Ancient geochemical cycling in the Earth as inferred from Fe isotope studies of banded iron formations from the Transvaal Craton

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1 Contrib Mineral Petrol (2003) 144: DOI /s x ORIGINAL PAPER Clark M. Johnson Æ Brian L. Beard Æ Nicolas J. Beukes Cornelis Klein Æ Julie M. O Leary Ancient geochemical cycling in the Earth as inferred from Fe isotope studies of banded iron formations from the Transvaal Craton Received: 14 February 2002 / Accepted: 23 September 2002 / Published online: 30 November 2002 Ó Springer-Verlag 2002 Abstract Variations in the isotopic composition of Fe in Late Archean to Early Proterozoic Banded Iron Formations (BIFs) from the Transvaal Supergroup, South Africa, span nearly the entire range yet measured on Earth, from 2.5 to +1.0& in 56 Fe/ 54 Fe ratios relative to the bulk Earth. With a current state-of-the-art precision of ±0.05& for the 56 Fe/ 54 Fe ratio, this range is 70 times analytical error, demonstrating that significant Fe isotope variations can be preserved in ancient rocks. Significant variation in Fe isotope compositions of rocks and minerals appears to be restricted to chemically precipitated sediments, and the range measured for BIFs stands in marked contrast to the isotopic homogeneity of igneous rocks, which have d 56 Fe= 0.00±0.05&, as well as the majority of modern loess, aerosols, riverine loads, marine sediments, and Proterozoic shales. The Fe isotope compositions of hematite, magnetite, Fe carbonate, and pyrite measured in BIFs appears to reflect a combination of (1) mineral-specific Electronic supplementary material to this paper can be obtained by using the Springer LINK server located at /s x. C.M. Johnson (&) Æ B.L. Beard Department of Geology and Geophysics, University of Wisconsin, Madison, WI 53706, USA clarkj@geology.wisc.edu Tel.: Fax: N.J. Beukes Department of Geology, Rand Afrikaans University, P.O. Box 524, Aucklandpark 2006, South Africa C. Klein Department of Earth and Planetary Sciences, 200 Yale Blvd NE, University of New Mexico, Albuquerque, NM 87131, USA J.M. O Leary Department of Geological and Planetary Sciences, MC , 1200 E. California Blvd., California Institute of Technology, Pasadena, CA 91125, USA Editorial responsibility: T.L. Grove equilibrium isotope fractionation, (2) variations in the isotope compositions of the fluids from which they were precipitated, and (3) the effects of metabolic processing of Fe by bacteria. For minerals that may have been in isotopic equilibrium during initial precipitation or early diagenesis, the relative order of d 56 Fe values appears to decrease in the order magnetite > siderite > ankerite, similar to that estimated from spectroscopic data, although the measured isotopic differences are much smaller than those predicted at low temperature. In combination with on-going experimental determinations of equilibrium Fe isotope fractionation factors, the data for BIF minerals place additional constraints on the equilibrium Fe isotope fractionation factors for the system Fe(III) Fe(II) hematite magnetite Fe carbonate. d 56 Fe values for pyrite are the lowest yet measured for natural minerals, and stand in marked contrast to the high d 56 Fe values that are predicted from spectroscopic data. Some samples contain hematite and magnetite and have positive d 56 Fe values; these seem best explained through production of high 56 Fe/ 54 Fe reservoirs by photosynthetic Fe oxidation. It is not yet clear if the low d 56 Fe values measured for some oxides, as well as Fe carbonates, reflect biologic processes, or inorganic precipitation from low-d 56 Fe ferrous-fe-rich fluids. However, the present results demonstrate the great potential for Fe isotopes in tracing the geochemical cycling of Fe, and highlight the need for an extensive experimental program for determining equilibrium Fe isotope fractionation factors for minerals and fluids that are pertinent to sedimentary environments. Electronic supplementary material to this paper can be obtained by using the Springer LINK server located at Introduction Understanding geochemical cycling of the redox-sensitive, major element Fe over the history of the Earth

