Petrological and stable isotope studies of carbonate and sulfide minerals from the Gunflint Formation, Ontario: evidence for the origin of early Proterozoic iron-formation

Precambrian Research, 52 ( 1991 ) 347-380

347

Elsevier Science Publishers B.V., Amsterdam

Petrological and stable isotope studies of carbonate and sulfide minerals from the Gunflint Formation, Ontario: evidence for the origin of early Proterozoic iron-formation William J. Carrigan a,m and Eion M. Cameron a,b a Derry Laboratory, Ottawa-Carleton Geoscience Center, Ottawa, Ontario, and Department of Geology, University of Ottawa, Ottawa, Ont. K IN 6N5, Canada b Geological Survey of Canada, 601 Booth St., Ottawa, Ont. K l A 0ES, Canada (Received August 21, 1990; revised and accepted March 26, 1991 )

ABSTRACT Carrigan W.J. and Cameron, E.M., 1991. Petrological and stable isotope studies of carbonate and sulfide minerals from the Gunflint Formation, Ontario: evidence for the origin of early Proterozoic iron-formation. Precambrian Res., 52: 347-380. The ~ 1.9 Ga Gunflint Formation, in the Thunder Bay district of Ontario, principally comprises iron-formation, overlain by an iron-poor limestone/doiostone member (the Upper Limestone Member). Rocks are virtually unmetamorphosed, retaining primary textures, except locally around diabase sills. The Gunflint presents an opportunity to study both unmetamorphosed and metamorphosed iron-formation and to study the changes from iron-rich to iron-poor sedimentation. Of the carbonate minerals in iron-rich units, siderite formed at or near the sediment/water interface, with ankerite and calcite forming later during diagenesis and during metamorphism around sills. Heaviest ~3C values of siderite, and of dolomite in the Upper Limestone Member, are consistently near 0%0 throughout the stratigraphic section, indicating that marine bicarbonate was the source of carbon. A spread of values between 0 and - 7%0 for unmetamorphosed rocks is the result of the incorporation of oxidized organic matter during diagenesis, whereas ~ aC values more negative than -7o/oo are the result of metamorphism. The heaviest ~ 8 0 values from unmetamorphosed rocks, -5.30/00 (PDB) for siderite, - 6. 1%o for ankerite, and - 6.7%o for calcite, are considered to represent the original marine composition. Lighter values (to - 17%o) represent isotopic exchange reactions with pore fluids at higher temperatures and/or isotopic exchange with 'SO-depleted meteoric water. Low S/C ratios and a relatively narrow range of positive ~34S values in pyrite throughout most of the formation are consistent with bacterial sulphate reduction of seawater containing sulphate at a significantly lower concentration than the modern ocean. Locally, higher sulfide contents and a wider range of ~34S values were the product of the introduction of sulfur after deposition. It is proposed that the Gunflint Formation was deposited in a stratified water column with anoxic bottom water. The depositional basin had restricted communication with the open ocean and was affected by distal volcanism. Hydrothermal activity associated with the volcanism provided a large source of dissolved iron, and possibly silica, which helped buffer 02 and sulphate in water to low levels. Low concentrations of sulphate limited the generation of H2S that would otherwise have restricted the solubility of iron. During periods of increased hydrothermal activity, the anoxic/oxic water boundary moved upwards, permitting transport of iron to the shallow shelf, where it was precipitated as siderite, iron hydroxides, iron-silicates, or pyrite depending on physico-chemical conditions. The transition to the overlying limestone marks a decrease in hydrothermal activity with a contraction of the redox boundary.

Introduction

Many aspects of Precambrian iron-forma~Present address: P.O. Box 62, Saudi Aramco, Dhahran 31311, Saudi Arabia.

0301-9268/91/$03.50

tions remain controversial. Among these are the composition of the primary sediment and the source of the chemical components for ironformation. The present mineralogy of iron-formations is the result of diagenetic and meta-

© 1991 Elsevier Science Publishers B.V. All rights reserved.

348

morphic modification of the primary sediment (e.g. Klein, 1983), so petrographic studies alone cannot always determine the original depositional phases. Stable isotope ratios of carbon and oxygen in carbonate minerals and sulfur in pyrite have often been used as indicators of depositional and diagenetic environments. However, compared to other sedimentary rocks, such as limestones, relatively few isotope studies have been carried out in iron-formations (e.g. Becker and Clayton, 1972, 1976; Perry et al., 1973; Goodwin et al., 1976, 1985; Schidlowski et al., 1979; Thode and Goodwin, 1983; Cameron, 1983a; Baur et al., 1985; Beukes et al., 1990). Of these studies, few have been carried out in conjunction with detailed petrography. The Gunflint Formation is a virtually unmetamorphosed Precambrian iron-formation in which primary textures have been well preserved, but which is locally metamorphosed near diabase sills, thus permitting the study of both unmetamorphosed and metamorphosed iron-formation. The top of the Gunflint Formation is an iron-poor limestone/dolostone member, allowing a comparison to be made between the deposition of iron-rich and ironpoor sediments. In this study we report the resuits of a study of the petrography and stable isotope geochemistry (~13C, 8180, and ~34S) of carbonate and sulfide minerals from the Gunflint Formation.

Geologic setting Early Proterozoic rocks of the western Lake Superior region in Ontario and Minnesota comprise the Mille Lacs and Animikie groups; correlative rocks in Michigan and Wisconsin are the Marquette Range Supergroup (MRS) (Morey, 1983). Major iron-formations, including the Gunflint Formation, are contained in the Animikie Group and correlative strata in the MRS. The tectonic evolution of the early Proterozoic rocks remains uncertain. Models involving sedimentation on a rifted passive

W.J. CARRIGAN AND E.M. CAMERON

margin during the opening of an ocean basin and subsequent deformation during the Penokean Orogeny have been proposed by Cambray (1978) and Larue and Sloss (1980). More recently it has been proposed that the rocks consist of a rifted passive margin sequence (Mille Lacs Group and correlative strata), overstepped northward by a synorogenic foredeep sequence (Animikie Group and correlative strata) (Hoffman, 1987, 1989; Southwick et al., 1988; Barovich et al., 1989). The Penokean Orogeny may have occurred during accretion of island arcs (Sims et al., 1989) and/or back-arc basins (Southwick et al., 1988 ) to the continental margin prism during closure of an ocean basin. Evidence for extensive magmatism within the Animikie Basin includes widespread volcanic ash in the Gunflint Formation (Goodwin, 1956) and mafic volcanic rocks of the Emperor Volcanic Complex, which underlies and is interlayered with the Ironwood Iron-Formation (Greenberg and Brown, 1983). The ~ 1.9 Ga (Morey, 1983) Gunflint Formation, located in the Thunder Bay district of northwestern Ontario (Fig. 1 ), is a Lake Superior type iron-formation (Gross, 1965 ). It is about 120 m in thickness and is conformably overlain by the Rove Formation, ~ 1000 m of dark grey to black argillite and interbedded argillaceous siltstone and greywacke (Morey, 1967). Together they comprise the Animikie Group in Ontario, which unconformably overlies Archean basement and is in turn unconformably overlain by the middle-Proterozoic Sibley Group. The Mille Lacs Group, which underlies the Atlimikie Group in parts of Minnesota, is not preserved in Ontario. To the southwest, the Gunflint and Rove formations are cut by intrusive rocks of the Duluth Complex. The Animikie Group in Ontario is essentially undeformed and has undergone subgreenschist metamorphism, except for thermal effects associated with the late Proterozoic Duluth Complex and associated thin diabase sills (Floran and Papike, 1975, 1978 ).

PETROLOGICAL AND STABLE ISOTOPE STUDIES OF CARBONATE AND SULFIDE MINERALS, GUNFLINT FORMATION

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Fig. 1. Geological map of the Thunder Bay region. Inset shows the distribution of the major early Proterozoic ironformations in the western Lake Superior region:GF= Gunflint Formation; BI= BiwabikIron Formation; TR = Trommald Formation;I W = IronwoodIron Formation; NG= NegauneeIron Formation;R V= Riverton Iron Formation; VU= Vulcan Iron Formation. A m a x i m u m burial temperature of 100 to 130°C has been estimated for the Gunflint by Miyano ( 1987 ). In this paper, only rocks that have undergone contact metamorphism will be referred to as metamorphosed. Goodwin (1956) determined that the Gunflint Formation was, in part, cyclically deposited. He defined four members: the Basal Conglomerate Member, overlain by the iron-rich Lower and Upper Members, in turn is overlain by the iron-poor Upper Limestone Member. The Conglomerate Member is a thin ( < 3 m ) , discontinuous polymictic conglomerate containing pebble- to boulder-size clasts derived from underlying Archean rocks. Goodwin (1956) divided each of the Lower and Upper Members into sub-members comprised of four facies, in ascending order: algal chert, tuffaceous shale, and taconite facies; a banded

chert-carbonate facies is laterally gradational with the taconite facies. However, Shegelski (1982) suggests that the sub-member divisions have become less useful as detailed stratigraphic investigations have uncovered additional units. In this paper, five lithofacies are recognized: black shale (tuffaceous shale of Goodwin), arenite (taconite of Goodwin), stromatolite (algal chert of Goodwin), banded chert-carbonate and laminated carbonate. Laminated carbonate is a newly defined lithofacies that was previously included with banded chert-carbonate, but contains very little chert. The Upper Member is gradationally overlain by limestone/dolostone of the Upper Limestone Member. The stromatolite facies is characterized by biohermal mounds, up to 1.5 m in height and 2 m in diameter, containing millimeter- to cen-

