Jacobsen, 1990). Oxygen, carbon and sulphur isotopic data on the Archaean BIF and associated shales and carbonates are extremely scarce (Perry, 1983, Goodwin et al., 1985; Beukes et al., 1990; Schidlowski, 1990). Recent efforts in this field emphasise that a much larger and detailed geochemical and paleobiological data base should be created (Walter and Hofmann, 1983) for minimising the speculative part of the proposed genetic models. Sometimes "pure" samples made up only of chert and iron minerals are chosen for geochemical studies and the associated "clastics", regarded as "contaminants" are ignored (e.g. Derry and Jacobsen, 1990). The pure chemical fraction of BIF (i.e. the chert and iron minerals) has features which render it a poor recorder of minor element trends. Therefore, if an imprint of a genetic process is to be detected it will be preserved in the so-called "contaminants", which are an integral part of the BIF sequences. Although it is generally agreed that BIFs are chemogenic sediments, it is not certain what the source of Fe and Si was and how the ox-
c. MANIKYAMBAET AL.
ides, carbonates, sulfides and silica were precipitated. Other questions concern the mode of solution and transport of Fe and Si; the source of 02 needed for precipitation of Fe; and how microbands were formed. Are bacteria involved in producing these microbands (Nealson and Myers, 1990)? BIFs of several basins have been studied in detail, addressing these and some other questions, but the processes that form BIFs still remain somewhat elusive; genetic models range from exhalative (Gross, 1980; Simonson, 1982; 1985; Goodwin et al., 1985), evaporative (Eugster and Chou, 1973; Trendall, 1973; Garrels, 1987), biologically mediated precipitation (Cloud, 1973; La Berge, 1986; Nealson and Myers, 1990), and ocean upwelling (Holland, 1973; Drever, 1974 ). During the past two decades a number of summaries of the work carried out on BIF have been published (Trendall and Morris, 1983 ). In these summaries and interpretations most of the data used are from the BIFs of early Proterozoic basins although the occurrence of BIF in Archaean greenstone belts is a worldwide phenomenon (Condie, 1981). The Archaean BIF of greenstone belts of Canada (Goodwin, 1973 ), India (Radhakrishna et al., 1986), China (Zhai and Windley, 1990) and Africa (Anhaeusser, 1990; Wilson and Nutt, 1990) are relatively less studied than those of early Proterozoic age. The relevant information on the BIFs of India is extremely meagre, and especially the geochemical and genetic aspects have seldom been attempted. BIFs in India are confined only to Archaean greenstone belts. In spite of their wide development, Proterozoic sedimentary basins are generally devoid of BIFs (Radhakrishna et al., 1986). Attempts have begun recently to understand the geochemistry and genesis of the Archaean BIFs (Majumdar et al., 1982; Chakraborty and Majumdar, 1986; Devaraju and Laajoki, 1986; Manikyamba, 1992; Khan et al., 1992; Rao, 1992 ). In the Dharwar craton, deposition of BIFs has taken place at
BIF OF THE ARCHAEANSANDURGREENSTONEBELT
139
five stratigraphic horizons (Fig. 1 ) in strata formed between 3.5 (?) and 2.6 Ga ago (Naqvi et al., 1988 ). The most extensive and thick horizon is developed in a sedimentary sequence of stable shelf environment between 3.0 and 2.8 Ga ago (BIF3 in Fig. 1); generally known as the Bababudan Group which forms, the lower division of the Dharwar Supergroup (Swami Nath and Ramakrishnan, 1981 ). These BIFs are associated with QPC (quartz pebble conglomerate), current bedded-ripple marked quartzites and carbonaceous phyllites, phyllites and basic-acidic volcanics. They exhibit classical meso- and microbanding and thin in-
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terbedding with several types of extremely fine grained volcaniclastic and terrigenous sediments. The rock suite forming the Sandur schist belt are geologically correlated with Bababudan Group (Radhakrishna, 1983 ). It contains very large amounts of BIF and BMF (banded manganese formations) which are economically exploited on a large scale. These BIFs and BMFs were deposited under stable shelf conditions and are very closely associated with stromatolites and microbiota bearing cherts (Murthy and Reddy, 1984; Naqvi et al., 1987; Venkatachala et al., 1990, 1991 ). In this re-
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BANDED IRON FORMATION
(BIF 1)
MAFIC ULTRAMAFIC KOMATITES PRIMORDIAL BASIC CRUST ( 7 )
Fig. 1. Model stratigraphic sequence of greenstone belts (schist belts) in the Karnataka nucleus,India showing five successive cycles of BIF deposition in older and younger schist belts. The Sandur BIFs represent BIF3 which are found in stable shelf association along with stromatolites. (Modified after Naqvi et al 1988. )
LEGEND
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5 Dolomitic Stromolitic and Mn-Fe Carbonite
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9 Metavolcanics
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BIF OF THE ARCHAEAN SANDUR GREENSTONE BELT
gion a transition from stromatolitic carbonates to GIF (granular iron formation) to BIF is well exposed in the western part of the boatshaped Sandur schist belt (Fig. 2). The basal strata are not repeated on the eastern part where maximum development of BIF is seen. We present the results of our geochemical studies on the BIFs of the eastern part of the Sandur schist belt to elucidate their origin and to infer the physico-chemical conditions which prevailed during the deposition.
Summary of the geology of the Karnataka nucleus The geology of the Karnataka nucleus has been summarised by several workers (Swami Nath and Ramakrishnan, 1981; Radhakrishna and Naqvi, 1986; see also Naqvi and Rogers, 1987 ). The sequences are classified into older and younger schist belts ranging in age from 3.5 (?) to 2.5 Ga. Older and younger schist belts are separated by a widespread crust forming event at or around 3.0 Ga represented by the final consolidation of the Peninsular gneiss (Swami Nath and Ramakrishnan, 1981). Older schist belts are largely made up of mafic and ultramafic (spinifex textured komatiites) volcanics, amphibolites, quartzites (meta cherts), metapelites, paragneisses, and finegrained argillaceous and volcanogenic sediments including thin bands of recrystallized garnet-bearing BIFs. These belts are exposed near Holenarasipur, Nuggihalli, Krishnarajpet, Hadnur and several other places and have been metamorphosed to amphibolite facies (Naqvi and Rogers, 1983, 1987). They generally exhibit a complex deformational history (Naha et al., 1986). Sequences in the younger schist belts start with an unconform-
141
ity represented by quartz pebble conglomerate (QPC), current-bedded and ripple marked quartzites interbedded and overlain by basic volcanics, phyllites of variable composition with or without carbon and rare carbonate. They are followed by a thick sequence of BIF (and BMF at some places) with interbedded shales/phyllites. These basal portions of the younger schist belts are designated as the Bababudan Group, which is followed by the Chitradurga Group, made up of oligomictic or polymictic conglomerates, arkoses, quartzites, stromatolitic carbonate rocks, graywackes, basic-acidic volcanics and BIFs (oxide, carbonate and sulphide facies). This group is intruded by 2.6 Ga old K-granites. The entire succession of the Dharwar schist belts is not exposed at any single locality and shows an interfingering relation with polyphase gneisses (Naqvi, 1981 ). This aspect has given rise to many controversies about the structure and stratigraphy of the schist belts. Several conflicting models of stratigraphy have been proposed and the correlation of rock suites of various belts is often subjective due to discontinuous outcrop and scarcity of radiometric data. Naqvi et al ( 1988 ) proposed (Fig. 1 ) that at least five horizons of BIF deposition may be recognized. The maximum and most extensive deposition of BIF took place in the Bababudan Group around 3.0-2.8 Ga. The BIF of this Group can be subdivided into two classes: (1) predominantly oxide facies BIFs of Bababudan, Kudremukh and Kushtgi schist belts, lacking stromatolitic carbonate rocks, Mn formations, widespread carbonaceous phyllites and sulfide facies BIFs; and (2) the BIFs of the Bababudan Group in the Chitradurga, Shimoga and Sandur schist belts which are associated with stromatolitic carbonates (Ven-
Fig. 2. Geological map of the Sandur schist belt including stratigraphic sequence. Inset (A) shows location of the Sandur schist belt with respect to Karnataka, India. (B) Detailed geological map of the Sandur schist belt. The transition from quartzite-stromatolite sequence to BIF is exposed at the eastern limb (simplified and modified after Roy and Biswas, 1983). (C) Location of the samples analysed from exploratory bore hole cores and railway and road cuttings. Location of stromatolites, microfossils and carbon phyllites is also shown. (D) Map of a strip giving details of transition zone.
142
katachala et al., 1991 ), Mn formations and extensive carbonaceous phyllites. Earlier, Srinivasan and Sreenivas ( 1972 ) designated it as a separate Dodguni Group. The BIF of different regions mentioned above show relative abundances of different facies (Manikyamba, 1992; Rao, 1992; Khan et al., 1992). The BIF of Chitradurga for example are made up of oxide, carbonate, sulfide and mixed carbonate sulfide facies, whereas BIF of Bababudan and Kudremukh are predominantly of oxide facies and mixed oxide silicate facies.
Geology of the Sandur schist belt and depositional environment
Distribution of rock types, stratigraphy, structure and metamorphism of the Sandur schist belt has been worked out in detail by Roy and Biswas ( 1979, 1983 ). Geological map and stratigraphic succession after Roy and Biswas (1983) is given in Fig. 2B. The stratigraphic sequence in the belt is divided into 4 formations: Yeshwanthanagar, Deogiri, Donimalai and Nandihalli. The Yeshwanthanagar Formation consists of amphibolites, metapyroxenites, metagabbro, quartzite and quartz-mica schist. The Deogiri Formation is believed to conformably overlie the Yeshwanthanagar Formation (Roy and Biswas, 1983). It is essentially a sedimentary sequence of manganiferous graywacke-argillite with bands of quartzite and arkoses with siliceous dolomite (stromatolitic) and banded ferruginous chert (BIF). The Donimalai Formation is made up of banded magnetite/haematite, chert/jasper and metavolcanics ranging from basic to acidic, metagabbros, graywackes, fuchsite-quartzite, carbonaceous schists, metapelites, and sulfide facies BIF. The youngest formation, Nandihalli, is characterized by the absence of BIF and is made up of metabasalts and metagabbros, acid volcanics and a graywacke-argillite sequence (Roy and Biswas, 1983 ).