2 524 remains an outstanding problem that has implications for the development of an oxygenated atmosphere and the origin of life. Iron-rich deposits known as iron formations, or where laminated, banded iron formations (BIFs), have been an important subject of discussions of ancient geochemical cycling of Fe because they are found on every continent and have formed periodically over much of the Precambrian (e.g., James 1954; Gross 1965; Trendall 1968; Beukes and Klein 1992). The largest BIF sequences are preserved in the Hamersley Range, Australia (e.g., Trendall and Blockley 1970) and the Transvaal Supergroup, South Africa (e.g., Beukes 1983), and these are generally thought to have been deposited during periods of relative tectonic stability (e.g., Cisne 1984). The average oxidation state of Archean through mid-proterozoic BIFs is Fe 2.4+, significantly higher than that of igneous rocks (Klein and Beukes 1992), and many studies have proposed that iron formations reflect oxidation of ferrous Fe-rich hydrothermal waters in a stratified ocean basin as they interacted with relatively oxygen-enriched layers in the upper water column (e.g., Drever 1974; Ewers 1983; Holland 1984; Morris and Trendall 1988). The mechanisms by which oxidation of ferrous Fe sources has occurred has been extensively debated. Abiologic photochemical oxidation of ferrous iron has been proposed as a mechanism for BIF formation (e.g., Braterman and Cairns-Smith 1987), which, if significant, would suggest an atmosphere that had low O 2 contents. An indirectly biological origin was proposed by Cloud (1965, 1968), as well as many others, where oxidation of ferrous Fe occurred during build up of photosynthetically produced oxygen in the atmosphere. A directly biological origin for BIFs was proposed by Widdel et al. (1993), where anoxygenic photosynthesis was responsible for oxidation of ferrous Fe; this proposal envisions oxidation to have occurred under relatively O 2 -poor conditions. In this contribution we present the first detailed Fe isotope study of BIFs from the superbly preserved Late Archean to Early Proterozoic Transvaal sequences. Although the field of Fe isotope geochemistry is still in its infancy, we can now set the Fe isotope variations observed in BIFs in the context of the global reservoirs that affect Fe isotope variations in the Earth, and compare these with a growing database of experimental determinations of Fe isotope fractionation factors in both biologic and abiologic systems. We show that Fe isotope variations in ancient chemical sedimentary rocks such as BIFs span almost the entire range yet observed on Earth, and that they have potential for distinguishing between inorganic precipitation of Fe-bearing minerals and precipitation that has been catalyzed by bacteria. In addition, Fe isotope differences between specific minerals can help constrain isotopic fractionations that have been predicted from spectroscopic data, and also bear on models for diagenetic modification of iron deposits after initial precipitation. Late Archean and Early Proterozoic Banded Iron Formations of the Transvaal Supergroup Banded iron formations have occurred at sporadic times throughout Earth s history, including well-know occurrences at 3.8 Ga (Isua), 2.5 Ga (Transvaal Hamersley), Ga (Superior Labrador Trough), and 0.7 Ga (Rapitan; e.g., Klein and Beukes 1992). Gross (1965) classified iron formations into four types, two of which continue to be used today, including Superior and Algoma types; the former are laterally extensive deposits and are thought to have been deposited in a passive continental setting, whereas the later have large volcanic components and seem likely to have been deposited in active tectonic settings, possibly related to volcanogenic mineral deposits. The Superior-type Transvaal Hamersley sequences are among the largest, most laterally extensive and best preserved deposits, which were subjected to only low-grade metamorphism, although some sections were considerably metamorphosed (e.g., Miyano 1982; Miyano and Beukes 1984). The iron formations of the Transvaal Supergroup that span the Archean Proterozoic boundary (2.5 Ga) are equivalent to the Hamersley Basin iron formations in size, and proposed correlations of the Pilbara and Transvaal cratons raises the possibility that the two BIF sequences are broadly correlative (Cheney 1996). In the Transvaal Supergroup, temporally correlative BIFs span the Kaapvaal craton, from the Griqualand West sequence in the SW, to the Transvaal Bushveld Basin in the NE (Beukes 1983; Eriksson et al. 1993). The best known are the Kuruman and Griquatown iron formations in the Griqualand West sequence, and the correlative Penge Iron Formation in the Transvaal Bushveld Basin. We focus here on the Griqualand West rocks because they did not experience regional metamorphism, and were generally heated only to 110 to 170 C (Miyano and Beukes 1984), in contrast to the Penge Iron Formation, which has been metamorphosed to temperatures of up to 500 C by the 2.05-Ga Bushveld complex (Miyano et al. 1987; Miyano and Beukes 1997). The Kuruman and Griquatown iron formations are chemically continuous over hundreds of kilometers, and