350

W.J. C A R R I G A N A N D E.M. C A M E R O N

(Dimroth, 1976; Markun and Randazzo, 1980; Lougheed, 1983). Bioherms are surrounded and overlain by arenite consisting of medium to coarse, sand-size, oolitic intraclasts composed mainly of chert with variable proportions of hematite, magnetite, greenalite, minnesotaite, and carbonate (Floran and Papike, 1975; Simonson, 1987). The arenite facies consists of cross-stratified, lenticular to wavy beds of medium to coarse, sand-sized, well-

timeter-sized planar, columnar, and domal structures (Hofmann, 1969). The stromatolites are mostly chert, but have laminae rich in iron-silicates, hematite, magnetite, and/or organic carbon. The chert is considered by some workers to have crystallized from a primary silica precipitate (Hofmann, 1969; Gross, 1972; Simonson, 1987; Simonson and Lanier, 1987); others, however, have proposed that silica replaced a primary carbonate rock

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Fig. 2. Photographs of polished slabs of carbonate-rich rocks. A. Laminated carbonate composed of alternating sideriterich (light colored) (see also Fig. 6, B and C) and elastic-rich (dark colored) laminae. B. Arenite composed of medium to coarse sand-size, iron-silicate and hematite-rich intraclasts, containing interbedded layers (S) and rip-up clasts (arrows) of fine-grained, laminated siderite. Siderite does not contain any relic outlines of intraclasts indicating that it did not replace intraclasts but was primary. C. Banded chert-carbonate composed of alternating fine-grained siderite (light colored ) (see also Fig. 5A) and laminated chert (dark colored) (see also Fig. 7C) bands. D. Arenite containing interbedded laminae and thin layers of fine-grained ankerite. Grain outlines are still visible in arenite (light colored) but have been extensively replaced by coarse-grained ankerite (see Fig. 7B). Dark laminae and layers are composed of very finegrained ankerite, which are interpreted as primary. Scale is in inches for all photos.

PETROLOGICAL AND STABLE ISOTOPE STUDIES OF CARBONATE AND SULFIDE MINERALS, GUNFLINT FORMATION

sorted and well-rounded intraclasts. The intraclasts are composed mostly of chert, with variable proportions of greenalite, hematite, minnesotaite, magnetite and carbonates (Floran and Papike, 1975). Many of these grains, especially those associated with stromatolites, are oolitic. The interstitial areas mostly contain early-formed chert cements, which are described in detail by Simonson ( 1987 ). Arenite commonly contains interbeds of shale or carbonate (Fig. 2, B and D). Banded chert-carbonate consists of alternating layers, which range from one to a few centimeters in thickness, of brown siderite and dark grey chert (Fig. 2C). Banding is parallel and may be laterally continuous for tens of meters, but more commonly layers pinch out or are erosionally truncated by overlying layers. Soft sediment deformation structures are common, which include intraformational breccias as well as slumped and folded beds sandwiched between undeformed beds (Gross, 1972). Laminated carbonate is made up of laminae to thin beds of silt- to sand-size clastic detritus containing interstitial carbonate altern/lting with carbonate-rich layers (Fig. 2A). Clastic grains consist of detrital quartz and feldspar, intraclasts similar in composition to those in the arenite facies, and volcaniclastic grains that may be epiclastic in origin. Carbonate-rich layers are composed of siderite, ankerire and fine-grained matrix. Thin beds of fiat pebble conglomerate are also present. Carbonaceous black shale is made up of laminae and very thin cross-stratified or graded beds of clayto silt-size clastic detritus. Clastic grains are probably of pyroclastic rather than terrigenous origin. Pyroclastic interbeds, composed of ash to lapilli-size detritus of basaltic to rhyolitic composition (Hassler and Simonson, 1989), are locally abundant, as are beds rich in carbonate. The Upper Limestone Member marks an abrupt decrease in the amount of chemically precipitated iron. It consists of impure beds of limestone/dolostone and contains variable proportions of pyroclastic detritus,

3 51

iron-formation intraclasts, and silt- to fine sand-size detrital quartz grains. Layers and lenses of chert are common. The Gunflint Formation was probably deposited on a shallow marine platform or shelf, similar to modern shelf carbonate environments (Shegelski, 1982; Simonson, 1985). Water depths during deposition are uncertain. The basal conglomerate represents an initial transgressive lag deposit (Goodwin, 1956 ) and the stromatolites that formed on the conglomerate, probably were deposited in the intertidal zone (Shegelski, 1982). Simonson (1985) proposes water depths of tens of meters for the iron-formation arenites. The fine-grained, laminated sediments may have been deposited in shallow depressions on the shelf or in deeper water during rises in sea level. The overlying Rove Formation represents the final drowning of the shelf.

Petrography The petrography and mineralogy of ironbearing minerals of both unmetamorphosed and metamorphosed rocks of the Gunflint Formation have previously been reported by Floran and Papike (1975, 1978). The petrology of silica cements in arenites of the Gunsyngenesis

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Fig. 3. Paragenetic sequence of carbonate and sulfide minerals from the Gunflint Formation.

352

W.J. CARRIGAN AND E.M. CAMERON

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flint is reported by Simonson (1987) and Simonson and Lanier (1987). Here, we give details on the petrography of carbonate and sulfide minerals in order that constraints can be placed on the interpretation of their isotopic data. The paragenetic sequence of the carbonate and sulfide minerals is shown in Fig. 3 and is discussed below.

Carbonate minerals Carbonate minerals in the Gunflint Formation are siderite, ankerite, dolomite, and calcite; but all four phases are never found together. The assemblages siderite-ankeritecalcite, siderite-ankerite, and ankerite-calcite occur in iron-rich rocks, whereas dolomite and calcite are found in the Upper Limestone Member. Microprobe analyses of representative carbonate minerals are shown on a

(FeMn)CO3-MgCO3-CaCO3 ternary plot (Fig. 4). Siderite and calcite form distinct fields, but ankerite and dolomite form a series from almost pure dolomite to ankerite containing about 30 mole% Fe. Mn is usually less than 2 mole% in siderite and ankerite. Siderite is the earliest-formed carbonate mineral in the iron-rich rocks and was followed paragenetically by ankerite and then calcite. In the Upper Limestone Member, dolomite is the oldest preserved carbonate mineral followed by calcite. Siderite. The abundance of siderite in the Gunflint Formation varies considerably between facies. Beds of essentially pure siderite alternate with beds of chert in the banded chert-carbonate facies. The amount of siderite in the black shale and laminated carbonate facies ranges from siderite-rich beds to beds containing only minor amounts of disseminated grains. Siderite is only a minor component in the arenite and stromatolite facies and is absent in the Upper Limestone Member. Siderite layers in banded chert-carbonate are composed of microspherical siderite grains, ~ 30 a m in diameter (Fig. 5A). The microspheres typically have a 10 to 20 a m thick double outer wall, composed of aggregates of microscopic siderite crystals, around a darker colored central core that is filled by fine crystalline silica. The core may contain aggregates of siderite crystals and rare inclusions of pyrite (see figs. 11-14 of LaBerge, 1973 and figs. 2 and 3 of Kazmierczak, 1979). The density of spheres within layers ranges from tightly packed with very little interstitial chert to loosely packed

Fig. 5. Micro-photographs of primary textures of siderite (all plane-polarized light). A. Siderite layer from banded chertcarbonate composed of microspheres, which consist of a dark inner core of silica surrounded by a lighter colored shell of siderite. Interstitial area between microspheres contains minor amounts of chert. Bar is 0.1 mm. B. Close-up of carbonate lithoclast from black shale. Interior of lithoclast consists of aggregate of dark siderite microspheres, similar to microspheres above, that probably formed at or near the sediment/water interface before being ripped up and redeposited. Lithoclast is partially coated by concentrically laminated, light colored siderite, probably after resedimentation. Bar is 0.2 mm. C. Microphotograph of siderite microspheres from black shale. Microspheres consist of a dark inner core of siderite and a concentrically laminated coating of light-colored siderite. One grain (left center) contains a double inner core surrounded by a single coating, implying that coatings probably formed after sedimentation of dark siderite (inner cores). Pyrite (black) partially replaces siderite and matrix. Bar is 0. l mm.

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with abundant interstitial chert. The spheres are dispersed within layers as individual spheres and less commonly as clusters of tightly adhering spheres. In the latter case, the interiors of some adjacent spheres are connected and the cluster is covered by a common coating of siderite. Individual and small clusters of siderite spheres are common in adjacent chert layers. Rhombic overgrowths of siderite are often observed around siderite spheres. Recrystallized beds contain coarse-grained siderite spar. Larger siderite microspheres, up to about 100 a m diameter, are locally observed in some layers within carbonaceous black shale (Fig. 5C). These spheres are dispersed throughout the shaly matrix and are composed of an inner dark-brown siderite core, which is similar to the spheres in the banded chert-carbonate, and an outer coating of light brown concentrically laminated siderite. Often, two or more of the dark siderite microspheres share a common coating of concentrically laminated siderite. Lithoclasts up to 2 m m in diameter, consisting of tightly packed dark brown siderite microspheres, have a discontinuous coating of light brown siderite (Fig. 5B). Siderite is also commonly observed as fine-grained rhombs within black carbonaceous matrix (Fig. 6A). Siderite near the base of the formation in the Kakabeka Falls area is commonly replaced by pyrite (Fig. 5C ), indicating that siderite formed above the base of the sulphate reduction zone. Laminae and thin layers of siderite interlayered with arenite are texturally similar to siderite layers in banded chert-carbonate. In some cases lithoclasts of siderite are contained within arenite layers adjacent to siderite layers (Fig. 2B). The presence of siderite lithoclasts indi-