C. MANIKYAMBA ET AL.
Transition from quartzite to BIF The transition from quartzites (ripple marked Mn-bearing sandstones) to BIF through stromatolitic carbonate rocks, Mn formations and GIF is found at the interface of the Deogiri and Donimalai Formations. Stable shelf strata of the Deogiri Formation gradually merge into relatively deeper water shelf microbanded strata of the Donimalai Formation in which volcanic rocks are relatively more abundant (Naqvi et al., 1992). Gradation and development of microbanded BIF and GIF at 5 km stone on Yeshwanthanagar to Sandur road, and deposition of graywackes indicates a relatively deeper environment of deposition than the strata having current bedded and ripple marked quartzites and stromatolites of small dimensions exposed about 5 km west of this outcrop (Fig. 2). Stromatolites are generally not found below the photic zone and ferythmites are not formed above the wave base. Graded bedding in GIF is indicating fluctuations of the wave base. These structures are clear indications of variation in the depth of deposition. The strata with its facies changes is believed to be folded (Fig. 2 ) into a synform (Roy and Biswas, 1983 ). On the western limb of this synform predominance of stable zone strata is seen whereas on the eastern limb BIF, GIF, interbedded with shales/phyllites graywackes and metavolcanics (all part of a relatively deeper shelf) dominate. Mn formations, quartzites, carbonates are not repeated on the eastern part (Fig. 2 ). It is in this eastern part, that both oxide and carbonate facies BIF are abundant along with a zone of carbonaceous phyllite. At a few places along this part the zones of sulfide (pyrite) facies and mixed sulfide (pyrite), oxide (jasper, haematite) facies or mixed sulfide (pyrite), carbonate (siderite) facies BIF are found. All along the eastern part, the BIF, both the cherty and shaly type, are interbedded with pillowed metavolcanic rocks. The Donimalai Formation appears to be a relatively deeper facies variant of
143
BIF OF THE ARCHAEAN SANDUR GREENSTONE BELT
the Deogiri Formation. The transition from quartzite to BIF occurs within a sequence that from the base upward grades from quartzites to stromatolitic dolomites and finally to BIF through Mn dolomite, Mn shales, Mn sandstones and GIF. Contacts between quartzites and stromatolitic/Mn dolomites, between dolomite and shales and sandstones are gradational, and between shales, sandstones and BIF vary from gradational to sharp. Within the first layer of BIF, which conformably rests over sandstone and shales, GIF and BIF are found exhibiting both vertical and lateral transition into each other (Naqvi et al., 1992 ). Grading of iron minerals in the chert is seen in thin sections. Gradual settling of the fine dust of haematite-magnetite and sometimes siderite in the fine grained quartzite/chert matrix is exhibited until a layer of coarse iron minerals is formed with sharp contact with another graded layer. Such sequences are repeated several times. Stromatolites of the Sandur schist belt are oval to elliptical in shape with convex laminae (see Venkatachala et al., 1991 ). The larger axis of the ellipsoids at the bedding plane varies from 2 to 20 cm and the length of the columns in sections perpendicular to the bedding plane varies from 5 to 15 cm. The carbonates of stromatolites are Fe-dolomite, siderite and Mnsiderite interbedded with chert. The sedimentary structure in the underlying quartzite indicates that these and the carbonates in IF ( iron formation) were deposited under passive shelf conditions and that some IF, especially of the eastern part of the Sandur belt, were deposited at a deeper shelf below the wave base. An exposed cratonic area, allowing for clastic input in the form of fine grained mature arenites, graywackes and highly aluminous shales, indicating deep and intense weathering, probably existed west of the Sandur belt. Similar lateral and vertical transition from current-bedded and ripple marked quartzites to oxide and carbonate facies BMF and BIF, through stromatolitic dolomite and Mn shales is seen north of
the Srinkatte dome and west of Bhimasamudra in the Chitradurga schist belt. This transition from quartzites to BMF and BIF, through dolomitic stromatolites, Mn shales and carbonaceous phyllites (shales) is also exposed in the Shimoga schist belt near the Kumsi manganese mines and Kalche (Naqvi et al., 1992 ). A stable shelf environment of deposition is clearly evident from the morphology of stromatolites and other sedimentary structures in these localities (see Srinivasan et al., 1989, 1990). Furthermore, the discoveries of stromatolites in the basal sections, grading into BMF and BIF, indicate extensive Archaean biogenic activity in Karnataka nucleus and therefore biogenic mediation for the precipitation of iron and Mn in the overlying formations. Earlier such a transition between stromatolitic carbonates and BIF has been described from the early Proterozoic Transvaal basin of South Africa, namely the Cambellrand-Kurman transition zone. From a sequence correlated with the Transvaal basin in Botsvana/Zimbabwe also Mn bearing stromatolitic carbonates have been described from Mn deposits (Litherland and Malan, 1973; Beukes 1983). Transitions between GIF (Granular iron formation ) and BIF are a common phenomenon in the early Proterozoic stratigraphic record. However, so far they have not been reported from Archaean greenstone belts (Gole and Klein, 1981 ).
Age metamorphism and deformation The Sandur schist belt is intruded by Closepet Granite, which is locally known as Torangal Granite. Closepet Granite is dated at 2.42.6 Ga and Torangal Granite has yielded a RbSr isochron of 2.4 Ga. The basic and acid volcanic suites of the Sandur belt have yielded a thermal resetting Rb-SR isochron age of 2.4 Ga (Bhaskar Rao et al., 1992). The basic volcanics of the Kudremukh belt (which are stratigraphically equated with the basic volcanics of the Sandur belt) have yielded a Sm-Nd age
144
of ~2.9 Ga (Drury et al., 1983). The gneissose basement on which this belt is resting is dated elsewhere at 3.1 Ga (Taylor et al., 1984). In view of the available radiometric age data, the age of the Sandur schist belt may be somewhere between 3.1 and 2.6 Ga. However, the exact radiometric age for the rock formations of this belt is not yet known. Efforts in this direction are being continued at NGRI. The Sandur schist belt has been deformed twice, represented by two distinct phases of (folding) deformation, designated as SD 1 and SD2, of which the earlier one SDI is quite pronounced producing the NW-SE trending regional synformal structure of the schist belt (Fig. 2 ). SD 1 has resulted in the development of macroscopic folds, tight to isoclinal mesoscopic folds, schistosity, mineral lineation, fold mullion and pebble lineation. Second generation S D 2 deformation gave rise to broad warps, tight mesoscopic folds and crenulation cleavage (Roy and Biswas, 1983). This interpretation has been recently disputed by Mukhopadhyay and Matin (1990). This belt is predominantly characterized by greenschist facies of regional metamorphism, which increases to amphibolite facies towards the outer margins of the schist belt.
Characteristics of the Sandur belt BIF Iron formations of the Sandur belt consist of cherts (C), ferruginous cherts (FC), cherty BIF (CBIF), shaly BIF (SBIF), ferruginous shales (FSH) and shales/phyllites (SH). In fact there are two end members namely cherts (C), and shales/phyllites (SH), in which a variation in the proportion of SIO2, Fe203 and A1203 bearing minerals has produced the observed compositional scatter, as has been earlier found by Beukes (1980) for the BIF of the Transvaal Basin. The Donimalai BIF from the eastern part of the Sandur belt (Fig. 2 ) is made up of meso- and microbands of chert and iron minerals. Iron oxides (haematite/magnetite) dominate over carbonates (siderite/anker-
c. MANIKYAMBAET AL.
ite). Sulfide (pyrite) and silicate (cummingtonite/grunerite) bands are also found but they are not as abundant as the iron oxide bands. Two types of banding are noticed, one in which along with the chert a great concentration of iron minerals is found. In this type the Fe and SiO2 rich layers are microlaminated. The other type is one in which laminated chert and carbonates with sporadic development of magnetite/haematite is found. In such cases a gradual change from an iron-rich layer to a chertrich layer is found. The individual thickness of the iron-rich laminae with abrupt change to chert varies from a few mm to 2-4 cm. The thickness of the laminae showing a gradual decrease of iron minerals into chert laminae are about 5-10 ram. Fine dust of iron oxide is seen gradually diminishing into pure chert. Irregular, wavy bedding and domal structures identical to stromatolites are very common. "Neither stromatolites nor convincing microfossils are known from Archaean IF" (Walter and Hofmann, 1983). However, Hofmann et al. ( 1991 ) have found unequivocal stromatolites from the carbonate facies BIF of the Michipicoten formation, Ontario, Canada. SBIF are generally microbanded with chert, iron mineral and shales. At several places due to interbedding of chert and iron minerals with shales, high order fissility has developed.
Sampling and analytical techniques Most of the samples studied come from exploratory bore hole cores of the National Mineral Development Corporation or from the recent railway or road cutting in the mining area. The location of the samples is shown in Fig. 2C. Special care has been taken to collect and study only fresh unaltered samples. One hundred and two samples of the oxide facies BIF and associated cherts, ferruginous cherts, shaly BIF and shales were chosen. Ore microscopy for opaques and thin section study for silicate minerals were carried out for all samples. Composition of the constituent minerals have
145
BIF OF THE ARCHAEAN SANDUR GREENSTONE BELT
been determined by an electron microprobe analyser. Major elements were determined on fused pellets by a Phillips 1400 XRF. Trace elements and REE (for 102 samples) were estimated on ICP-MS and AAS. Fe203 in all samples was estimated on AAS and by colorimetric method. Among trace elements Zr and Y were determined on XRF. For details of the analytical techniques used and precision of the estimates refer to Govil (1985) and Balaram et al. (1991).
Petrology and mineralogy BIFs of the Donimalai region are characterised by simple mineralogy. CBIF consists of microcrystalline quartz/chert, haematite, specularite, with minor cummingtonite and grunerite. Magnetite is also found in some microbands. Hornblende is restricted to mixed silicate-oxide BIFs. Haematite and jasper are often present as ultrafine spheroids resembling coccoids (10 # m and less). Haematite and magnetite are also found as tetrahedral and cubic crystal clusters in the layers. Needle shaped haematite is seen in a few sections. In some samples veins of pyrite and chalcopyrite are present. Carbonates such as dolomite, ferroan dolomite, siderite, magnesian siderite, magnesian dolomite are found in a few sections in minor quantity. SBIFs are composed of chert and shaly layers with iron minerals. In addition to chert, haematite _ magnetite and jasper, they contain kaolin, sericite, and chlorite in varying proportions. In these formations minor carbonate is present. Ferruginous shales are made up of mainly kaolin _+ sericite, chlorite and iron oxides. Extremely fine grained and finely laminated varved types of shales/phyllites chiefly composed of kaolin are found interbedded with both types of BIFs. Some shale samples are thinly laminated and rhythmically bedded with chert having very little or no iron oxide, carbonate or sulphide minerals. In the iron ore zone of Donimalai where supergene enrichment has occurred and mining is in
progress, haematite, magnetite, specularite, martite, goethite and lepidocrocite are found. Syngenetic iron-rich bands are mainly made up of haematite and often magnetite. Haematite and chert appear to be primary minerals. Siderite occurs as primary and diagenetic varieties. Diagenetic siderite is interstitial and is often found as well developed rhombohedras. Magnetite appears to be of metamorphic origin. Details of mineralogical work form a separate paper (Manikyamba, 1992). Specularite, martite, goethite, limonite and lepidocrocite, etc. are secondary minerals. Samples containing secondary minerals were not included for petrological and geochemical studies.