the Kuruman Iron Formation has a similar sequence of BIF- and S-macrobands as seen in the Hamersley Iron Formation (Beukes 1983; Horstmann and Ha lbich 1995; Cheney 1996). The depositional age of the Kuruman and Griquatown Iron Formations is constrained by a U Pb zircon age of 2,521±3 Ma for a volcanic ash bed from the underlying Gamohaan Formation of the Campbellrand Subgroup (Sumner and Bowring 1996), as well as volcanic layers within the Kuruman Formation that have a U Pb zircon age of 2,432±31 Ma (Trendall et al. 1990). Development of iron formations reflects a dramatic departure from typical marine conditions. Sedimentary rocks of the Campbellrand Subgroup that underlie the Kuruman and Griquatown iron formations provide an

3 525 exceptionally detailed record of clastic and carbonate sedimentation prior to iron-formation deposition, including ferruginous carbonate turbidites and Ca Mg Fig. 1 Chemical and stable isotope (C and O) variations of the upper Gamohaan Formation, Kuruman Iron Formation and lower Griquatown Iron Formation. The stratigraphic section is interpreted to reflect a facies transition from shallow marine stromatolite environments to anoxic deep-water conditions (Beukes et al. 1990). The decrease in d 13 C values for carbonates that accompanies an increase in Fe-rich deposition may reflect greater influences by low-d 13 C organic material in the water column (Beukes et al. 1990). The generally lower d 18 O values measured for iron-formation carbonates as compared with those of high-ca carbonates is consistent with lithologic changes, although the lowest d 18 O values (which occur near a diabase sill in the upper sections of the AD-5 core) may reflect the effects of metamorphism (Kaufman 1996). The generally lower organic carbon contents and higher d 13 C values for kerogen in the iron-formations, as compared with the Gamohaan Formation, is consistent with distillation of CH 4 during early diagenesis, whereas the decrease in d 13 C values for kerogen in the Groenwater Member likely reflects contact metamorphism and CO 2 loss in proximity to the diabase sill. Fields for data from core WB-98 (upper Gamohaan to lower Kuruman Formations) also shown (scaled for differences in stratigraphic thicknesses). Samples used for Fe isotope analyses are shown by arrows on the right; samples from near the diabase sill in core AD-5 were avoided. Gray arrows indicate equivalent stratigraphic positions of samples taken from DW-19A core, which is adjacent to the AD-5 core. Similar symbols used in later figures, although units are somewhat generalized here. Data from Klein and Beukes (1989), Beukes et al. (1990), Beukes and Klein (1990), and Kaufman (1996) carbonates (Beukes 1984). Facies changes in the Campbellrand rocks are interpreted to reflect lateral changes in deposition, from shallow water Ca Mg carbonates to deeper Fe-carbonates, to iron oxide precipitation in the deepest sections, in an open shelf setting (Beukes 1984; Beukes et al. 1990). The occurrence of Fe-carbonate sequences in the transition from the Campbellrand to Kuruman sequences has been taken as evidence for a strongly stratified ocean shelf in terms of oxygen and ferrous Fe contents (Beukes et al. 1990; Ha lbich et al. 1992; Sumner 1997). For example, Fe(II) contents are interpreted to have varied from >770 lmol/l in the region where iron oxides were deposited, to 50 to several hundred lmol/l at depths of siderite deposition, to <20 lmol/l for Ca Mg carbonate precipitation (Sumner 1997). The transition zone between the extensive Ca Mg carbonate sequences of the Gamohaan Formation and the Kuruman Iron Formation is marked by the Tsineng and Kliphuis members of the Gamohaan and Kuruman Formations, respectively, and are characterized by interbedded shale and siderite-rich banded iron formation, with decreasing abundance of Ca Mg carbonate rocks and increasing Fe contents upward in the sequence (Fig. 1). The Kliphuis member is overlain by the magnetite siderite banded iron formations of the Groenwater member of the Kuruman Iron Formation (Fig. 1),