W.J. CARRIGANAND E.M.CAMERON

cates that siderite formed close enough to the sediment/water interface to be reworked. Laminated carbonate contains alternating carbonate-rich and clastic-rich laminae and thin beds. Carbonate-rich layers that contain abundant fine-grained matrix, also contain abundant siderite as microspheres and fine-grained rhombs (Fig. 6B ), whereas layers that have little matrix, consist mostly of fine- to coarsegrained ankerite. Clastic-rich layers consist of angular silt- to sand-size quartz and feldspar, well-rounded iron-rich intraclasts that are similar in composition to those in the arenite facies, and, less commonly, elongated sand- to pebble-size intraclasts composed of altered volcanic ash. In these layers siderite is associated with clasts composed of volcanic ash and occur as fine-grained rhombs that tend to be preferentially located around the clast margins (Fig. 6C), suggesting a diagenetic origin for these siderite grains. Siderite was the earliest-formed carbonate mineral in all rocks, except the Upper Limestone Member (Fig. 3). Siderite beds are cut by synsedimentary erosional surfaces and by intraformational breccias. This, and the presence of siderite intraclasts, indicates that siderite was a primary mineral which precipitated from seawater or from pore waters close to the sediment/water interface. The presence of interstitial chert in siderite layers and of siderite grains in chert layers suggest that silica and siderite precipitation was coeval. The origin of siderite microspheres is uncertain. LaBerge ( 1973 ) interpreted them as relics of the microfossil Eosphaera tyleri (Barghoorn and Tyler, 1965), whereas Kazmierczak (1979) compared them to Devonian calcispheres in shallow-water limestones. Others suggest that the

Fig. 6. Micro-photographs of primary textures of siderite (all plane-polarized light). A. Black shale composed of finegrained rhombs of siderite and ankerite in very fine-grained black matrix. Small poikiloblasts of calcite (white) enclose some rhombs of siderite and ankerite indicating calcite is latest carbonate phase. Bar is 0.2 mm. B. Siderite-rich layer from laminated carbonate, composed of siderite microspheres and rhombs (S), surrounded by coarser-grained rombs of ankerite (A), within a fine-grained matrix. Matrix (M), which is isotropic under crossed polars, is probably made up of volcanic ash. Bar is 0.2 mm. C. Elongate clasts (C), possibly of pyroclastic origin, are fringed by rhombs of siderite (S). Siderite and clasts are enclosed by coarse-grained ankerite cement (A). Bar is 0.2 mm.

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PETROLOGICAL AND STABLE ISOTOPE STUDIES OF CARBONATE AND SULFIDE MINERALS, GUNFLINT FORMATION

microspheres are o f inorganic origin and are either siliceous or sideritic structures that were modified by diagenesis or m e t a m o r p h i s m (e.g. Oehler, 1976). During contact metamorphism, siderite recrystallized or was replaced by ankerite (Floran and Papike, 1978 ), or with increasing m e t a m o r p h i c grade, dolomite and calcite. Magnetite, which is a minor component of u n m e t a m o r p h o s e d rocks, becomes significant during the formation o f m e t a m o r p h i c ankerite and increases in abundance with metamorphic grade. Ankerite. Within siderite layers in the banded c h e r t - c a r b o n a t e facies, ankerite often forms fine- to medium-grained rhombs that overgrow or replace earlier-formed siderite, particularly in beds that have been extensively recrystallized. Coarse-grained ankerite is often observed replacing chert, especially adjacent to stylolites that occur at the contacts between siderite and chert bands. Isolated, coarsegrained rhombs of ankerite are c o m m o n within chert layers, where they overgrow primary lamination (Fig. 7B ). In arenite, ankerite exists as layers, cements, and replacements o f earlier-formed minerals. Arenite containing laminae and thin beds o f fine-grained siderite a n d / o r ankerite often contain abundant coarser-grained ankerite, which occurs interstitial to siderite grains and to intraclasts in adjacent arenite beds. In some cases, intraclasts in arenite beds are partially (Fig. 7A) or completely (Fig. 7C ) replaced by coarse-grained ankerite. Isolated lenses and nodules of m e d i u m - to coarse-grained ankerite

357

c o m m o n l y replace volumes o f arenite in beds that do not contain interlayered fine-grained siderite or ankerite. In laminated carbonate, ankerite exists mostly as a c e m e n t to siderite and to clastic grains (Fig. 6, B and C ) and m a y partially replace clastic grains. Ankerite also forms a cement to some pyroclastic beds interbedded with black shale (Fig. 7D ) and partially replaces pyroclastic grains. Ankerite rhombs are dispersed throughout the shale matrix (Fig. 6A). The formation o f ankerite appears to be complex, beginning during early diagenesis and continuing into late diagenesis (Fig. 3 ). Where ankerite coexists with siderite, textures indicate that ankerite formed later than siderite. Coarse-grained ankerite in contact-metamorphosed rocks is recrystallized and shows a decrease in iron content with increasing grade; there is an associated increase in magnetite. Calcite. Calcite is the latest-formed carbonate mineral in the iron-rich rocks (Fig. 3). It exists as small poikiloblasts containing inclusions o f siderite or ankerite, as coarse-grained crystals that partially or completely replace preexisting grains, and as fillings in small veins and fractures. In m e t a m o r p h o s e d rocks, coarsegrained calcite is often intergrown with ankerite or dolomite and with magnetite. Upper Limestone Member. The proportion o f dolomite in the U p p e r Limestone M e m b e r varies from individual medium-grained ( ~ 0.2 m m ) rhombs, dispersed within a loosely packed framework of pyroclastic material (Fig. 8A), to beds o f interlocking coarse-grained

Fig. 7. Micro-photographs of primary and diagenetic textures of ankerite (plane-polarized light except where noted). A. Arenite intraclasts at top left are composed of chert and iron-silicates and are cemented by chert. Intraclasts in right half of photo have been partially replaced by coarse-grainedankerite. Note bleached zone in center, consisting mainly of chert, in which iron has been leached out ofintraclasts to form ankerite during diagenesis. Bar is 2.0 mm. B. Finely laminated chert layer from banded chert-carbonate contains elongated inclusions of organic matter and dispersed siderite microspheres (dark-colored spheres ). Large rhombs of ankerite (center) overgrowchert and siderite during diagenesis;crossed polarizers. Bar is 0.5 mm. C. Arenite containing fine-grainedankerite layer and laminae from Fig. 2D. Intraclasts adjacent to ankerite layer (center of photo) have been replaced by coarse-grainedankerite (light colored), but still preserve grain outlines. Laminae of ankerite (dark laminae in top half of photo) are often deformed around clast outlines suggesting ankerite precipitated coeval with deposition of intraclasts. Bar is 2.0 mm. D. Relatively uncompacted pyroclastic grains, which consist of rock fragments, accretionary lapilli and crystals, are cemented by coarse-grainedankerite, indicating early cementation. Bar is 1.0 mm.

PETROLOGICAL AND STABLE ISOTOPE STUDIES OF CARBONATE AND SULFIDE MINERALS, GUNFLINT FORMATION

rhombs containing interstitial chert and deformed pyroclastic grains (Fig. 8B). In the former case, the interstitial areas to both pyroclastic grains and dolomite rhombs are filled by chert. An isopachous coating of parallel, fibrous chalcedony mantle dolomite and clastic grains with remaining pore spaces filled by radial fibrous chalcedony and, in the centers of the larger pores, blocky quartz. The abundant interstitial chert indicates that silica was an early-formed mineral, as suggested by Simonson ( 1987 ) for chert cements in arenitic ironformation. Therefore, dolomite either formed directly during very early diagenesis or was a very early replacement of primary calcite or aragonite. Calcite occurs as coarse-grained spar that contains abundant fine-grained inclusions, probably of pyroclastic origin, and rhombs of dolomite (Fig. 8C). Calcite also partially replaces chert cement and pyroclastic grains, indicating that it formed later than chert, possibly during late diagenesis.

Sulfide minerals Pyrite is the most abundant sulfide mineral. In the vicinity of diabase dikes pyrite has been partially converted to pyrrhotite. Locally, lateformed cross-cutting fractures are filled with pyrite a n d / o r marcasite or with carbonate containing minor amounts of galena. Pyrite is most abundant in carbonaceous black shale, mainly as very fine-grained euhedral crystals disseminated throughout the organic-rich matrix. In some cases pyrite overgrows siderite grains. Contents vary from trace amounts up to about 2% of the rock. Locally, higher amounts of pyrite are present as thin layers and

359

laminae. At one location, irregular-shaped pyrite concretions are found that display textures suggestive of displacive growth after the sediment had become compacted. However, these concretions are associated with late-formed fractures. Pyrite is generally a minor c o m p o n e n t in other rock types. Disseminated very finegrained pyrite occurs in banded chert-carbonate, in laminated carbonate, and in carbonate layers interbedded with arenite. Inclusions of pyrite in siderite microspheres were noted by Kazmierczak ( 1979 ). Local accumulations of coarser-grained pyrite are observed along stylolites and in siderite and ankerite beds that have undergone significant recrystallization. In arenite and stromatolites, scattered coarsegrained crystals of pyrite occur both within intraclasts and in interstitial areas and often cross-cut grain boundaries. In the Upper Limestone Member, fine-grained pyrite is mostly found associated with mafic pyroclastic grains or disseminated throughout the porelining chalcedony cement. These textures indicate that much of the pyrite formed during very early diagenesis, possibly at or just below the sediment/water interface, and is approximately coeval with siderite. Coarse-grained pyrite was a result of recrystallization during burial. During contact metamorphism, pyrite was converted to pyrrhotite. The lower part of the Gunflint Formation in the Kakabeka Falls area contains an unusually large amount of pyrite. Here, undeformed beds of black shale overlie a sequence of banded chert-carbonate and arenite that have been folded and brecciated as a result of syndepositional faulting. Pyrite is abundant in the deformed beds and is texturally complex, occur-

Fig. 8. Micro-photographsfrom the Upper Limestone Member (all plane-polarizedlight). A. Loose framework of pyroclastic material (dark grey) with interstitial rhombs of dolomite (D). Both are mantled by early-formedparallel fibrous chalcedony.Remaining pore spacesare filled by radial fibrous chalcedony.Thin film of black inclusions,which separates rim and pore-fillingcement, contains abundant very fine-grainedinclusions of pyrite. Bar is 0.2 mm. B. Coarse-grained dolomite with deformed interstitial pyroclasticgrains (dark grey) and minor amounts of calcite and chert. Bar is 0.5 mm. C. Coarse-grainedcalcite contains inclusions of dolomite (D) and abundant inclusionsof fine-grainedpyroclasticmaterial (black), suggestingcalcite formed late. Bar is 0.5 mm.