Geochemistry Behaviour of major elements Mean and standard deviations of major, trace, and rare earth element contents of BIF and associated chert, shale/phyllites are given in Table 1. There is a large scale variation in the SiO2/Fe203 (T) molar composition. On the basis of SiO2, A1203 and Fe203 content the samples analysed are grouped as cherts (C), ferruginous cherts (FC), cherty BIF (CBIF), shaly BIF (SBIF) and phyllite/shales (SH). Samples containing less than 15% Fe203 and 80-75% SiO2 are classified as cherts. Ferruginous cherts contain 5-14.9% Fe203 and rest SiO2; whereas cherty BIF contains SiO2 and more than 15% Fe203; shaly BIF contains more than 15% Fe203 and more than 2% A1203; Shales (phyllites) contain less than 5% Fe203 and more than 5% A1203. A clear distinction between CBIF and SBIF can be made on the basis of A1203 content (Dimroth, 1986). Almost all major element constituents exhibit a large scale variation in their abundances in SBIF, whereas CBIF for a major part is made up of SiO2 and Fe203(T). The alkali content of both the varieties, with the exception of a few samples of SBIF, is lower than 1.0%.
146
C. MANIKYAMBAET AL.
TABLE 1 Average chemical composition of BIF from the Sandur schist belt Cherts (7) Mean
FC (8) SD
Mean
CBIF (51) SD
SBIF (27)
Mean
SD
Mean
Shales ( 9 ) SD
Mean
SD
Major elements (wt%) SiO2 TiO2 A1203 Fe203 MnO MgO CaO Na20 K20
P205
96.71 0.01 0.02 2.50 0.04 0.27 0.14 0.02 0.11
1.31 0.01 1.52 0.05 0.43 0.31 0.01 0.09
86.36 0.01 0.05 11.01 0.10 0.16 0.12 0.05 0.01 0.11
4.07 0.05 4.52 0.11 0.13 0.26 0.06 0.02 0.11
51.80 0.03 0.23 44.30 0.09 0.50 0.09 0.19 0.07 0.08
20.25 0.03 0.29 21.30 0.14 0.70 0.26 0.23 0.35 0.10
41.84 1.77 17.47 29.14 0.06 2.03 1.11 0.28 1.08 0.14
11.33 1.74 7.35 13.93 0.07 4.00 1.61 0.44 1.70 0.19
62.18 0.88 17.63 8.09 0.08 1.46 1.31 0.75 2.16 0.06
9.75 0.76 8.40 3.90 0.13 2.06 2.50 1.00 1.87 0.04
Trace elements (ppm) Sc V Cr Co Ni Cu Zn Rb Sr Y Zr Nb Ba Hf
0.16 7.53 11.40 98.56 7.53 23.75 3.76 0.07 1.61 0.46 0.71 0.15 1.08 0.09
0.11 11.86 12.09 15.99 3.61 41.78 3.96 0.04 1.63 0.37 0.47 0.06 0.72 0.10
0.38 2.36 5.48 73.01 10.73 35.54 3.55 0.18 0.54 0.67 0.92 0.22 3.61 0.12
0.36 4.52 5.10 33.02 5.30 81.96 2.25 0.12 0.43 0.39 0.70 0.13 3.79 0.08
1.30 11.28 55.74 33.72 10.23 12.51 11.21 2.66 3.78 1.78 4.01 0.29 14.55 0.18
1.40 16.62 82.00 28.54 5.70 16.12 10.47 11.92 5.41 1.72 9.05 0.25 26.32 0.28
20.58 210.20 276.01 16.79 82.91 76.13 59.89 40.89 43.09 29.21 181.16 6.27 172.20 3.64
8.83 145.34 187.84 19.58 106.90 66.92 60.96 77.12 49.52 26.99 164.58 5.03 268.99 3.40
23.82 307.79 354.94 24.70 96.02 149.66 98.82 86.52 101.53 19.50 258.73 9.73 249.73 3.75
13.37 220.20 368.35 16.84 81.41 156.12 88.94 96.01 202.78 14.58 262.65 8.05 221.22 2.17
La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu ~REE
3.57 1.14 0.66 5.79 0.49 0.16 0.16 0.03 0.53 0.10 0.29 0.03 0.15 0.02 13.32
4.80 1.13 1.41 9.13 0.89 0.23 0.07 0.02 0.48 0.08 0.21 0.01 0.07 0.02 16.18
3.66 0.97 0.10 0.24 0.14 0.09 0.08 0.04 0.19 0.04 0.15 0.02 0.14 0.02 6.11
9.65 1.47 0.15 0.17 0.13 0.06 0.04 0.06 0.14 0.03 0.10 0.02 0.10 0.02 12.21
5.97 2.10 0.26 1.30 0.34 0.18 0.26 0.14 0.37 0.07 0.25 0.03 0.19 0.03 11.42
9.41 1.92 0.26 1.36 0.32 0.13 0.26 0.16 0.28 0.05 0.35 0.02 0.14 0.02 10.70
20.29 39.11 3.78 20.39 4.13 1.51 3.77 0.64 4.07 0.89 2.73 0.38 2.68 0.58 96.78
22.39 40.88 3.82 22.32 4.04 1.55 3.06 0.71 4.60 0.94 2.72 0.37 2.95 0.65 107.37
14.49 38.33 2.73 9.35 2.20 0.73 2.22 0.40 2.22 0.47 1.49 0.21 1.48 0.24 74.82
9.68 30.09 1.87 6.12 1.55 0.48 1.39 0.24 1.43 0.27 0.77 0.10 0.81 0.17 49.80
( 1 ) Number of samples for each type analysed is given in parentheses. (2) Data for the samples will be made available on request.
However, MgO and CaO are fairly abundant and show great scatter due to the presence of carbonates (Table 1 ). A1203 appears to be one of the most diagnostic constituents o f the dif-
ferent varieties of IF in the belt. T i O 2 also is found to distinguish between the two main varieties of the BIFs. MnO and P205 are of very low abundance and exhibit an overlap be-
BIF OF THE ARCHAEAN
SANDUR
GREENSTONE
147
BELT
elevated to more than 0.4% and 0.2% respectively (Fig. 3A). Even at elevated concentrations of TiO2 and A1203, as in SBIF, the scatter in TiO2 and A1203 concentrations shows that the clastic input for different layers and at different levels has been inconsistent and may probably represent different types of source rocks. Thus the A1203 and TiO2 abundances in Sandur BIFs indicate that clastic sedimentation from variable compositional sources has also occurred simultaneously with chemical precipitation of SiO2 and Fe203, and the quantity and quality of this clastic sedimentation has been responsible for the observed compositional variation in both types of BIF. The
tween the two varieties (Table 1 ). Elevated A1203 and TiOe reflects a clastic component (Ewers and Morris, 1981 ). Most of our CBIF samples have 0.01 to 0.1% TiO2 with A1203 values ranging between 0.01% and 1.35%. In SBIF both the TiO~ and A1203 values are elevated from 0.2 to 8.5% and from 3.4% to 33% respectively. CBIF and SBIF form seperate groups, though the TiO~/A1203 ratio for both types of BIF is always less than one (Fig. 3A). Linear relationships between A1203 and TiO~ of S macrobands has been observed by several workers (Ewers and Morris. 1981). In our samples we observe crude linearity, but only after the concentration of TiO2 and A1203 are 100
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148
scatter in the TiO2/AI203 ratios (Fig. 3A) and the composition of the thinly laminated varvelike shales interbedded with oxide facies BIF at Donimalai, containing more than 15% A1203 due to the presence of kaolin, clearly indicate that terrigenous sedimentation continued with chemical precipitation. In thick layers of shales interbedded with CBIF several sedimentary cup-like structures, are present indicating deposition of shales on shallow shelves as such cupsuate structures are formed by fluctuations of the wave base (Naqvi et al., 1992). The TiO2/Zr of the shales and SBIF are similar to sedimentary clays and slates. Higher TiO2/ A1203 ratios in some of our samples also indicate that in addition to terrigenous sedimentation mafic volcanic activity also contributed to the shaly part of the BIF. The MgO content of the SBIF has a bimodal distribution. Twenty-two samples have a MgO content between 0.01 to 0.86% whereas five samples have from 2.9 to 13.4% MgO. A considerable amount of MgO is in the form of ferromagnesian minerals like chlorite and amphiboles contributed from basic volcanism, probably volcanic ash. Generally samples of SBIF having very high A1203 contents have low MgO abundances and vice versa (Table 1 ).
D&tribution and behaviour of trace elements Data for fourteen trace elements are given in Table 1 indicating a large scatter in their values as reflected by their standard deviations. The chromium content of both types of BIF and associated shales/phyllites generally varies from 4 to 400 ppm and their Ni abundances range from less than 2 ppm to 88 ppm. In one sample of shale Cr is found at 976 ppm level. A wide scatter of Ni and Cr contents (r=0.63) of these rocks is illustrated in Fig. 3B, where it can be observed that on the basis of Ni and Cr abundances and their ratios, no distinction can be made between CBIFs, SBIFs and shales. Even cherts and ferruginous chert, exhibit fairly high Ni and Cr abundances, es-