4 526 and records cyclic deposition of stilpnomelane lutite and magnetite hematite banded iron formation (Beukes 1983). These depositional cycles have been interpreted to reflect alternating deposition of volcanic ash and iron formation, respectively, similar to the cyclic deposition in the Hamersley Basin (Beukes 1983). Overlying the Groenwater member is the greenalite- and siderite-rich iron formation of the Riries member (Fig. 1), followed by the overlying Ouplass member that consists of clastictextured iron formation. In aggregate, the Kuruman Iron Formation is m thick (Beukes 1980). The Griquatown Iron Formation conformably overlies the Kuruman Iron Formation, has an average thickness of 250 m (Beukes 1980), and reflects a change in iron-formation deposition to include granular banded iron formation that is largely absent from the Kuruman Iron Formation. The granular iron formation of the main member, the Danielskuil member, is not oolitic, but instead consists of peloidal magnetite-rich iron formation, and is interpreted to represent an upwardshallowing sequence of storm-dominated deposits (Beukes and Klein 1990), reflecting deposition in an open continental shelf. Hematite is generally absent in the Griquatown Iron Formation, in contrast to the Groenwater member of the Kuruman Iron Formation. Geochemical variations Major and trace element compositions for the Kuruman and Griquatown iron formations are similar, suggesting that the source of metals was the same, despite their different depositional settings (Klein and Beukes 1989; Beukes and Klein 1990; Horstmann and Ha lbich 1995). Very little clastic component exists in the iron formations, as indicated by low Al and Sc contents, as well as low high-field-strength (Zr, Hf, Nb, Ta) or actinide (U and Th) element contents. Rare earth element (REE) patterns for the iron formations are typical of Archean and Early Proterozoic BIFs, characterized by high heavy REE/light REE ratios and positive Eu anomalies, which have been interpreted to largely reflect a mid-ocean ridge (MOR) hydrothermal source (Klein and Beukes 1989; Beukes and Klein 1990; Bau and Dulski 1996), and this is supported by modestly positive Nd values calculated at 2.5 Ga (Bau et al. 1997). Because Fe/Nd ratios of BIFs are similar to those of MOR hydrothermal fluids, most workers interpret the source of REEs to be similar to that of Fe (e.g., Jacobsen and Pimentel-Klose 1988). The transition from Ca Mg carbonate deposition to iron-formation deposition was accompanied by a decrease in d 13 C and d 18 O values for carbonate, a decrease in total organic carbon contents, and an increase in d 13 C values for organic carbon (Fig. 1). The marked decrease in d 13 C values for carbonates, from d 13 C=0 in the Gamohaan Formation to values as low as 13 in the Groenwater member Fe carbonates, is interpreted to reflect development of a stratified marine shelf or basin, where the d 13 C values of total dissolved CO 2 decreased downward toward the site of iron-formation deposition (Beukes et al. 1990). Beukes et al. (1990) interpret the origin of the low d 13 C values to largely reflect hydrothermal sources, consistent with trace element compositions (Klein and Beukes 1989), although equilibration of dissolved CO 2 with organic carbon, or generation of low d 13 CCO 2 from diagenetic reactions, may also have been factors. The low d 13 C values for the iron-formation carbonates cannot be explained by changes in the equilibrium isotope fractionation factor with mineralogy, in as much as D 13/12C siderite dolomite=0±1& over the temperature range of C (Sheppard and Schwarcz 1970; Carothers et al. 1988; Romanek et al. 1992). The modest lowering of d 18 O values from Ca Mg carbonate deposition to Fe-carbonate deposition may, however, reflect differences in O isotope fractionation factors or contrasts in early diagenetic histories. For example, at 25 C, D 18/16O dolomite siderite=+1.4 to +2.8& (Northrup and Clayton 1966; O Neil and Epstein 1966; O Neil et al. 1969; Sheppard and Schwarcz 1970; Carothers et al. 1988), which can largely explain the differences in d 18 O values that are observed for the carbonates (Fig. 1). The variations in the abundance and isotopic composition of total organic carbon upward in the sections can be largely explained by diagenetic processes, and in the case of the AD-5 core, contact metamorphic effects with proximity to a diabase sill at 87.5 to 92.5 m depth (Fig. 1). The d 13 C values of 35 to 40& for organic carbon in the Gamohaan Formation are interpreted to reflect autotrophic carbon fixation during primary production (Beukes et al. 1990; Strauss and Beukes 1996), which was thought to be greatest in the shallow stromatolite environments that are recorded in the Gamohaan Formation (Klein and Beukes 1989). The decrease in total organic carbon contents, accompanied by an increase in d 13 C values up through the Tsineng, Kliphuis, and the base of the Groenwater members may be explained by CH 4 loss during progressive kerogen maturation, low-grade metamorphism, or biogenic CH 4 production. For example, assuming D 13/12C methane carbon= 15 to 45&, which is appropriate for equilibrium fractionations between 50 and 200 C (Bottinga 1969), or fractionations inferred from field and laboratory studies of natural materials (Stahl 