360

I111Itll IIII fill IIII Itll IIII Ilil IIII III1t111IIII III11111till 1111111 0 11 12 13 14 i5 ta 17 18 Fig. 9. Ellipsoidal pyrite concretion in black shale from the Kakabeka Falls area. Pyrite concretions occur along coarser-grained laminae (lighter colored) within black shale. Inner core of concretion contains coarse-grained pyrite within a matrix of shale, which is texturally similar to coarse-grained pyrite in shale below concretion. The outer shell of the concretion consists of bladed crystals of pyrite that radiate outward from and surround the inner core. Bedding laminae are deformed around concretion indicating that they formed during early diagenesis. Scale in inches.

ring as coarse-grained aggregates and rounded nodules that overgrow recrystallized ankerite, as a general cement to brecciated fragments, and as small veins. The contact with the overlying undeformed shale is marked by a zone, about 0.5 m thick, of thin layers of massive pyrite within a black shaly matrix. Above this zone, over an interval of about 5 m, the black shale contains three types of pyrite. Type 1, very fine-grained, similar to that found in shale elsewhere in the formation, is disseminated throughout the shale matrix. Type 2 is coarse, euhedral grains, dispersed singly throughout the shale matrix, or as aggregates of several grains, or as small lenses elongated parallel to bedding. Coarse-grained pyrite often overgrows the fine-grained pyrite and replaces siderite rhombs and microspheres indicating that it formed later than Type 1. However, primary bedding laminae are deformed around lenses

W.J. CARRIGANAND E.M.CAMERON

of pyrite, indicating that pyrite formed while the sediment was still relatively uncompacted. Type 3 pyrite is ellipsoidal concretions, up to 4 cm in diameter, which are often flattened parallel with bedding. Bedding laminae are also deformed around the concretions, indicating that they formed as the sediments were undergoing compaction. Some contain an inner core of coarse-grained pyrite in a matrix of shale, similar to the Type 2 pyrite, whereas other concretions have a core entirely of coarsegrained pyrite with mutually interfering grain boundaries (Fig. 9 ). The cores are surrounded by a shell of euhedral bladed pyrite crystals that grow radially outward in all directions. The similarity in the textures of Type 2 pyrite and some of the concretion cores suggest that they formed at the same time. Therefore, the concretion rims formed later than Type 2 pyrite, but also during early diagenesis. These concretions were only observed in the Kakabeka Falls area.

Analytical methods Chemical separation techniques similar to those used by Epstein et al. (1964) for coexisting calcite and dolomite and by A1-Aasm et al. (1990) for coexisting calcite, dolomite, and siderite were used in this study. These are based on the differential reaction rates of various carbonate minerals with phosphoric acid at different temperatures. Carbonate mixtures were reacted with 100% phosphoric acid at 25°C for 4 h, after which the CO2 was extracted and labeled calcite. The unreacted portion of the sample are then allowed to continue reacting at 50°C for an additional 16 to 20 h. The CO2 extracted at this time was labeled ankerite, or dolomite for samples from the Upper Limestone Member. Reaction times for ankerite are unknown and are dependent on iron content, but for this study it is assumed that ankerite reaction was similar to dolomite. For samples that contained siderite, the unreacted portion was reacted at 50°C for an additional

PETROLOGICAL AND STABLE ISOTOPE STUDIES OF CARBONATE AND SULFIDE MINERALS, GUNFLINT FORMATION

4 days, after which time the CO2 was collected. Although siderite may take up to 11 days for complete reaction, A1-Aasm et al. (1990) have shown that the isotopic composition of the CO2 extracted after 4 days does not change. The isotopic ratios of carbon and oxygen were determined on the CO2 using a VG Isogas SIRA 12 mass spectrometer at the University of Ottawa. The ~ values are reported with respect to the PDB (Peedee belemnite) reference scale and are precise to + 0.1%. Oxygen isotope fractionation factors for acid-released CO2 are 1.01025 for calcite at 25°C (Freidman and O'Neil, 1977), 1.01065 for dolomite and ankerite at 50°C (Rosenbaum and Sheppard, 1986 ), and 1.010453 for siderite at 50 ° C (Rosenbaum and Sheppard, 1986). Pyrite was physically separated from samples whenever possible; otherwise sulfur was extracted chemically using the KIBA method (Sasaki et al., 1979 ), which produces sulfur as Ag2S. Pyrite and Ag2S was burnt directly, with Cu20, in an electric furnace at 1100 °C to produce SO2, which was analysed on a Micromass 602E mass spectrometer at the University of Ottawa. The ~ values are reported with respect to the CDT (Canyon Diablo troilite) reference scale and are precise to + 0.2%o.

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~ 3C of carbonate minerals The carbon isotope ratios of the various carbonate minerals from the Gunflint Formation have a wide range of values, between + 0.5 and -20o/oo (PDB) (Figs. 10-13). However, unmetamorphosed rocks have a more restricted range, all heavier than -7%0. In unmetamorphosed rocks, siderite has ~I3C values between +0.5 and - 5.2%o (Fig. 10), ankerite between - 0 . 4 and -6.6%o (Fig. 11), and calcite between - 1.2 and -7.7%o (Fig. 12). In the Upper Limestone Member, dolomite is between + 0.5 and - 2.70/oo, and calcite - 1.2 to - 3.9°/oo (Fig. 13 ). There are no significant differences

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Fig. 10. in 8'3C between different facies and no systematic variations with respect to stratigraphy. The different carbonate minerals all show variations of up to 5°o over short vertical distances; however, the heaviest values at any given stratigraphic interval are usually close to 0%o.

362

w.J. CARRIGAN AND E.M. CAMERON

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Fig. 14. A. Histogram of d34S values from the Kakabeka Falls area. Pyrite textures include fine-grained pyrite in black shale (Type 1 ), both fine and coarse-grained pyrite in black shale (both Type 1 and Type 2), pyrite concretions (Type 3), and coarse-grained pyrite from folded and brecciated beds underlying black shale (coarse-grained). B. Histogram ofd34S values for pyrite for all locations from the Gunflint Formation, except the lower 20 m in the Kakabeka Falls area, and for the Rove Formation.

- 8.7%0 a n d calcite b e t w e e n - 3.8 a n d - 8.8%0. S a m p l e s c o l l e c t e d near d i a b a s e sills s h o w a p r o g r e s s i o n to lighter v a l u e s t o w a r d s i n t r u s i v e contacts.

PETROLOGICAL AND STABLE ISOTOPE STUDIES OF CARBONATE AND SULFIDE MINERALS, GUNFLINT FORMATION

~80 of carbonate minerals ~ 8 0 of carbonate minerals lie between - 5.3 and -18.0%o (PDB) (Figs. 10-13). Metamorphosed and unmetamorphosed rocks do not show the same clear-cut distinction as for ~3C: unmetamorphosed rocks range between - 5.3 and - 15.6%o, metamorphosed rocks between - 9 . 6 and - 18.0%o. In unmetamorphosed rocks, siderite is between - 5 . 3 and -13.8%o (Fig. 10), ankerite - 6 . 1 to -15.0%o (Fig. 11 ), and calcite - 6 . 7 to - 15.7%0 (Fig. 12 ). In the Upper Limestone Member, dolomite is between - 9 . 2 and - 14.6 and calcite between - 13.1 and - 15.5%o (Fig. 13). Like the carbon isotope ratios, ~180 values from different rock types show a significant amount of overlap, although most from arenitic rocks tend to fall in the more ~80-depleted range. All carbonate minerals show variations of up to 9%o over short vertical distances; however, the range of values is consistent throughout the entire section. The heaviest values for each type of carbonate are believed to represent the closest to the original value, with lighter ratios the result of diagenetic alteration. Metamorphosed rocks have siderite c~80 values between - 9 . 6 and -13.8%o, ankerite values between - 11.1 and - 16.9%o, and calcite values between - 12.0 and - 17.5%o. In the Upper Limestone Member, dolomite is between - 1 5 . 6 and -18.0%o and calcite between - 17.1 and - 17.5%o. Samples collected in proximity to intrusive sills generally have a more uniform range of c~80 values and have values that overlap or are more depleted in ~80 than unmetamorphosed rocks. ~34S

363

stratigraphy, with most of the variation restricted to the lower 20 m of the Gunflint Formation near Kakabeka Falls (Fig. 14A). Elsewhere, with only a few exceptions, pyrite has a narrow range of values between about + 4 and + 12%0 (Fig. 14B ). Eight samples of pyrite analysed from the 600 m thickness of the overlying Rove Formation have d34S values between +13 and +21%o (Fig. 14B). In the Kakabeka Falls area (Fig. 1 ), coarsegrained pyrite from folded and brecciated beds of banded chert-carbonate and arenite from the lower 15 m of the Gunflint Formation are texturally complex and more than one generation may be present. ~34S values vary from - 18.2 to + 22.0%0 (Fig. 15 ). In the overlying, undeformed black shale, three types of pyrite have previously been described. It is not possible to separate Type l (fine) pyrite from Type 2 (coarse); such samples vary between - 2 and + 35%o and show an upward trend to heavier values over a vertical distance of 5 m (Fig. 15 ). Detailed traverses across pyrite concretions (Type 3 pyrite) yield a complex distribution. Four of these concretions were cut in half and micro-samples taken at 2 m m intervals from rim to core to rim (Fig. 16). These display internal variations between + 7 to + 39%0, with the heaviest values generally from the cores and the lightest values from the rims. These concretions also show a general upward trend to heavier values (Fig. 15 ). Pyroclastic beds appear about 20 m above the base of the formation (Fig. 15), after which only finegrained pyrite (Type 1 ) is present, which has a very narrow range of ~34S values between + 4 and +10%o, similar to the rest of the formation.