C. MANIKYAMBA ET AL.
pecially chromium. Ferruginous sediments, formed by chemical precipitation from normal seawater do not have such high concentrations of these elements. Therefore, higher abundances of these elements indicate that enrichment in Cr and Ni contents is probably due to the contribution of volcaniclastic input. The presence of 88 ppm Ni in cherts and 300 ppm in shales, 976 ppm Cr in shales and 300 ppm in CBIF probably cannot be explained satisfactorily unless the fumerolic and explosive volcanic activity producing ash is assumed to occur in this basin. Ni abundances vary between 2 and 250 ppm and Zr varies from 0.1 ppm to 716 ppm in cherts to SBIF through CBIF. A gradual and simultaneous increase ( r = 0.28 ) of these two elements is seen from CBIF and cherts to SBIF and shales (Fig. 3C). Zr abundances of SBIF and SH are distinctly higher than those of CH, FH and CBIF and thus the N i / Z r ratios and behaviour demonstrate the clastic contribution to these rocks. Maximum concentration of Zr found in SBIF is 280 ppm and that of SH is 700 ppm. In CBIF it varies between 0.1 and 20 ppm. Evidently Zr is contributed from terrigenous sediments. In the Isua BIF also Zr values are highest in the aluminous variety (Dymek and Klein, 1988). The Zr and Cr relationship (r=0.52) brings out this aspect more clearly (Fig. 3D). In general both Zr and Cr enrichment has taken place simultaneously in CBIF and SBIF. SBIFs relative to CBIF are enriched in both Cr and Zr suggesting that hydrothermal, volcaniclastic and terrigenous sources have contributed to the observed composition of BIF. This is further substantiated from the Zr and Cr data of shales, where shales enriched in MgO have higher Cr and shales relatively depleted in MgO and enriched in A1203 are more enriched in Zr (Fig. 3D). Similarly very low concentrations ofZr ( < 3 ppm ) in the majority of C, FC, and CBIF samples (Fig. 3D) are suggestive that contribution of solutions derived from chemical leaching of the land outside the basin to the Fe and Si of the
BIF OF THE ARCHAEAN SANDUR GREENSTONE
149
BELT
BIF may not be significant but it cannot be ruled out as Zr would have been left in the leached rock suites and this will come to the basin only as clastic component. Wherever terrigenous contribution has increased the Zr value has increased too. This is further strengthened by the Zr and Y behaviour (Fig. 3E) and Zr and YREE patterns (Fig. 3F). A gradual increase of both Zr and Y (r=0.73) is observed (Fig. 3E) and also no compositional gap between the two is seen. Similarly a sympathetic positive relationship (r=0.66) between Zr and YREE (Fig. 3F) with no gap in between suggests that mixing of chemical precipitates and clastics has been the main reason for the observed compositional variations. Zr and Hf both increase simultaneously ( r = 0.77) in our samples and show an excellent linear relationship (Fig. 4A) between their values from CH, FC, CBIF, SBIF, and SH. There is no gap between cherts to shales through CBIF or SBIF. This suggests that a continuous supply of a varying quantity of the divergent types of terrigenous and volcaniclastic debris has been maintained to the basin and the mixing of hydrothermal solutions with clastics derived from two divergent sources has taken place. Zr, Y and Hf are provided by felsic rocks and if felsic volcanic ash would have provided these elements, appreciably high concentrations of K20 and NazO of the order of their abundances in felsic volcanics should be preserved in SBIFs and SH, as the settling of volcanic ash being a rapid processes without weathering and fluvial abrasive transportation will not allow removal of alkalies from feldspars. Contrary to this, only in four samples of SBIF and one sample of SH relatively higher concentrations of alkalies are found (Table 1 ). Abnormally high AI203 contents from 8 to 33% in SBIF and shales and low alkali contents in the corresponding samples, suggests that probably a considerable part of the clay of the BIFs has been derived from a deeply weathered continental source and not from within the basin felsic volcanic source.
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150
XRD data have shown that most of it is kaolin and hence indicates to a feldspar-rich source. Rb and Sr in all geochemical process generally accompany alkalies. In general CBIF and SBIF appear as separate groups on the Rb-Sr plot (Fig. 4B ). Most of the CBIF contain Rb below one ppm. A few samples have Rb more than 13 ppm. Whereas the Rb content of SBIFs generally remains below 100 ppm but goes up to 257 ppm in two samples. Similarly the maximum Sr content of CBIF is found to be only 19 ppm and in SBIF it goes up to 66 ppm. In general CBIF have lower abundances of Sr, whereas SBIF and shales have high Sr content which reaches up to 638 ppm in sample No. SM 17. The large and random scatter of Rb and Sr (r=0.12) in both varieties of BIF, shows a multicomponent contribution for the constituents of BIF (Fig. 4B). Vanadium and scandium show sympathetic linearity (r=0.87) and a noticeable distinction between CBIF and SBIF. Maximum enrichment of V and Sc has occurred in SBIF and a few samples of chert and ferruginous chert and CBIF are extremely depleted in both V and Sc. A complete gradation of CBIF into SHBIF is seen on the V-Sc plot (Fig. 4C ). Similar distinctive behaviour is seen in the Ta and Nb distribution (Table 1 ). Cu and Zn concentrations and sulfide veining in source samples indicate a late stage hydrothermal syngenetic or postgenetic emplacement of these elements.
REE geochemistry REE are estimated in 102 samples but sixty seven chondrite normalised patterns are presented in Figs. 5 and 6. Based on REE pattern shape and La, Ce and Eu anomalies, these samples are divided into the following twelve groups. Since both La and Ce enrichment and/ or depletion has taken place, N d N / Y b N ratios are used, instead of La and Lu ratios to get an idea about REE fractionation. Groups 1 to 8 are of the samples belonging to chert, FC, and
C. MANIKYAMBAET AL.
CBIF. Groups 9 and 10 are of SBIF and Groups 11 and 12 represent shales. (1) Group 1 (Fig. 5A) consists of eight samples with extremely depleted EREE, loworder La enrichment, positive Eu anomalies, and NdN/YbN= 1 and LREE/HREE > l, LREE from La to Sm show fractionation, HREE are flat. (2) Group 2 (Fig. 5B) consists of 4 samples with extremely depleted YREE, low-order La enrichments, negative Ce anomalies, negative Eu anomalies. NdN/YbN 1. Sample No. SM24 is a pure chert exhibiting negative Ce and Eu anomalies. (3) Group 3 consists of four samples (Fig. 5C) with extremely depleted ~REE, depletion in La, positive Ce and Eu anomalies, NdN/YbN ~<1, and LREE/HREE slightly < 1. (4) Group 4 consists of seven samples (Fig. 5D) in which all REEs, except La and Eu, are extremely depleted. La is highly enriched, Ce is depleted and shows strong negative anomalies. Eu exhibits weak to strong positive anomalies. NdN/YbN>~ 1, including La, HREE/ L R E E < I , excluding La, L R E E / H R E E < I . One sample, No. CM40, shows negative Eu anomaly. From Nd to Lu the pattern shapes are fiat. NdN/YbN and HREE/LREE, excluding La, are close to one. This group of samples is very important because of its high-order La enrichment. (5) Group 5 consists of six samples (Fig. 5E) having moderate to very high ZREE abundances. This group is characterized by very high order La enrichment and negative Ce anomalies, positive Eu anomalies. NdN/ Y b N > I , excluding La and Ce, LREE/ HREE > 1. Sample No. MM24 shows very high N d N / Y b N and LREE/HREE ratios. Sample No. CM36 shows a very strong negative Ce anomaly. In fact negative Ce anomalies are the most characteristic and important feature of this group. (6) Group six (Fig. 5F) consists of six samples which are characterized by moderate ~REE abundances with positive Eu anomaly,
BIF OF THE
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Fig. 5. (A) Extremely depleted REE pattern in ferruginous cherts and cherty BIF showing slight La enrichment and strong positive Eu anomalies. (B) Extremely depleted REE patterns in cherts and cherty BIF showing negative Ce and Eu anomaly. (C) REE patterns with slight La depletion strong Eu enrichment and weak positive Ce anomalies. (D) Extremely depleted REE patterns showing high order La enrichment, slight negative Ce anomalies, strong positive Eu anomalies, with almost flat pattern shape between Nd to Yb. (E) REE patterns moderately enriched in •REE with very strong negative Ce anomalies, showing fractionated pattern from Pr to Lu and positive Eu anomalies. Four samples show Gd depletion. Sample No. MM24 shows extremely fractionated pattern. Extrapolation from Pr to La in each case shows enrichment in La. (F) REE patterns of cherty BIF with moderate ~ REE abundances, showing fractionation from La to Lu with positive Eu anomalies. No La enrichment or Ce depletion is recorded in these samples. For legend see also Fig. 3.
NdN/YbN> 1 and LREE/HREE> 1. Sloping REE patterns due to slight depletion in HREE and positive Eu anomaly are characteristic features of this group. No La enrichment and Ce depletion are recorded in this group. (7) Group seven (Fig. 6A) consists of six samples which are characterized by moderate ZREE abundances, fractionated pattern from
La to Sm, positive Eu anomalies, flat HREE except in one sample (SM 19), NdN/YbN> 1 and LREE/HREE > 1 fractionation from La to Sm is the most characteristic feature of this group. (8) Group eight (Fig. 6B) consists of five samples with moderate •REE abundances, moderate negative Ce anomalies and moder-
152
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ate positive Eu anomalies. Three samples show fiat pattern shape from Ce to Sin, whereas two other samples show fractionated pattern shape from Pr to Lu, NdN/YbN and LREE/ HREE> 1. A general fractionation pattern shape is the significant characteristic of this group. (9) Group nine consists of seven samples of SBIF (Fig. 6C) this group is characterized by enrichment in ~REE compared to CBIF, no La or Ce anomalies, weak positive or no Eu
anomalies and NdN/YbN and LREE/ HREE> 1. (10) Group ten (Fig. 6D) consists of eight samples of SBIF having enrichment in Y.REE with fiat to slightly sloping pattern shapes and weak positive and negative Eu anomalies. Enrichment in total REE is the main distinguishing feature of this group. NdN/YbN and HREE/LREE< 1. In this group 2 samples which are not plotted in this illustration also exhibit positive cerium anomalies.
BIF OF THE ARCHAEAN SANDUR GREENSTONE BELT
( 11 ) Group eleven (Fig. 6E) consists of 3 samples of shales showing enriched ZREE with strong positive Ce and weak negative Eu anomalies. N d y / Y b y and H R E E / L R E E < 1. The appearance of positive cerium anomalies is one of the most outstanding aspects of this group. (12) Group twelve (Fig. 6F) consists of three samples of shales/phyllites with enrichment in EREE relative to CBIF. They show both very weak positive and negative Eu anomalies. NdN/YbN and LREE/HREE are slightly greater than one. These patterns of shales resemble those of NASC.