of pyrite Sulfur-carbon ratios

Pyrite has a total range o f c~345of 70%0, from - 32.5 to + 35%o (CDT). Sulfur isotope ratios do not show any systematic variations between different rock types or between metamorphosed or unmetamorphosed rocks. Ins t e a d , (~345 variations appear to be related to

A plot of weight percent sulfur versus carbon from black shales of the Gunflint Formation is shown in Fig. 17. These shales contain very fine-grained, disseminated pyrite, except locally where thin laminae and layers are pres-

364

W.J. CARRIGAN AND E.M. CAMERON

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ent. Sulfide layers were not analyzed for sulfur and carbon. Unlike most marine shales of Phanerozoic age (e.g. Berner, 1984), there is no correlation of S with C; the points are mostly distributed along a narrow band near the C axis, at low S concentrations. Black shales in the lower 20 m of the Gunflint Formation at Kakabeka Falls are also shown in Fig. 17 (as triangles). The samples that were analyzed contain variable propor-

tions of both fine- and coarse-grained pyrite (Types 1 and 2 ), but are free of pyrite concretions (Type 3 ). The samples have higher S/C, and the plot shows a great deal of scatter. From petrographic observations, fine-grained pyrite-sulfur contents are generally less than 1%, the scatter being due to high contents of coarsegrained pyrite. At the contact with underlying brecciated banded chert-carbonate, the shale contains layers of massive pyrite containing

PETROLOGICAL AND STABLE ISOTOPE STUDIES OF CARBONATE AND SULFIDE MINERALS, GUNFLINT FORMATION

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366

W.J. C A R R I G A N

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relatively small carbon contents (not shown in Fig. 17).

Interpretation of isotopic data ~13Cdata A feature of many Precambrian iron-formations is the presence of carbonate minerals depleted in 13C relative to normal marine carbonates (Becker and Clayton, 1972; Perry and Tan, 1973; Baur et al., 1985; Beukes et al., 1990). Different explanations have been offered, which have been related to genetic interpretations for the iron-formations. These resolve into two main interpretations: ( l ) ironformation carbonates precipitated from a 13Cdepleted water column (Becker and Clayton, 1972; Beukes et al., 1990); and (2) the ~3Cdepleted carbonates are a result of diagenetic or metamorphic processes (Perry and Tan, 1973; Baur et al., 1985). Becker and Clayton (1972) proposed that carbonate in iron-formations of the Hamersley Group in Australia, which have ~ 1 3 C values between - 6.5 and - 15%0, precipitated from isotopically light bicarbonate in a basin sepa-

AND E.M. CAMERON

rated from the open ocean reservoir of bicarbonate. Oxidation of organic matter, possibly coupled with the reduction of ferric iron, provided the source of light carbonate. Dolomite units in the Hamersley Group, with 6~3C values close to 0%0, are considered to have been deposited during periods when the basin was open to the ocean. Beukes et al. (1990) propose a stratified water column for the deposition of the Kuruman Iron Formation. In their model, siderite and oxide-rich iron-formation was deposited in a deeper ~3C-depleted water mass, whereas limestone and shale was deposited in shallow surface water having a normal marine 6~3C value close to 0%0. They suggest that the ~3C-depleted carbon in the deep water mass was derived in part from hydrothermal water circulating through older organic-rich sediments and in part from volcanic or mantle-derived carbon. Perry and Tan ( 1973 ) noted that carbonate in rocks of the Biwabik Iron Formation that contain magnetite are significantly more ~3Cdepleted than are rocks that do not contain magnetite. They proposed that iron was initially precipitated as a ferric oxide phase and was partially reduced to form magnetite, coupled to the oxidation of organic carbon. This produced light CO2, which exchanged with primary siderite that initially had normal marine isotope ratios. Baur et al. (1985) proposed that free oxygen, produced by seasonal upsurges of photosynthesis, caused the precipitation of iron as a ferric hydroxide, which was buried in the sediment along with organic matter. Bacterial oxidation of organic matter within the sediment zone provided the source of isotopically light carbonate. Dissolved ferrous iron, required to form siderite and ankerire, may have been derived by the bacterial reduction of the ferric hydroxide. If the wide range of carbon isotope ratios in the Gunflint Formation is the result of variations in the 6~3C of the water mass, this should be reflected in stratigraphic variations in the 6~3C values of the carbonate minerals. The pet-

PETROLOGICAL AND STABLE ISOTOPE STUDIES OF CARBONATEAND SULFIDE MINERALS, GUNFLINT FORMATION

rographic evidence indicates that siderite, as well as dolomite in the Upper Limestone Member, was the earliest-formed, and still preserved, carbonate mineral. However, unlike the Australian and South African iron-formations, the range ofd ~3C ratios for siderite and dolomite are the same, regardless of their stratigraphic position, and heaviest values are consistently near 0%o throughout the section, which is the value that would be expected for marine bicarbonate (Veizer, 1983; Anderson and Arthur, 1983 ). Thus the wide range of isotopic ratios most likely do not record fluctuations in the 8~3C of the water mass, but are the result of either diagenetic processes or metamorphism. The restriction of values lighter than about - 6%o to metamorphosed rocks indicate that the most depleted values are a product of metamorphism. The heaviest isotopic ratios for both siderite and dolomite, near 0%o, reflect their primary dl 3C values. Equilibrium fractionation factors predict that siderite should be enriched by about 5%o relative to cogenetic calcite (Kaufman et al. cited in Beukes et al., 1990) and that dolomite should be about 2%o heavier than cogenetic calcite (Veizer, 1983 ). Therefore, under equilibrium conditions, siderite should be 3%o heavier than dolomite. If the 813C ratios near 0%o for both siderite and dolomite represent their primary values, this might indicate that the water mass from which siderite formed was up to 3%o lighter than that for dolomite (e.g. Beukes et al., 1990). However, fractionation in natural systems often does not follow theoretical equilibrium fractionation factors (Veizer, 1983). For example, rapid rates of precipitation may cause 8~3C of the carbonate to approach that of the bicarbonate. Because of the similarity in isotopic compositions of siderite and dolomite in the Gunflint, we interpret these results to indicate a common source of bicarbonate having a similar isotopic composition. The shift to more negative isotopic ratios was most likely the result of diagenetic processes.

367

T h e 8 1 3 C composition of pore waters in organic-rich sediments becomes significantly altered during diagenesis by the oxidation of organic matter to CO2 (Claypool and Kaplan, 1974). Metabolization of organic matter is principally by aerobic bacteria in sediments underlying oxic waters and by dysaerobic bacteria in sediments underlying the aerobic zone. As conditions become truly anoxic, sulphatereducing bacteria become the dominant agent for the oxidation of organic matter, usually at depths of tens of centimeters. In anoxic basins, the top of the sulphate reduction zone may coincide with or rise above the sediment/water interface. Once sulphate is depleted, carbonate reduction and fermentation reactions become involved in the oxidation of organic carbon. During late diagenesis, above about 75 °C, the oxidation of organic carbon is accomplished mainly by thermocatalytic reactions. These processes lead to a depth zonation in the isotopic composition of pore waters (Irwin et al., 1977; Gautier and Claypool, 1984). Bicarbonate in pore water just below the sediment/water interface has the same concentration ( ~ 2 - 3 mM) and 8t3C (~0%o) as the overlying water column. In sediments with abundant organic matter, the content of bicarbonate will increase steadily with depth, as organic matter is bacterially oxidized, ~13C being rapidly shifted to values approaching that of organic carbon ( ~ - 2 5 ° o ) . Oxidation of upward diffusing 13C-depleted methane may cause 8~3C to become even more depleted than the organic carbon (Gautier and Claypool, 1984). Below the sulphate reduction zone, CO2 is used by methanogenic bacteria to produce methane, accompanied by a strong isotope fractionation, causing the residual CO2 to become more enriched in ~3C with depth. Pore water 813C may be as heavy as + 15%o (Irwin et al., 1977 ), but more typically is in the range - 5 to +50/oo (Gautier and Claypool, 1984). During late diagenesis and metamorphism, the 8~3C of C02 produced by the thermocatalytic oxidation of organic matter is similar to the

368

composition of the parent organic matter. The heaviest isotopic ratios, close to 0%o, for both siderite and dolomite of the Upper Limestone Member reflect that of the bicarbonate of the water mass. Values more negative than this indicate that some siderite and dolomite continued to form during early diagenesis or else underwent isotopic exchange with 13C-depleted pore water. Siderite and dolomite that have been significantly recrystallized during metamorphism, have the most depleted 613C values, between - 6 and -10%o, indicating isotopic exchange during metamorphism. Petrographic evidence suggests that ankerite and calcite formed as diagenetic cements and during the recrystallization of primary carbonate minerals. Their ratios, between 0 and -7%0, reflect precipitation from a n d / o r equilibration with 13C-depleted pore waters. Migration of 13C-depleted pore water from organic-rich sediments would explain the presence of isotopically light ankerite and calcite in organicpoor sediments such as the arenite facies. The isotopic data are consistent with the petrographic evidence that siderite formed very early, at burial depths shallow enough to be reworked, even perhaps in the water column. Also, many siderite grains are partially replaced by pyrite, indicating that siderite formed above the base of the sulfate-reduction zone, which may indicate that Fe 2+ was in excess over S2-. In summary, siderite precipitation began either above or just below the sediment/water interface, with marine bicarbonate the principal source of carbon. The principal source of carbon for dolomite in the Upper Limestone Member, which formed during very early diagenesis, was also marine bicarbonate. Ankerite and calcite formed during early to late diagenesis and some siderite recrystallized at this time. The isotopic composition of these minerals is slightly lighter than that of the primary siderite, reflecting the presence in the pore water of bicarbonate produced during the bacterial oxidation of organic matter or methane.