Enrichment of XREE: consequence of clastic contribution to BIF The absolute REE content of the Archaean Sandur schist belt BIF shows great variation in abundance, from very low to very high (Table 1 ). Such patterns (Figs. 5 and 6) represent the end product of a complex series of events that record the properties of the solutions that precipitated them with clastic sediments. A contemporaneous input of clastic material appears to have greatly effected the absolute abundance of REE in these rocks. The overall shape of the REE patterns of the oxide facies BIF samples comprising only chert and iron minerals is generally similar to patterns of modern seawater, with the exception of La, Ce and Eu which show anomalous behaviour. Derry and Jacobsen (1990) have also found a similarity in the overall shape of REE patterns of Archaean oxide facies BIF and modern seawater, with the exception of Ce and Eu. Despite a large scatter in the LREE/HREE ratio (Fig. 7A) ( < 1 to > 4 0 ) , LREE and HREE increase simultaneously when the entire population ofCH, FC, CBIF, SBIF and SH is considered together. The N d / Y b ratio shows the same scatter and a simultaneous increase of Nd and Yb from pure CBIF to SBIF. The lowest Nd and Yb values are found in chert and pure CBIF (irrespective of the mutual varia-
15 3
tion of SiO2 and Fe203 content, both these oxide comprise almost 100% of the sample) whereas the highest values of Nd and Yb occur in shales and SBIF (Fig 7B). A complete gradation between these values is seen in the LREE/HREE and N d / Y b plots. A change in ZREE from < 1.00 ppm to > 421 ppm from CBIF to SBIF is apparent in Figs. 5 and 6. These data clearly demonstrate that the varied amount of clastic input into the chemogenic sediments has increased the YREE, and the variation in LREE/HREE and Nd/Yb. Fractionation from La to Lu, a characteristic of felsic rocks is not observed and the REE patterns and shapes with La, Ce and Eu anomalies appear to be coming from from three different sources. The modern oceans receive most of their REEs from three sources: (1) the dissolved load of rivers (Goldstein and Jacobsen, 1988a, b), (2) hydrothermal alteration of the oceanic crust (Michard and Albarede, 1986), and ( 3 ) solutions derived from sediments undergoing diagenesis. The magnitude and pattern of the diagenetic REE contribution is uncertain, but on global scale it is probably small relative to erosional and hydrothermal fluxes (Sholkovitz et al., 1989). REE data from metalliferous sediments from the East Pacific Rise (EPR) show evidence of scavenging of REE by Fe oxyhydroxides and a clear pattern of gradual decrease of LREE depletion with distance from the ridge axis (Ruhlin and Owen, 1986). According to Derry and Jacobsen (1990) these relations imply that metal oxy-hydroxides precipitating from the hydrothermal plumes are preferentially enriched in LREEs such that by the time the REEs in the plume are far removed from the ridge axis, the residual water mass is LREE depleted. The similarities between the REE patterns of modern metalliferous sediments and the oxide facies BIFs suggest that scavenging by Fe hydroxides was the mechanism responsible for the incorporation of the REEs into BIF. Derry and Jacobsen ( 1990 ) further conclude that the overall trend
154
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Fig. 7. (A) Mutual relationship between LREE and HREE indicating simultaneous increase of LREE and HREE from ferruginous cherts to SBIF. (B) Nd-Yb relationship showing the scatter in the ratio from 1 to 40 and simultaneous increase of Nd and Yb from CBIF to SBIF. (C) Eu-Eu* relatioship indicating the change in the nature and magnitude of Eu anomalies from CBIF to SBIF and SH. (D) (ZREE-La) are plotted against La abundances to show the highly La enriched characteristics of some CBIF samples, (E) La-Fe203 relationship shows that between 40 and 60% of Fe203 a large number of samples of CBIFs have high order La content similar to those of SBIF in which anomalous behaviour of La is not recorded. (F) Ce-Fe203 relationship shows that at 40-60% level of Fe203, maximum Ce depletion has taken place. For legend see also Fig. 3. o f REEs in BIFs are compatible with a mechanism o f oxide facies BIF formation in which Fe hydroxides precipitated from seawater with a REE pattern similar to the modern ocean (except for Ce and Eu ). The variability o f REE patterns in the IF samples probably results from differences in scavenging efficiency, with the most depleted patterns resulting from the most efficient REE removal. Our data, however, indicate that the variability o f REE patterns in BIF samples results from the mixing o f the clastic c o m p o n e n t s with that o f the sea-
water in which hydrothermal solutions were emplaced at an Archaean mid-oceanic ridge (AMOR).
Eu, La and Ce anomalies." possible signatures of AMOR hydrothermal activity and biogenic oxidation Eu anomalies Except 5 samples o f C, CBIF and 7 samples o f SBIF all samples o f BIF, chert, and ferruginous chert from the Sandur belt exhibit posi-
BIF OF THE ARCHAEAN SANDUR GREENSTONE BELT
tive Eu anomalies with variable magnitude (Fig. 7C). Such positive Eu anomalies are observed in BIFs irrespective of their age or locality and can only result from the input of Eu enriched hydrothermal fluids into the water column of the oceans (Graf, 1978; Fryer et al., 1979; Barret et al., 1988). Kerrich and Fryer (1979) and Fryer et al. (1979) suggested that the Eu enriched nature of Archaean and early Proterozoic seawater was due to a major contribution of strongly reducing hydrothermal fluid discharging on to the sea floor during these times. Experimental work of Flynn and Burnham ( 1978 ) supports this view. Modern seawater does not typically display Eu anomaly except near hydrothermal vents. REEs in present day ocean water occur in exceedingly low concentrations with patterns that show depletion in Ce and Eu and LREE < HREE. These characteristics are interpreted to indicate a control by continental input on the abundance of Eu, while the relative depletion of the LREE, especially Ce, is due to selective scavenging of seawater REE during the formation of Mn nodules. The source of anomalous Eu concentration in the BIFs requires that the hydrothermal REE flux comprised a much larger fraction of the total REE flux than today (Derry and Jacobsen, 1990). Present day hydrothermally active regions in EPR and M P R (MidPacific Rise) are characterized by low EREE abundances, extreme positive Eu anomalies and LREE enrichment, because the REEs of the fluids are controlled by high temperature rockwater interactions in the volcanic sequence at depths (Michard et al., 1983; Campbell et al., 1988). Such conditions favour the stability of Eu 2+ in aqueous solution and the development of positive Eu anomalies (Sverjensky, 1984). The signature of hydrothermal REE appears to die out rapidly by mixing with seawater and the fact that Pacific sea water near EPR is relatively enriched in Eu compared to North Atlantic water may represent the effect of a hydrothermal component (Klinkhammer et al., 1983). Dymek and Klein (1988) have
15 5
computed that a 1:100 mix of North Atlantic seawater and EPR hydrothermal fluid is able to retain positive Eu anomaly. This 1:100 mix of EPR and NASW (North Atlantic Sea Water) yielded patterns similar to uncontaminated BIFs of Isua Formation. The deep (2500 m) hydrothermal plume waters directly above the EPR at 19°S (Klinkhamer et al., 1983) have an REE pattern completely different from that of the vent waters, yet still distinct from normal seawater (De Baar et al., 1985). Although vent driven water in this plume is strongly diluted by ambient sea waters, a hydrothermal component is recognized on the basis of tracers such as Mn and He 3, (Weiss, 1977; Lupton and Craig, 1981 ) for distance of hundreds meters above and hundreds of kilometers west of EPR. By contrast, the hydrothermal Eu anomaly is lost immediately following the discharge ofhydrothermal solutions into seawater. The Eu anomaly is also absent in metalliferous sediments from various ridge related and DSDP sites in the eastern Pacific (Ruhlin and Owen, 1986; Barrett and Jarvis, 1988). Clearly this anomaly is not transmitted in modern oxidized open oceans, not even to plumes only a few hundred meters above the ridge axis. This is apparent because of rapid oxidation of hydrothermal Eu 2÷ to E u 3 + in the water column; any precipitates formed also absorb sea water REE that lack a positive Eu anomaly. However, ifArchaean oceans were relatively reducing, transmittance of positive Eu anomaly away from its hydrothermal source seems plausible (Barrett et al., 1988). The total range of the REE content of the BIFs of the Sandur schist belt is < 1.00 p p m to >421 p p m (Table 1) and as the )~REE increase from CBIF to SBIF, the amplitude of the positive Eu anomaly gradually diminishes and even becomes slightly negative (Fig. 7C ). This shows that the REE patterns and Eu anomalies of BIFs of the Sandur schist belt are greatly affected as a result of mixing, not only with the seawater but also with the clastic component.
156
La enrichment and Ce depletion Apart from the positive Eu anomalies, which are characteristic of BIF all over the world (Fryer, 1983; Derry and Jacobsen, 1990), the most important and significant aspect of the REE data from the Sandur belt BIF are La enrichment and negative Ce anomalies. Thirtysix samples of CBIF, chert and ferrugenous chert exhibit La enrichment of low to very high order. Since La is behaving abnormally, its concentration is deducted from YREE and thus a plot of ( ~ R E E - L a ) against La is made to illustrate the behaviour of La in the total population of samples analysed. Among the samples of SBIF where YREE is very high, La and ( E R E E - La) show a well defined sympathetic linear relationship whereas most of the CBIF, chert and FC samples fall below that line of linearity showing La enrichment (Fig. 7D). In CBIF a maximum variation in the La content (from 0.025-47 ppm) is found at 40-60% Fe203 content (Fig. 7E). Ce is also anomalous and at 40-60% Fe203 many samples show strong Ce depletion. The maximum depletion in the Ce (0.08 ppm) content is recorded in sample No. CM36 having about 85% Fe203 (Fig. 7F). Ce depletion is excellently projected in Fig. 8A in which ~ R E E - ( L a + C e ) is plotted against Ce. Between 0.3 to 1.0 ppm Ce depletion level the 3~REE- ( L a + C e ) varies from 0.56 to 7.0 ppm. SBIFs depict a very clear cut sympathetic linear relationship between Ce and ~ R E E - ( L a + C e ) (Fig. 8A). Nd is not showing any abnormal behaviour in our samples. Therefore, ratios between Nd, Ce and La in a ternary plot illustrate the La enrichment and Ce depletion (Fig. 8B). The other locality from where La enrichment consequent of hydrothermal spiking has been recorded is the Archaean Greenstone belt of Canada. Barrett et al. (1988) have interpreted that this La enrichment is caused by the La spiking from hydrothermal solutions. In support of their inferences, Barrett et al. have cited an example of Red Sea hydrothermal sediments where all chemical precipitates show
C. MANIKYAMBA ET AL.
distinct La enrichment. They postulate that IF with La enrichment and strong positive Eu anomalies have been formed from reduced bottom waters carrying a significant hydrothermal signal. Accordingly, IF with no La enrichment and smaller Eu anomalies may have been precipitated from solutions that were diluted strongly by ambient ocean water during periods of lessened discharge of hydrothermal solutions or from solutions far removed from vent sources. Barrett et al. (1988) proposed that the ambient seawater, although probably characterized by positive Eu anomalies, had an otherwise near fiat pattern with no La enrichment. Archaean seawater in general is expected to be more reducing, which should be favourable for the retention of hydrothermal signatures, but the precipitation of iron oxides in jaspery chert and haematite/magnetite indicate that away from the locus of hydrothermal input an oxygenated environment existed and that the FeO + SiO2 were transported from the reduced environment and were precipitated in the oxygenated environment. During this transport, mixing with the ambient seawater and clastic sediments took place resulting in the obliteration of La spiking in many samples. Whenever mixing of ocean water and clastic sediments was incomplete and ineffective La and Eu enrichment is pronounced. On the other hand, samples with less pronounced or nil La and/or Eu enrichment indicate that ocean water and clastic mixing with the hydrothermal solutions has obliterated and masked the hydrothermal signature. In some samples of chert and cherty BIF (Fig. 5B)extremely depleted ~REE abundances are associated with negative Ce and negative Eu anomalies. These samples could have been deposited from oxygenated ocean waters near the shoreline having their Fe and REEs brought by rivers as dissolved load from continental source to a zone where oxygen was produced more rapidly by stromatolites developing photosynthetic bacteria. Seven samples of FC and CBIF also show La depletion and Ce enrichment. These data
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Fig. 8. (A) Relationship between ~REE- (Ce+La) and Ce. The linear relationship is shown by the straight line from SBIF to ferruginous chert, most of the CBIF samples are falling above the line indicating Ce depletion in CBIFs and ferruginous cherts suggesting oxidation of the hydrothermal solutions. (B) shows La:Ce:Nd ratio distribution in the total population of the samples analysed. Since La and Ce both are showing abnormal behaviour with respect to Nd this ratio is extremely useful for illustrating La enrichment and Ce depletion with respect to Nd. (C) Relationship between EREE and AI203 illustrates that variation and scatter in REE is consequent of mixing with the divergent clastic source. (D) Relationship between EREE and sum of Ni, Cu and Co. Fields of hydrothermal deposits and hydrogenous sediments are after Klein and Beukes (1989). For legend see also Fig. 3. indicate that precipitation o f CBIF has taken place in a weak and intermittently oxygenated environment. The inference that both types namely reducing and oxidizing environments prevailed in Archaean oceans is further substantiated by the size o f the negative Eu anomalies found in some cherts and cherty BIF samples. Based on limited data available at that time Fryer ( 1983 ) inferred that significantly anomalous Ce abundances are not known from Archaean IF, whereas in Proterozoic BIFs Ce is definitely anomalous with examples o f both enrichment and depletion being observed. Thus, Fryer (1983) suggested that Ce was being separated from the rest of the REE by
oxidation to the Ce 4+ state, at least back to the early Proterozoic, which would require that areas of moderately to strongly oxidizing conditions must have existed in the Proterozoic seas. REE patterns of IF of the early Proterozoic Transvaal Supergroup are characterized by strong depletion in LREE relative to H R E E , strong negative Ce anomalies and marked positive Eu anomalies (Beukes and Klein, 1990). Also, Ce depletion in the modern seawaters is attributed to the removal of Ce to Mn nodules by oxidation of Ce 3+ to Ce 4+ (Addy, 1979) for which a typical marine p H of ~ 8 and Eh values o f 0.15-0.45 V are required. The early Proterozoic IFs, which exhibit anomalous Ce, are also enriched in Eu compared to their clas-
158
tic sediments. Interpreting these data Fryer (1983) suggests that there was a significant hydrothermal input to the Proterozoic seas from which these deposits are formed. Our data (Figs. 5 and 6) profusely demonstrate the presence of modest to strong negative Ce anomalies in many samples of CBIF along with positive Eu anomalies and La enrichment. Two SBIF samples from the same locality exhibit positive Ce anomalies similar to some samples of shales interbedded with CBIF and SBIF, (Fig. 6E). A few samples of CBIF (Fig. 5 ) also show weak to moderate positive Ce anomalies. Discussion
The relationship between ZREE and A1203 is very convincing in demonstrating that increases in ZREE are a direct consequence of simultaneous clastic deposition and chemogenic precipitation (Fig. 8C). A mixing of clastic and chemogenic components is illustrated in Figs. 3 and 4. The clastic component was probably contributed in in a volcaniclastic and a terrigenous form. This aspect is demonstrated by the Ni, Cr, Zr, MgO, Ti, V, Hf and other elements (Figs. 3 and 4). High Cr and Zr contents for example in highly aluminous SBIF can be contributed to the two divergent clastic sources. Since the end product is an outcome of three divergent processes absolute linearity between any two distinguishing parameters will not develop, because the third will interfere with the other two. Positive Eu anomalies, La enrichment and depleted patterns of REE indicate that hydrothermal solutions brought these REEs and supplied FeO and SiO2 from deeper sources (see Morey, 1983). The inference that the REE, FeO and SiO2 of the BIF was mainly supplied by hydrothermal solutions emplaced at AMOR is strongly substantiated by Fig. 8D. Here, almost all data points of CBIF fall within the field of hydrothermal deposits, whereas those of SBIF and SH fall outside and away from this field. These hydrothermal solutions were em-
C. MANIKYAMBA ET AL.
placed into the ambient seawater which was largely reducing at that time. There is convincing evidence of production of photosynthetic oxygen through stromatolites in the shallower part of the Sandur basin. In other regions the evidence of stromatolites push the record of life and photosynthesis back as far as 3.4 or 3.5 Ga (Lowe, 1980; Walker et al., 1983). Biogenically produced oxygen reacted with solutions containing REE, FeO and SiO2 and precipitated iron as Fe hydroxides and oxidized Ce to its tetravalent state resulting in negative Ce anomalies as in modern oceans. The Ce thus removed from BIF was accummulated in some shales (Fig. 6E ). This reaction appears to have taken place at the shallow shelf region at the wave base interaction level not very much below the photic zone, where the precipitation of ferrous iron to ferric iron was mediated by biogenic oxidation; simultaneous volcanic activity and terrigenous sedimentation resulted in a ficterogenous product. Thus the observed compositional characters in EREE, La enrichment, Ce depletion, positive Eu anomalies, variation in A1203, TiO2, MgO, Rb, Sr, Cr, Ni, Zr, Hf, Sc, V, Y and other elements appear to be controlled by divergent clastic input to the ambient ocean, its mixing with the hydrothermal solutions and oxidation by photosynthetic oxygen. Combination of these four processes has diluted the effect of each one of them and a hybrid sequence of chert to shales through CBIF and SBIF is developed in a shelf sequence. The hydrothermal activity appears to have taken place in a relatively deeper portion of the basin and was accompanied by intermittent volcanism. Hydrothermal solutions were enriched in dissolved ferrous iron. Due to thermal and chemical gradient and upwelling, these solutions were mixed with ocean water which was reducing except in places where photosynthetic bacteria were producing oxygen. As a result of ocean circulation and thermo-chemical gradients these solutions moved towards the shoreline (see Borahert, 1960) and met oxygen below the wave base or
BIFOFTHEARCHAEANSANDURGREENSTONEBELT
159
TABLE 2 Flow chart of the Sandur belt BIFs RAMORHydrothermal solutions ] Exposed continental nuclei educing andalkalinetoneutralenvironment (depleted S~REE,I IAnaerobic environmental, intense chemical and physical a, Eu enrichment, Higher concentration of SiO2 and FeO) J ~veathering I
dTransport towards shoreline due to thermal and chemical gra~ Dissolved load in river waters suspended load of terrigenous / material (negative europium anomalies, high YREE, A1203, Zr ient and ocean circulation 'and Hf ) Volcanism within the basin, I volcaniclastic input (high Cr and Ni ! e a r stable shelf Biogenically mediated oxida- I ntermittently oxidizing and alternating acidic tion. (negative Ce anomalies,[ nd alkaline environment I presence of stromatolites mi-] [crobiota organic carbon) ] I Si02 precipitation and rhythmic intermittent Fe/03 ontinuous below wave base and photic zone. Fluctuations in [precipitation Iwave base
~
at a fluctuating wave base level, got oxidized and precipitated as Fe203, deposited below the wave base and the photic zone. However, fluctuations in the wave base and probably the photic zone also produced some chert bands in which microfossils could be preserved (Naqvi et al., 1987; Venkatachala et al., 1990). The precipitation of silica in such an enormous quantity, in the absence of any known Archaean biogenic activity, capable of silica deposition remains an enigma, although LaBerge (1986) believes that some form of silica precipitating microorganism existed in that period. Some workers believe that changes in Eh and pH values have caused silica precipitationu (Morey, 1983), whereas others suggest that evaporitic conditions and seasonal changes have caused precipitation of iron and silica (Trendall, 1983; Garrels 1987 ). The clay composition of the shaly beds in the Sandur, Chitradurga and Shimoga belts is mostly kaolinitic, indicating its derivation from a highly weathered source. Iron present i n the continental rocks which gave rise to kaolin must have been brought in solution to the basin and precipitated there with the iron from hydrothermal solutions. Some of the Fe would also
have been provided by leaching of the contemporaneous basic volcanic rocks. This may also explain the variation in the Eu anomalies from strongly positive to even negative in both CBIF and SBIF. We propose that the Sandur belt BIFs have been generated by a combination of factors represented in a flow chart given in Table 2. In this model, sedimentation of BIF occurred on the shallow stable shelf while intermittent volcanism and hydrothermal activity was in progress in the eastern part of the basin probably on an Archaean mid-oceanic ridge.
Conclusions From the above discussion and presentation of geochemical data on the oxide facies BIFs the following conclusions can be drawn: ( l ) Cherty and shaly BIFs are interbedded in the Sandur schist belt. Cherts and shales are two end members in which mixing of minerals containing A1203 and Fe203 in various proportions have given rise the observed compositional variations from cherts to ferruginous cherts, cherty BIF, shaly BIF to shales. These variations caused by mixing of volcaniclastic and terrigenous sediments with chemogenic
160 precipitates is illustrated by Ni, Cr, Zr, Y, V, Sc, Ti, FREE and NdN/YbN ratios. (2) The main source for the iron and silica were hydrothermal solutions generated at AMOR. The hydrothermal sources of iron and REE are reflected by La and Eu enrichment and extremely depleted EREE in CBIF. These patterns are similar to the REE patterns of EPR and Red Sea hydrothermal solutions. Fluvial contribution of FeO and REEs in the form of dissolved load from the land is not completely ruled out. (3) Signatures of the volcaniclastic debris are preserved in the high Cr and Ni content of both the CBIF and SBIF, and terrigenous sedimentation is reflected in Zr, Y, V, Sc, Hf, Ta and Nb abundances. (4) A large number of samples show negative Ce anomalies and indicate oxidation of hydrothermal solutions on stable shelf zones. The magnitude of La enrichment is variable suggesting the mixing of hydrothermal solutions with ambient seawater. ( 5 ) Strong positive Eu anomalies are found in CBIF and weak positive to negative Eu anomalies are found in SBIF indicating the diminishing nature of the hydrothermal signature from CBIF to SBIF. (6) Based on the geochemical data and field evidence it is suggested that the Sandur schist belt BIFs were deposited by a combination of four processes. Hydrothermal solutions and fluvial solutions provided the SiO2 and FeO and REE of the deposits. Hydrothermal solutions were emplaced in a reducing and neutral to alkaline environment near vents at AMOR. These solutions due to the thermal and chemical gradient convected towards the shoreline and got mixed with ambient ocean water and clastic input of volcanic and terrigenous origin. Existence of stromatolites in the underlying carbonates indicates photosynthetic activity which provided oxygen for the oxidation of ferrous iron to ferric iron and C e 3+ to C e 4+ and thus biogenically mediated precipitation of iron took place along with terrigenous and vol-
c. MANIKYAMBAETAL. caniclastic varved type beds. Near the shoreline the environment was acidic and alkaline with alternating precipitation of SiO2 and Fe203 in response to changes in Eh and pH. The banding of the BIF represents the break in precipitation of iron due to the nonavailability of photosynthetic 02 or hydrothermal FeO or both. Therefore, four processes, namely (1) hydrothermal activity in a deeper reducing alkaline environment and its movements towards the shoreline, (2) terrigenous sedimentation, (3) volcanic sedimentation, and (4) oxidation from biogenically produced oxygen, have produced the Sandur schist belt BIFs.