W.J. CARRIGAN AND E.M. CAMERON

During contact metamorphism, isotopic exchange with CO2, produced by the thermal oxidation of organic carbon, shifted 613C values to between - 6 and - 20%o. (~180 data The 6180 of carbonate minerals is mainly determined by the 6180 of the water in which the mineral formed and a fractionation factor that varies with temperature. Present-day ocean water has a mean 6180 value of 0%o (SMOW). Fresh waters vary from 0 to - 50°o (SMOW) as a result of evaporation and condensation cycles that concentrate 180 in the condensed phase. Thus precipitation becomes more depleted in 180 with increasing latitude and elevation (Craig, 1961 ). Marine carbonate rocks, and other marine chemical sediments, show a decrease in mineral 6180 with increasing geologic age. Three explanations have been offered, none of which has been universally accepted. Degens and Epstein ( 1962 ) and Keith and Weber (1964) proposed that rocks have undergone progressive isotopic exchange with meteoric waters since their deposition, causing a time-dependent depletion in 180. Perry and Tan (1972) and Veizer et al. ( 1982 ) propose that seawater has become progressively more enriched in 180 towards the present, due to cycling of seawater through the mantle. Knauth and Epstein (1976), Knauth and Lowe (1979), and Karhu and Epstein (1986 ) suggest that the lighter isotopic ratios reflect a higher temperature for the earlier oceans, with 6180 for the waters remaining constant. The heaviest 6180 values of carbonate minerals from the Gunflint Formation ( - 5 to 7%0 PDB ) are similar to the values reported for least altered calcite and dolomite from late Archean and early Proterozoic marine carbonates, which Veizer et al. ( 1989, 1991 ) consider to be reasonable estimates of carbonate phases formed in equilibrium with coeval seawater. More 180-depleted isotope ratios are the result -

PETROLOGICAL AND STABLE ISOTOPE STUDIES OF CARBONATE AND SULFIDE MINERALS, GUNFLINT FORMATION

of post-depositional alteration. If the fractionation factor of Becker and Clayton ( 1972 ) for siderite is used, the maximum siderite d~80 value of -5.3%o (PDB) indicates either a water temperature of about 45°C, assuming marine d 180 = 0°/oo (SMOW), or else a marine d180 of - 5.5°/o0 (SMOW), assuming a water temperature of 20 ° C. If the heaviest isotopic ratios represent the original marine values, then lighter d180 values are the result of post-depositional alteration or the formation of new carbonate minerals during diagenesis or metamorphism. Perry and Tan ( 1973 ) suggested that the light d~80 values were a result of oxidation-reduction reactions which produced shifts of oxygen from hematite or ferric hydroxides to carbonate. This would take place via the pore water. If the fractionation factor for ferric hydroxide is similar to magnetite, its d ~80 will be, at most, 5%0 lighter than seawater (Becker and Clayton, 1976 ). However, the oxygen exchanged is only a small percentage of that present in the pore water reservoir. Thus it is unlikely that oxidation-reduction reactions caused the wide range of d~80 values in the Gunflint Formation. Also, hematite is only a minor component of the Gunflint, occurring mainly in the arenite facies. Beukes et al. (1990) observed that siderite in the Kuruman Iron Formation is slightly more depleted in 180 than calcite in the underlying limestone. Because siderite should theoretically be enriched by about 1o/oorelative to calcite, if isotopic equilibrium was attained, they propose that the water column during deposition of the Kuruman Iron Formation may have been stratified with regard to d~80, with deeper water having lighter isotopic ratios due to the addition of 180-depleted hydrothermal water. However, in the Gunflint Formation, siderite values are as much as 4%o heavier than dolomite from the Upper Limestone Member, opposite to that predicted by the above model. There are two realistic possibilities for the wide range of ~80-depleted ratios in the Gun-

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flint. The first is re-equilibration at higher temperature in a system closed to oxygen; the second is exchange with meteoric waters in an open or partially open system. Given the maximum burial temperature of 130°C for unmetamorphosed Gunflint rocks (Miyano, 1987), the most 180-depleted siderite value of -13.8%0 (PDB) falls within the expected range for the closed system re-equilibration given the starting composition noted above of - 5 to -7%0 (SMOW). However, it is not possible to distinguish this process from opensystem exchange discussed below, d ~80 values of carbonate minerals from metamorphosed rocks are more uniform and generally lighter than values from unmetamorphosed rocks, so it seems reasonable to ascribe this to equilibration at metamorphic temperatures. Gregory (1986) and Gregory and Criss (1986) proposed that most iron-formations reflect open-system disequilibrium systematics during early diagenesis or metamorphism. They suggest that isotopic exchange with a meteoric fluid having a d~80 value of about - 10% (SMOW) could produce the isotopic ratios observed in iron-formations. Alternatively, the isotopic ratios may reflect partiallyopen, multiphase systems where the d180 of the pore water is buffered internally by the isotopically heavy carbonate minerals (Veizer et al., 1990). Knauth and Epstein ( 1976 ) report a 619 of -73°/oo and a dlSO of 23.6°/oo (SMOW) for a sample of chert from the Gun flint Formation, which may indicate a component of meteoric water. The shallow shelf setting and proximity to a shoreline implies that meteoric water could have been introduced.

Sulfur-carbon ratios Elemental abundances of organic carbon and sulfide-sulfur in normal marine sediments typically show a positive correlation that passes through the origin on a S-C plot, caused by bacterial sulphate reduction during early diagenesis (Berner, 1970, 1984; Goldhaber and

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Kaplan, 1975; Raiswell and Berner, 1986 ). The amount of sulfide-sulfur in sediments is dependent upon the sulphate supply, the amount of organic matter that can be readily metabolized, and the abundance of reactive iron to fix HzS as iron sulfide. In normal marine environments (oxygenated bottom waters), clastic sediments contain abundant reactive iron and most sulphate reduction takes place within the bioturbation zone where sulphate is abundant. Therefore, organic matter is the controlling factor on the amount of pyrite formed. Rapidly deposited sediments, and sediments underlying areas of high productivity, will have higher contents of organic matter, and hence sulfide-sulfur, whereas slowly deposited sediments and those that underlie areas of low productivity will have low contents of carbon and sulfur. Recent sediments have an average S/C ratio of 0.36 (Sweeney and Kaplan, 1980 ). In marine environments with anoxic bottom waters, the waters may contain H2S, permitting pyrite to form in the water column. This can lead to high S/C ratios with positive S intercepts on S-C plots but a poor correlation (Raiswell and Berner, 1985 ). Sediment-hosted ore deposits often contain sulfide-sulfur contents in excess of 10%, but contain organic carbon contents less than 1% (Ohmoto et al., 1990), which indicates that the accumulation of sulfide in these sediments did not take place by simple biogenic reduction of sulphate during early diagenesis. In fresh waters, which typically contain very low sulphate concentrations, only minor amounts of pyrite can form. This leads to C versus S plots where the sample points are distributed in a narrow band along the C axis at low S contents. In such rocks, S is essentially uncorrelated with C and the S/C ratio averages ~ 0.03 (Berner, 1984). Carbonaceous black shale of the Gunflint Formation was deposited under anoxic conditions in which easily metabolizable organic matter, such as bacteria and simple algae, were likely abundant. Black shale samples, except for the basal 20 m of the Kakabeka Falls section,

W.J. CARRIGAN AND E.M. CAMERON

mostly contain less than 1% sulfide-sulfur and low S/C ratios. A S-C plot (Fig. 17) is flat along the C axis, partly simulating the plot for freshwater sediments of Berner (1984). The principal difference is that the freshwater sediments are even lower in S, with values < 0.5%. Given this lack of correlation between S and C, there are two possible origins for the sulfide in the Gunflint. The first is that the sulfide is of hydrothermal-exhalative origin as suggested by Cameron (1983a). However, the low S/C ratios are atypical of such mineralization. Also, at least part of the pyrite is of early diagenetic origin, since it replaces siderite. The second, and more likely, explanation is that the sulfide derived from biogenic reduction of a limited supply of sulphate. Reduction may have occurred in both the water column and the upper part of the sediment. The local presence of sulfide laminae suggests some precipitation within the water column. Formation of pyrite was likely completed in the topmost part of the sedimentary column, limited by exhaustion of the sulphate supply. Many samples contain excess amounts of carbon and all samples have excess iron in the form of siderite. Therefore, the low S/C ratios were the result of low sulphate concentrations. Black shale from undeformed shale in the lower 20 m of the Kakabeka Falls section contain coarse-grained and nodular pyrite, in addition to the usual fine-grained pyrite. Data for these samples, but excluding nodular pyrite, is shown in Fig. 17. These have higher S/C ratios and sulfide contents of up to 18%. The very high sulfide contents occur in sediments that contain much less organic carbon, which is inconsistent with simple biogenic reduction of marine sulphate. These undeformed shales overlie highly pyritic, brecciated and folded arenite and banded chert-carbonate beds, which are associated with a synsedimentary fault. This suggests that sulfur may have been introduced to these sediments by fluids moving upward along this fault.