Acknowledgements We express our profound gratitude to Prof. D. Gupta Sarma FNA, Director NGRI, for his encouragement, keen involvement, support and permission for publication of this work. The work has been supported under the DST Grant No. SP 12/PC2/86 and NGRI project on Lithos 5.1. Mr. T. Gnaneshwar Rao, Mr. S.L. Ramesh and Mr. K.V. Anjaiah have provided help in the analysis of the samples. Mrs. Nancy Rajan has prepared the manuscript. We are deeply indebted to Dr. B.P. Radhakrishna, Dr. M. Ramakrishnan, Dr. R. Srinivasan, Dr. V. Divakara Rao, Dr. B.L. Narayana for a critical review of the manuscript and their suggestions for the improvement. We express our gratitude to Dr. G.L. Reddy, Mr. K.K. Reddy, M. M.H. Shareef, Mr. Masood Rizvi of NMDC, Donimalai iron ore mines who have greatly helped in sampling and provided core samples from the core library of NMDC. Dr. A.K. Chaterjee has extended all help and facilities at the Donimalai mine area. An earlier version of the manuscript has been greatly improved by extremely valuable review, comments and suggestions from Prof. H.L. James and A.M. Goodwin. We are extremely grateful to both of them.
BIF OF THE ARCHAEAN SANDUR GREENSTONE BELT
References Addy, S.K., 1979. Rare earth element patterns in manganese nodules and micronodules from northwest Atlantic. Geochim. Cosmochim. Acta, 43:1105-1115. Anhaeusser, C.R., 1990. Precambrian crustal evolution and metallogeny of South Africa. In: S.M. Naqvi (Editor), Precambrian Continental Crust and its Economic Resources. Elsevier, Amsterdam, pp. 123-156. Balaram, V., Manikyamba, C., Ramesh, S.L. and Anjaiah, K.V., 1991. REE and trace elements analysis of iron formation refrance samples by inductively coupled plasma mass spectrometry. Proc. 5th Natl. Symp. on Mass Spectrometry, Ahmedabad EPS, pp. 1-3. Barrett, T.J. and Jarvis, I., 1988. Rare-earth elements geochemistry of metalliferous sediments from DSDP Reg 92: the East Pacific Rise transect. Chem. Geol., 67: 243-259. Barrett, T.J., Fralick, P.W. and Jarvis. I., 1988. Rare-earth element geochemistry of some Archaean iron formalions north of Lake Superior, Ontario. Can. J. Earth Sci., 25: 570-580. Beukes, N.J., 1980. Suggestion towards and classification of and nomenclature for iron formation. Trans. Geol. Soc. S. Aft., 83: 285-290. Beukes, N.J., 1983. Paleoenvironmental setting of iron formations in the depositional basin of the Transvaal Supergroup, South Africa. In: A.F. Trendall and R.C. Morris (Editors), Iron Formations: Facts and Problems. Elsevier, Amsterdam, pp. 131-209. Beukes, N.J. and Klein, C., 1990. Geochemistry and sedimentology of a facies transition from microbanded to granular iron formation in the Early Proterozoic Transvaal Supergroup, South Africa. Precambrian Res., 47: 99-139. Beukes, N.J., Klein, C., Kaufman, A.J. and Hayes, J.M., 1990. Carbonate petrography, kerogen distribution and carbon and oxygen isotope variations in an early Proterozoic Transition from limestone to iron formation deposition, Transvaal Supergroup, South Africa. Econ. Geol., 85: 663-690. Bhaskar Rao, Y.J., Sivaraman, T.V., Pantulu, G.V.C., Gopalan, K. and Naqvi, S.M., 1992. Rb-Sr ages of Late Archaean metavolcanics and granites, Dharwar craton, evidence for Early Proterozoic thermotectonic event (s). Precambrian Res., submitted. Borahert, H., 1960. Genesis of marine sedimentary iron ore. Trans. Inst. Min. Metall., 69: 261-277. Campbell, A.C. and 10 co-authors, 1988. Chemistry of hot springs on the Mid-Atlantic Ridge. Nature, 335:514519. Chakraborty, K.L. and Majumdar, T., 1986. Geological aspects of the banded iron formation of Bihar and Orissa. J. Geol. Soc. India, 28: 71-91. Cloud, P.E., 1973. Palaeobiological significance of the banded iron formation. Econ. Geol., 68:1135-1143.
161
Condie, K.C., 1981. Archaean Greenstone Belts. Amsterdam, Elsevier, 434 pp. Davy, R., 1983. A contribution on the chemical composition of Precambrian iron formations. In: A.F. Trendall and R.C. Morris (Editors), Iron Formations: Facts and Problems. Elsevier, Amsterdam, pp. 325-343. De Baar, HA.W, Bacon, M.P. and Brewer, P.G., 1985. Rare earth elements in the Pacific and Atlantic Oceans. Geochim. Cosmochim. Acta, 49: 1943-1959. Derry, L.A. and Jacobsen, S.B., 1990. The chemical evolution of Precambrian sea water: Evidence from REEs in banded iron formations. Geochim. Cosmochim. Acta, 54: 2965-2977. Devaraju, T.C. and Laajoki, K., 1986. Mineralogy and mineral chemistry of the manganese poor and manganiferous iron formations from the high grade metamorphic terrain of southern Karnataka, India. J. Geol. Soc. India, 28: 134-164. Drever, J.I., 1974. Geochemical model for the origin of Precambrian banded iron formations. Geol. Soc. Am. Bull., 85:1099-1106. Dimroth, E., 1986. Depositional environments and tectonic settings of cherty iron formations of the Canadian Shield. J. Geol. Soc. India, 28: 239-250. Dymek, R.F. and Klein, C., 1988. Chemistry, petrology and origin of banded iron formation lithologies from the 3800 Ma Isua supracrustal belt, West Greenland. Precambrian Res., 39: 247-302. Drury, S.A., Holt, R.N., Vanclastern, P.C. and Beckinsale, R.D., 1983. Sm-Nd and Rb-Sr ages of Archaean rocks in western Karnataka, South India. J. Geol. Soc. India, 24: 454-459. Eugster, H.P. and Chou, I.M., 1973. Depositional environments of Precambrian banded iron formations. Econ. Geol., 68:1144-1168. Ewers, W.E. and Morris, R.C., 1981. Studies on the Dales Gorge Member of the Brockman Iron Formation, Western Australia. Econ. Geol., 77:1929-1953. Flynn, R.T. and Burnham, C.W., 1978. An experimental determination of rare earth partitioning coefficients between a chloride containing vapour phase and silicate melts. Geochim. Cosmochim. Acta, 42: 685-701. Fryer, B.J., Fyfe, W.S. and Kerrich, R., 1979. Archaean volcanogenic oceans. Chem. Geol., 24: 25-33. Fryer, B.J., 1983. Rare earth elements in iron formations. In: A.F. Trendall and R.C. Morris (Editors), Iron Formation: Facts and Problems. Elsevier, Amsterdam, pp. 345-358. Garrels, R.M., 1987. A mode for the deposition of the microbanded Precambrian iron formations. Am. J. Sci., 287: 81-106. Gole, M.J. and Klein, C., 1981. Banded iron formation through much of time. J. Geol., 89:169-183. Goodwin, A.M., 1973. Archaean iron-formations and tectonic basins of Canadian Shield. Econ. Geol., 68:915933.
162 Goodwin, A.M., Thode, H.G., Chou, C.L. and Karkhansis, S.N., 1985. Chemostratigraphy and origin of the late Archaean siderite-pyrite rich Helene Iron Formation, Michipicoten belt, Canada. Can. J. Earth Sci., 22: 72-84. Goldstein, S.J. and Jacobsen, S.B., 1988a. Nd and Sr isotopic systematics of river water suspended material: implications for crustal evolution. Earth Planet. Sci. Lett., 87: 249-265. Goldstein, S.J. and Jacobsen, S.B., 1988b. REE in the great Whale river Estuary, NW Quebec. Earth Planet. Sci. Lett., 88: 241-252. Govil, P.K., 1985. X-ray fluorescence analysis of major, minor and selected trace elements in new IWG refrance rock samples. J. Geol. Soc. India, 26: 38-42. Graf, J.L., 1978. Rare earth elements, iron formations and sea water. Geochim. Cosmochim. Acta, 42:1845-1850. Gross, G.A., 1980. A classification of iron formation based on depositional environment. Can. Mineral., 18:215222. Hargraves, R.B., 1986. The metallogenetic significance of iron formation and related stratifer rocks. J. Geol. Soc. India, 28: 92-108. Hofmann, H.J., Saxe, R.P. and Herdusco, E.N., 1991. Archaean stromatolites in Michipictoen group siderite ore at Wawa, Ontario. Econ. Geol., 86: 1023-1030. Holland, D.H., 1973. The oceans: a possible source of iron in iron formations. Econ. Geol., 68:1169-1172. James, H.L., 1983. Distribution of banded iron-formation in space and time. In: A.F. Trendall and R.C. Morris (Editors), Iron Formation: Facts and Problems. Elsevier, Amsterdam, pp. 471-490. Kasting, J.F., 1987. Theoretical constraints on oxygen and carbondioxide concentrations in the Precambrian atmosphere. Precambrian Res., 34: 205-229. Kerrich, R. and Fryer, B.J., 1979. Archaean precious metal hydrothermal systems Dome mine, Abitibi greenstone belt: II REE and oxygen isotope relations. Can. J. Earth Sci., 16: 440-458. Khan, R.M.K., Govil, P.K. and Naqvi, S.M., 1992. Geochemistry and genesis of BIF from Kudremukh schist belt, Karnataka Nucleus, India. J. Geol. Soc. India, submitted. Klein, C. and Beukes, C., 1989. Geochemistry and sedimentology of facies transition from limestone to iron formation deposition in the early Proterozoic Transvaal Supergroup, South Africa. Econ. Geol., 84:17331773. Klinkhammer, G., Elderfield, H. and Hudson, A., 1983. Rare earth elements in seawater near hydrothermal vents. Nature, 305: 185-188. La Berge, G.L., 1986. A model for the biological precipitation of Precambrian iron formation. LPI Conference on Greenstone belts (Abstract) LPI Rep., Houston, TX, pp. 71-75. Litherland, M. and Malan, S.P., 1973. Manganiferous
c. MANIKYAMBAET AL. stromatolites from the Precambrian of Botswana. J. Geol. Soc. London, 129: 543-544. Lupton, J.E. and Craig, H., 1981. A major Helium-3 source at 15 °S on East Pacific Rise. Science, 214:13-18. Lowe, D.R., 1980. Stromatolites, 3400 Myr old from the Archaean of Western Australia. Nature, 284: 441-443, Majumdar, T., Chakraborty, K.L. and Bhattacharyya, 1982. Geochemistry of banded iron formation of Orissa, India. Miner. Deposit, 17:107-118. Manikyamba, C., 1992. Geochemistry of iron and manganese formations of Sandur schist belt, Dharwar craton, India. Ph.D. thesis, Osmania Univ., 327 pp. Michard, A. and Albarede, F., 1986. The REE content of some hydrothermal fluids. Chem. Geol., 55:51-60. Michard, A., Albarede, F., Michard, G., Minster, J.F. and Charlan, J.L., 1983. Rare earth elements and uranium in high temperature solutions from the East Pacific Rise hydrothermal vent fields. Nature, 314: 325-330. Morey, G.B., 1983. Animikie basin, Lake Superior region, USA. In: A.F. Trendall and R.C. Morris (Editors), Iron Formation: Facts and Problems. Elsevier, Amsterdam, pp. 13-68. Mukhopadhyay, D. and Matin, A., 1990. A new interpretation of the structural pattern in the Sandur schist belt in the Dharwar craton, Karnataka (Abstract). Seminar on Evolution of Precambrian crust in India, Calcutta, p. 33. Murthy, P.S.N. and Reddy, K.K., 1984. 2900 million year old stromatolite from Sandur greenstone belt of Karnataka Craton, India. J. Geol. Soc. India, 25: 263-266. Naha, K., Srinivasan, R. and Naqvi, S.M., 1986. Structural unity in the Early Precambrian Dharwar tectonic province, Peninsular India. Q. J. Geol. Min. Metal. Soc. India, 58: 219-243. Naqvi, S.M., 1981. The oldest supracrustal of Dharwar craton. J. Geol. Soc. India, 22: 458-469. Naqvi, S.M. and Rogers, J.J.W., 1983. Precambrian of South India. Mem. Geol. Soc. India, 4: 1-575. Naqvi, S.M. and Rogers, J.J.W., 1987. Precambrian Geology of India. Oxford Univ. Press, New York, 223 pp. Naqvi, S.M., Venkatachala, B.S., Manoj Shukla, Kumar, B., Natarajan, R. and Mukund Sharma, 1987. Silicifled cyanobacteria from the cherts ofArchaean Sandur schist belt, Karnataka, India. J. Geol. Soc. India, 29: 535-539. Naqvi, S.M., Sawkar, R.H., Subba Rao, D.V., Govil, P.K. and Gnaneshwar Rao, T., 1988. Geology and geochemistry and tectonic setting of Archaean greywackes from Karnataka nucleus, India. Precambrian Res., 39: 193-216. Naqvi, S.M., Manikyamba, C. and Gnaneshwar Rao, T., 1992. Geology of transition from stromatolitic carbonates to banded iron and manganese formations in the stable shelf zone of the Archaean Greenstone belts of Karnataka Nucleus, India. submitted. Nealson, K.H. and Myers, C.R., 1990. Iron reduction by
163
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bacteria: A potential role in the genesis of banded iron formations. Am. J. Sci., 290A: 35-45. Oehler, D.Z., Shopf, J.W. and Kvenvolden, K.A., 1972. Carbon isotope studies of organic matter in Precambrian rocks. Science, 175:1246-1248. Perry, E.C., Jr., 1983. Oxygen isotope geochemistry of iron formations. In: A.F. Trendall and R.C. Morris (Editors), Iron Formation: Facts and Problems, Elsevier, Amsterdam, pp. 359-371. Radhakrishna, B.P., 1983. Archaean granite greenstone terrain of South India. In: S.M. Naqvi and J.J.W. Rogers (Editors), Precambrian of South India. Mem. Geol. Soc. India, 4: 1-46. Radhakrishna, B.P. and Naqvi, S.M., 1986. Precambrian continental crust of India and its evolution. J. Geol., 94: 145-166. Radhakrishna, B.P., Devaraju, T.C. and Mahabaleshwar, 1986. Banded iron formations of India. J. Geol. Soc. India, 28: 1-17. Roy, A. and Biswas, S.K., 1979. Metamorphic history of the Sandur schist belt, Karnataka. J. Geol. Soc. India, 20: 179-187. Roy, A. and Biswas, S.K., 1983. Stratigraphy and structure of Sandur schist belt, Karnataka. J. Geol. Soc. India, 24: 19-28. Rao, T.G., 1992. Geology and geochemistry of BIFs of the central part of the Chitradurga schist belt, Karnataka, India. Unpublished Ph.D. thesis, OU, Hyderabad, 350 pp., unpubl. Ruhlin, D.E. and Owen, R.M., 1986. The rare earth element geochemistry ofhydrothermal sediments from the East Pacific Rise: examination of a seawater scavenging mechanism. Geochim. Cosmochim. Acta, 50: 393400. Schidlowski, M., 1990. Early evolution of life and economic mineral and hydrocarbon resources. In: S.M. Naqvi (Editors), Precambrian Continental Crust and its Economic Resources. Elsevier, Amsterdam, pp. 605-630. Sholkovitz, E.R., Piepgras, D.J. and Jacobsen, S.B., 1989. The pore water chemistry of rare earth elements in Buzzards Bay sediments. Geochim. Cosmochim. Acta, 53: 2847-2856. Simonson, B.M., 1982. Sedimentology of Precambrian iron formation with special reference to the Sokoman Formation and associated northeastern Canada. Ph.D dissert., John Hopkins Univ., Baltimore, MD, 359 pp. Simonson, B.M., 1985. Sedimentological constraints on the origins of Precambrian iron formations. Geol. Soc. Am. Bull., 96: 244-252. Srinivasan, R. and Sreenivas, B.L., 1972. Dharwar stratigraphy. J. Geol. Soc. India, 13: 72-83. Srinivasan, R., Manoj Shukla, Naqvi, S.M., Yadav, V.K., Venkatachala, B.S., Uday Raj, B. and Subba Rao, D.V., 1989. Archaean stromatolities from Chitradurga schist belt, Dharwar craton, South India. Precambrian Res., 43: 239-250.
Srinivasan, R., Naqvi, S.M. and Vasantha Kumar, B., 1990. Archaean shelf-facies and stromatolite proliferation in Dharwar Supergroup, North Kanara District, Karnataka, J. Geol. Soc. India, 35:203-212. Sverjensky, D.A., 1984. Europium redox equilibria in aqueous solution. Earth Planet. Sci. Lett. 67: 70-78. Swami Nath, J. and Ramakrishnan, M., 1981. Present classification and correlation. In: J. Swami Nath and M. Ramakrishnan (Editors), Early Precambrian Supracrustals of Southern Karnataka. Mem. Geol. Surv. India, 112: 23-38. Taylor, P.N., Moorbath, S., Chadwick, B., Ramakrishnan, M. and Viswanatha, M.N., 1984. Petrography, chemistry and isotopic ages of Peninsular gneiss, Dharwar acid volcanic rocks and the Chitradurga granite with special reference to the Late Archaean evolution of Karnataka craton, Southern India. Precambrian Res., 23: 349-375. Trendall, A.F., 1973. Iron formations of the Hamersley Group of Western Australia: type examples of varved Precambrian evaporites. In: Earth Sci. (UNESCO, Paris), 9: 257-268. Trendall, A.F., 1983. The Hamersley Basin. In: A.F. Trendall and R.C. Morris (Editors) Iron Formation: Facts and Problems, Elsevier, Amsterdam, pp. 69-130. Trendall, A.F. and Morris, R.C. (Editors), 1983. Iron Formation: Facts and Problems. Elsevier, Amsterdam, 558 pp. Venkatachala, B.S., Shukla, M., Sharma, M., Naqvi, S.M., Srinivasan, R. and Uday Raj, B., 1990. Archaean microbiota from the Donamalai Formation, Dharwar Supergroup, India. Precambrian Res., 47: 27-34. Venkatachala, B.S., Naqvi, S.M., Chadha, M.S. and 13 coauthors, 1991. Geology and geochemistry and paleobiology of Precambrian stromatolites and associated sedimentary rocks from the Dharwar craton: constraints on Archaean biogenic processes. Proc. IndoSoviet Symp. on Stromatolites and Stromatolitic Deposits. Wadia Inst. of Himalayan Geology, Dehradun, pp. 1-20. Walker, J.C.G., Klein, C., Schidlowski, M., Schopf, J.W., Stevenson, D.J. and Walter, M.R., 1983. Environmental evolution of the Archaean-Early Proterozoic Earth. In: J.W. Schopf (Editor), Earth's Earliest Biosphere Its Origin and Evolution. Princeton Univ. Press, Princeton, pp. 260-290. Walter, M.R. and Hofmann, H.J., 1983. The palaeobiology and palaeoecology of Precambrian iron formations. In: A.F. Trendall and R.C. Morris (Editors), Iron Formation: Facts and Problems. Elsevier, Amsterdam, pp. 373-400. Weiss, R.F., 1977. Hydrothermal manganese in the deep sea: scavenging residence time and M n / H e relationships. Earth Planet. Sci. Lett., 37: 257-262. Wilson, J.F. and Nutt, T.H.C., 1990. The nature and occurrence of mineralization in the Early Precambrian -
-
164 Crust of Zimbabwe. In: S.M. Naqvi (Editor), Precambrian Continental Crust and its Economic Resources. Elsevier, Amsterdam, pp. 555-592. Zhai, M. and Windley, B.F., 1990. The Archaean and early
C. MANIKYAMBA ET AL.
Proterozoic banded iron formations of North China: their characteristics, geotectonic relations, chemistry and implications for crustal growth. Precambrian Res., 48: 267-286.