PETROLOGICAL AND STABLE ISOTOPE STUDIES OF CARBONATE AND SULFIDE MINERALS, GUNFLINT FORMATION (~348

data

t~34S of H2S produced during bacterial sulphate reduction is controlled by the t~345of the dissolved sulphate, kinetic isotope effects, and the open or closed nature of the system with respect to sulphate. The ~34S of marine sulphate has fluctuated between + l0 and + 35%0 since the late Proterozoic, as recorded by the isotopic composition of marine evaporites (Claypool et al., 1980). Because of the scarcity ofevaporites older than about 900 Ma, the isotopic composition of Precambrian oceans is not well known. For the Archean, many workers consider that the rarity of evaporites, together with most sedimentary sulfides and sulphates having compositions near to 0%o, to indicate low dissolved sulphate concentrations in the oceans a n d / o r the absence of sulphate-reducing bacteria (Goodwin et al., 1976; Donnelly et al, 1977; Lambert et al., 1978; Cameron, 1982; Skyring and Donnelly, 1982, Hattori et al., 1983). In these circumstances, sedimentary sulfides were mainly the product of inorganic processes. There is some disagreement concerning the time when bacterial reduction of a significant oceanic reservoir of sulphate became important. However, most workers have indicated ages between about 2.7 Ga (Goodwin et al., 1976; Ripley and Nicol, 1981) and 2.3 Ga (Cameron, 1982, 1983b). There are a limited number of estimates of the isotopic composition of the oceanic reservoir of sulphate in early Proterozoic time. Three estimates from the interval 2.0 to 2.3 Ga are uniform in the range + 14_+3%o. These comprise data from the Gordon Lake Formation, Ontario (Cameron, 1983b), Malmani Formation, Transvaal (Buchanan and Rouse, 1982), and Kona Dolomite, Michigan (Hemzacek et al., 1984). An estimate of + 27.8%0 is reported by Ueda et al. (1990) for a dolomite-mudstone sequence containing sulphate molds and halite casts near the base of the ~1.9 Ga Belcher Group, N.W.T., Canada. The age of the Belcher Group

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is poorly constrained, so this estimate should not be directly applied to the Gunflint Formation. These data do, however, indicate that the isotopic composition of seawater sulphate during lower Proterozoic time varied within the same general range as during the Phanerozoic. The kinetic isotope effect associated with the bacterial reduction of sulphate to H2S ranges between about - 7 5 and 0%0 (Ohmoto et al., 1990). Fractionation associated with the formation of pyrite from H2S is usually negligible, thus the pyrite ~34S closely approximates that of the H2S. Reduction of sulphate by anaerobic bacteria of the genus Desulfovibrio proceeds by a series of reactions that can be divided into four principal steps, each with a characteristic isotope effect (Kaplan, 1983; Holser et al., 1988 ). The first two steps are the uptake of external sulphate by the cell followed by its activation by ATP (adenosine triphosphate) to form APS (adenylyl sulphate). The isotope effect associated with these two steps is small, in the order of a few per rail (Holser et al., 1988 ). The following steps constitute the enzymatic reduction of APS to sulfite, followed by the reduction of sulfite to sulfide. Isotope effects for these are larger, estimated as -25°/oo by Rees (1973) and - 1 0 to - 1 5 by Chambers et al. (1979) for step 3 and 0 to -33%0 (Holser et al., 1988) for step 4. The overall isotope effect will depend on which of these steps is rate limiting. Laboratory experiments, with an unlimited sulphate supply, show that the overall isotope effect is inversely proportional to the rate of sulphate reduction (Harrison and Thode, 1958; Kaplan and Rittenberg, 1964; Kemp and Thode, 1968), which is dependent upon bacterial species, availability of nutrients, and temperature. In modern sediments the kinetic isotope effect falls in the range -45_+20%o (Ohmoto et al., 1990), indicating that steps 3 and 4 are the rate-limiting reactions. Harrison and Thode (1958) found that the kinetic isotope effect decreased as sulphate concentrations were decreased, suggesting that at low

WJ. CARRIGANAND E.M.CAMERON

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sulphate concentrations steps 1 and 2 are the rate-limiting reactions. It was argued in the previous section that the low S/C ratios of the Gunflint shales indicated that the formation of sulfide was limited by the amount of sulphate present in the water column. The isotopic composition of sulfide, in a narrow range of + 4 to + 12%o, is consistent with this. At low sulphate concentrations, steps 1 and 2 above, with low fractionation effects, become the rate-limiting reactions. The resulting sulfide will have t~345 values closer to the parent sulphate than will "normal" marine sulfide. Experimental data show that at sulphate concentrations of 10 and 0.6 mM (compared to 28 mM for the modern ocean), the fractionation effect is reduced to - 1 5 and -4%o, respectively (Harrison and Thode, 1958), and is reversed at 0.01 mM. In the absence of large fractionation effects, the sulfides will tend to have more uniform isotopic compositions than modern marine sulfide. Given a low concentration of sulphate and abundant organic carbon, the formation of diagenetic pyrite will tend to be confined to the topmost few centimeters of sediment, below which sulphate in pore water will be exhausted. This total reduction of sulphate will also cause the sulfide to be close to the parent sulphate in isotopic composition. Petrographic evidence was presented earlier for the presence of both syngenetic and diagenetic pyrite in the Gunflint shales. This is similar to modern anoxic basins, such as the Black Sea. Because the t~345 of massive sulfide deposits of post-Archean age (Franklin et al., 1981 ), which are presumed to be ofhydrothermal-exhalative origin, is often similar to the finegrained sulfide of the Gunflint shales, and may be homogeneous in isotopic composition (Cameron, 1983a ), it is difficult to rule out its presence. However, given the petrographic observations that at least some pyrite is of diagenetic origin, and since a biogenic origin accords with other evidence discussed above, there is no need to invoke any significant ex-

ternal flux of sulfide of hydrothermal origin. The lower 20 m of the Gunflint, in the Kakabeka Falls area, contains texturally complex pyrite of highly variable isotopic composition. Coarse-grained, discordant pyrite in brecciated and folded arenite and banded chertcarbonate beds have ~34S values between - 18 and +22%o. Because these sulfides formed later than fine-grained pyrite, an external source of sulphate is favored, which may have been introduced by fluids moving upward along synsedimentary faults. Black shales overlying the folded and brecciated beds contain three types of pyrite: fine-grained disseminated pyrite, coarse-grained pyrite, and pyrite concretions. Shale containing Types 1 and 2 h a v e ~348 values between - 2 and + 35%o that show an upward trend to heavier values. Pyrite concretions overgrow the first two types of pyrite, but formed during early diagenesis. The concretions have a similar ~34S distribution to coarse-grained pyrite in both the brecciated beds and undeformed shales. In addition, an unusual feature of these concretions is the internal isotopic zoning, whereby (~345 values usually show an decrease outward from the cores to the rims. This zoning is the opposite of that expected for pyrite that had formed as a result of continuous sulphate reduction from an initially open to a progressively more closed system, which is typical for most sedimentary systems (e.g. Coleman and Raiswell, 1981; Raiswell, 1982). The similarity in t~345 distribution between brecciated beds and undeformed shales, their association with possible synsedimentary faults, and the timing of pyrite formation implies a common external source of sulfur. Discussion

The primary depositional phases and the source of the chemical components of iron-formations is still a matter of some controversy. The primary, or earliest-formed, minerals in the Gunflint Formation are considered to be

PETROLOGICAL AND STABLE ISOTOPE STUDIES OF CARBONATE AND SULFIDE MINERALS, GUNFLINT FORMATION

siderite, chert, greenalite, hematite, and pyrite as well as dolomite in the Upper Limestone Member (Floran and Papike, 1975, 1978; Klein, 1983; Simonson, 1987; this study). Other common minerals, such as ankerite, minnesotaite, and magnetite most likely formed as replacements of primary minerals during diagenesis or metamorphism. Siderite is the most abundant iron-bearing mineral in the banded chert-carbonate and laminated carbonate facies and is a common mineral in black shale, but is a minor component of the arenite and stromatolite facies. Greenalite, followed by hematite, is the most abundant ironbearing mineral in the latter two facies, although ankerite is locally abundant. Chert is a dominant component of banded chert-carbonate, arenite and stromatolites and a minor component in laminated carbonate and black shale. Pyrite is sparsely disseminated throughout all rock types but increases in abundance in black shale. The average chemical compositions of Archean ( > 2.5 Ga) and early Proterozoic ironformations are similar (Gole and Klein, 1981; Davy, 1983), perhaps indicating a common source of dissolved iron, derived either from a continental or a marine source. Lateritic weathering of the continent under an anoxic atmosphere, and transport of Fe 2+ in surface waters, has been suggested by some researchers (Garrels et al., 1973; Garrels, 1987; Lepp and Goldich, 1964). However, the large amounts of terrigenous detritus that would be associated with rivers large enough to transport the quantity of Fe 2+ required are not present (Holland, 1973). Moreover, the atmosphere during the early Proterozoic appears to have been at least mildly oxidizing (Walker et al., 1983; Holland, 1984; Kasting, 1987; Holland et al., 1989), which would limit the transport of Fe 2+ in surface waters. Evidence for an oxidizing atmosphere in Gunflint time is that the dominant iron minerals in clastictextured iron-formation are hematite and greenalite. These sediments were deposited on

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a marine platform in a shallow energetic, relatively oxidizing environment. The upward transition from iron-rich sediments of the Gunflint to iron-poor limestone/ dolostone, and then to iron-poor shale of the Rove Formation, indicates a cessation in the supply of Fe 2+, although the precipitation of silica continued during deposition of the Upper Limestone Member. fil3C and 6180 compositions for the earliest formed carbonate minerals deposited both above and below this transition are consistent with precipitation from the same marine source of bicarbonate. This implies that the iron-rich rocks were not deposited during a period of basin emergence as suggested by Becker and Clayton ( 1972 ). A marine source of Fe 2÷ is favored by Holland (1973, 1984), Drever (1974), Morris and Horwitz (1983), and Simonson (1985), among others. There are different interpretations as to the source of iron within the marine environment, with opinion divided between reduction of detrital Fe 3÷ and a volcanic-hydrothermal component. In modern marine waters, the sulphate concentration, at 28 mM, is in sufficient excess that it does not limit sulphate reduction. Boudreau and Westrich (1984) show that sulphate must drop to 4% or less of this concentration to significantly influence rates of reduction. Thus given the high organic matter content of anoxic water columns, sufficient H2S is generated to fix all available Fe 2÷. M o d e m anoxic waters, such as in the Black Sea (Brewer and Spencer, 1974) and Saanich Inlet (Jacobs and Emerson, 1982), contain <0.01 ppm Fe. It is estimated that 1 ppm (Holland, 1984) to 20 ppm (Ewers, 1983 ) of dissolved iron were required to form iron-formation. The S/C data given here indicate that, compared to m o d e m anoxic waters, the Gunflint shales were deposited in waters of limited sulphate concentration. This permitted the generation of sufficient H2S to fix only a small amount of Fe 2+. Thus iron could be kept in solution and transported. This interpretation of low sulphate

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concentrations is supported by the rather uniform and positive 8 3 4 S values for the widely distributed, fine-grained pyrite of the Gunflint. Pyrite formed in marine sediments, including some of lower Proterozoic age (Cameron, 1982; Lambert and Donnelly, 1990), is usually depleted in 34S because of the strong isotopic fractionation effects during biogenic reduction. The isotopic composition of the fine-grained pyrite in the Gunflint suggests only a moderate degree of fractionation from the parent sulphate. Data by Harrison and Thode (1958) show a sharp drop in the degree of fractionation as the sulphate concentration falls from 10 to 0.6 mM. Drever (1974) proposed that sulfide was quantitatively removed from the water column in the deep anoxic portions of the ocean, which would allow iron to buildup in solution. Dissolved sulphate would have been held to concentrations of ~ 0.2 mM. Cameron (1982) suggested that sulphate was maintained at very low levels in Archean ocean water by a greater rate of exchange of reduced mantle material with the ocean. A substantial proportion of the Earth's heat flow is dissipated by the passage of seawater through ocean ridges. The higher heat flow of Archean time would thus have led to a high flux of water through ridges and other centers of submarine volcanism. Although heat flow declined with time, the flux would have remained substantial during lower Proterozoic times, buffering sulphate, as well as dissolved oxygen, in seawater to lower levels than today. In a stratified ocean or basin, with ongoing volcanism, the lower, anoxic layer may have been maintained at even lower sulphate concentrations than the ocean as a whole. Volcanic activity and hydrothermal solutions have been proposed as the source of iron for iron-formations (e.g. Trendall and Blockley, 1970; Goodwin, 1973; Gross, 1980; Cameron, 1983a), particularly for the Archean deposits which are associated with volcanic rocks. A direct volcanic source of iron for the Hamersley iron-formations was rejected b y Hol-

w.J. CARRIGAN AND E.M. CAMERON

land ( 1973 ), who favored the reduction ofdetrital ferric iron minerals to soluble ferrous iron in the deep anoxic ocean. However, it is now known that submarine hydrothermal systems can release substantial amounts of iron. For example, Bowers et al. ( 1988 ) report up to 10.8 m M Fe in vent fluids from the East Pacific Rise. Such fluids have been considered to be a plausible source of ferrous iron for iron-formations (Morris and Horwitz, 1983; Fryer, 1983; Dymek and Klein, 1988; Barrett et al., 1988; Jacobsen and Pimentel-Klose, 1988). Submarine hydrothermal systems thus provide a means for both providing a large supply of dissolved iron and to help buffer the sulphate content of anoxic waters to sufficiently low levels that the transport of iron is not limited by its precipitation close to the source as sulfide. Beukes et al. (1990) propose a hydrothermal source for both iron and isotopically light bicarbonate, which were concentrated in the deeper water mass of a stratified water column. Our interpretation differs in that we consider the z3C and lgO.depleted carbonate minerals in the Gunflint Formation to have been a result of alteration during diagenesis and metamorphism. Hydrothermal fluids may also have provided a source of dissolved silica. However, because silica-secreting organisms probably did not exist at this time, seawater was probably already saturated with respect to amorphous silica, or some other siliceous phase (Drever, 1974; Ewers, 1983 ). This might account for the lack of a direct correlation between iron and silica in the Gunflint and other ironformations. Considering an individual iron-formation, like the Gunflint, the question arises as to whether there was a general, ocean-wide source of iron or if the source was local. Recently, it has been proposed that the Animikie Group was deposited in a foredeep or backarc basin (Hoffman, 1987; Southwick et al., 1988), rather than on a passive margin having open circulation with the ocean. Magmatism was

PETROLOGICALAND STABLEISOTOPE STUDIESOF CARBONATEAND SULFIDE MINERALS,GUNFLINT FORMATION

active in the Animikie basin, as evidenced by interbedded tuff in the Gunflint Formation. The source of volcanic activity was most likely the Emperor Volcanic Complex, which is interlayered with iron-formation in northwestern Michigan (Greenberg and Brown, 1983), and which appears correlative with the Gunflint. Restricted circulation with the open ocean would have the result that iron would be concentrated within this basin rather that being dispersed throughout the ocean. Also, oxygen and sulphate would be more effectively removed from seawater. However, other than showing that ~ 13C and ~180 are consistent with deposition of the primary carbonate minerals of the Gunflint from marine waters of normal composition, this study does not provide evidence relating to basis restriction versus open ocean conditions. Conclusions Siderite was the earliest formed carbonate mineral in the iron-rich portion of the Gunflint Formation. Precipitation was initiated at or close to the sediment/water interface. Ankerite and calcite formed during early to late diagenesis, as authigenic minerals and as replacements of earlier formed minerals, and during contact m e t a m o r p h i s m by the breakdown of siderite. Dolomite is the earliest formed and still preserved carbonate mineral in the Upper Limestone Member. Disseminated, fine-grained pyrite formed during very early diagenesis and, also, partly in the water column. Locally, in the Kakabeka Falls area, coarse-grained pyrite and pyrite concretions post-date fine-grained pyrite, but also formed during early diagenesis. The OI3C values from u n m e t a m o r p h o s e d carbonate minerals near 0%o indicate that marine bicarbonate was the d o m i n a n t source of carbon. The shift to lighter values was the result of incorporation of oxidized organic carbon during diagenesis (0 to - 7%o) and metamorphism (<-70/oo). The heaviest 8 1 8 0

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values for unmetamorphosed carbonate minerals, -5.3%0 (PDB) for siderite, -6.1°/0o for ankerite, and -6.7%o for calcite, are similar to 180 values reported for least altered calcite and dolomite in Precambrian marine carbonate formations, that are considered to represent original marine values. The shift to lighter values probably represents isotopic exchange reactions with pore fluids at higher temperatures a n d / o r isotopic exchange with 180-depleted meteoric water. Low S / C ratios from black shales in the Gunflint Formation are consistent with bacterial sulphate reduction within an anoxic water column where sulphate concentrations were low, limiting the formation of sulfide. A relatively uniform range of positive ~345 values between + 5 and + 12%o is consistent with bacterial sulphate reduction from a low concentration of dissolved sulphate. Locally, high S / C ratios in the Kakabeka Falls area, with ~345 values between - 33 and + 35%0, suggest formation by different means. It is suggested that sulphate here was partly supplied by fluids moving along syndepositional faults. The overlying Rove Formation has 834S values between + 13 and + 21 o/o0,which indicates that the basin was partially closed to the open ocean supply of sulphate over long intervals of time. This is consistent with the interpretation that the Animikie Group was deposited in a basin, such as a foredeep or backarc basin, with restricted communication with the open ocean. The trend to heavier ~345 values in the Rove Formation and the similarity in ~I3C and ~180 composition of carbonates in the Gunflint indicates that the transition from iron-formation to the iron-poor limestone member does not represent deposition of iron-formation during a period of basin isolation (e.g. Becker and Clayton, 1972 ) followed by an incursion of the open ocean during deposition of the limestone member. Instead, a stratified ocean in which iron was in solution below a redox boundary is favored. Expansion of the redox boundary to shallow water would have allowed the trans-

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port of dissolved ferrous iron to the shallow shelf. The transition to iron-poor conditions of the Upper Limestone Member could only have been caused by the contraction of the redox boundary. The widespread evidence of penecontemporaneous volcanism suggests a hydrothermal origin for the iron and possibly silica. Expansion and contraction of the redox boundary may have been related to cycles of tectonic activity, i.e. transgressions and regressions. During deposition of iron-formation, siderite precipitation would be favoured on deeper or sheltered parts of the shelf where conditions were anoxic. A limited amount of pyrite may have formed above the sediment/water interface in anoxic water, but most pyrite probably formed within the sediment by bacterial sulphate reduction of porewater during very early diagenesis. In shallower or more energetic environments, oxidizing conditions would favour the precipitation of iron as a ferric hydroxide. Greenalite, or some amorphous precursor, would precipitate in transitional environments. Evaporation of seawater would enhance the precipitation of silica in all environments. Acknowledgements Funding for this study was provided by an NSERC Scholarship and a University of Ottawa Scholarship to WJC and by an NSERC Grant to EMC. We would like to gratefully acknowledge discussions with R.J. Shegelski and P.W. Fralick; the assistance of B.E. Taylor, P. Middlestead, and G. St. Jean during the isotopic analyses; and E. Hearn for some of the photographic work. We would also like to thank I.R. Jonasson and two anonymous journal reviewers for reviewing earlier versions of the manuscript.

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