1977; Whiticar et al. 1986), the order of magnitude decrease in total organic carbon contents and 20& increase in d 13 C values can be explained by CH 4 loss; such correlations are found in sedimentary rocks throughout the geologic record (e.g., McKirdy and Powell 1974; Stahl 1977; Whiticar et al. 1986). The anti-correlation in total organic contents and d 13 C values in the Riries member of the Kuruman Iron Formation, as well as the Griquatown Iron Formation, is also consistent with CH 4 loss in those sections sampled by the CN-109 core (Fig. 1). Loss of CO 2 through diagenetic or metamorphic reactions, or kerogen breakdown, will change the isotopic compositions of residual kerogen and carbonates, through reactions such as:

5 527 3FeCO 3 þh 2 O! Fe 3 O 4 þ3co 2 þh 2 FeCO 3 þfe 2 O 3! Fe 3 O 4 þco 2 6Fe 2 O 3 þc! 4Fe 3 O 4 þco 2 (e.g., Perry et al. 1973; Kaufman 1996). In the case of CO 2 loss in the presence of kerogen, assuming D 13/12C carbon dioxide carbon=+12.7 to +14.4&, appropriate for temperatures between 100 and 300 C (Chacko et al. 1991), decreasing the d 13 C values of residual kerogen from 20 to 35 may be accomplished by 70% reaction and CO 2 loss; this seems to be a likely explanation for the decrease in d 13 C values in the upper part of the Groenwater member in the AD-5 core with increasing proximity to the sill at m depth (Fig. 1). In the case of CO 2 loss in the presence of carbonate, d 13 C and d 18 O values of residual siderite would increase between 1.5 and 4& and decrease between 5 and 6&, respectively, assuming a reaction of 50% and using the fractionation factors of Carothers et al. (1988) and Chacko et al. (1991); although such effects are difficult to determine from the scatter in d 13 C values for the upper part of the Groenwater member in the AD-5 core, the trends in d 18 O values support such an interpretation (Fig. 1). Such reactions may have occurred in the lower part of the Kuruman Iron Formation, but C isotope variations may have been masked by the effects of CH 4 loss (Fig. 1). Kaufman (1996) interprets mineralogical and stable isotope fractionations in the upper part of the Groenwater member in the AD-5 core to reflect the effects of contact metamorphism with proximity to the sill, where the largest effects have occurred within m of the sill ( m depth). Iron isotope geochemistry The field of Fe isotope geochemistry is relatively new, and here we briefly review the current understanding of Fe isotope variations in the Earth, as well as constraints on Fe isotope fractionation provided by theoretical calculations and experiments. Variations among the four isotopes of Fe ( 54 Fe, 56 Fe, 57 Fe, and 58 Fe) may be expressed in several ways. We favor using the d notation in units of per mil (&), following the accepted convention in light stable isotope geochemistry. Because 58 Fe abundances are quite small, Fe isotope variations are most usefully described using the 56 Fe/ 54 Fe and 57 Fe/ 54 Fe ratios, where the respective d values may be defined as: d 56 Fe ¼ and d 57 Fe ¼ 56 Fe= 54 Fe Sample 56 Fe= 54 Fe Whole Earth Fe= 54 Fe Sample 57 Fe= Fe Whole Earth ð1þ ð2þ ð3þ The whole-earth ratios are defined using the average Fe isotope composition of an extensive survey of igneous rocks from oceanic and continental settings (Beard et al. 2002a), where it was recognized, for example, that the 56 Fe/ 54 Fe ratio of igneous or whole- Earth Fe is homogeneous within the ±0.05& precision (1r) of the current state-of-the-art analytical methods. Because Fe isotope variations on Earth are expected to reflect only mass-dependent fractionations, variations in d 57 Fe values will be approximately 1.5 times those of d 56 Fe variations. We report both values for our analyses, which provides a critical internal check of data quality. We use the d 56 Fe value in our discussion, however, because it is defined using the two most abundant Fe isotopes and, therefore, has the highest precision. It is important to note that d values between laboratories may need to be adjusted to a common standard value, and some laboratories use the parameter, which defines Fe isotope variations in units of parts per 10,000 instead of per mil. The issue of standards is further discussed below in the Analytical methods section. Terrestrial igneous volcanic and plutonic rocks from a wide variety of tectonic settings, including ocean islands, mid-ocean ridges, and continental environments, spanning compositions from peridotite to rhyolite, have a homogeneous d 56 Fe value of 0.00±0.05& (1r; Fig. 2; Beard et al. 2002a). Moreover, modern and Cenozoic weathering products, including loess, aerosols, continental sediments, suspended river loads, and marine turbidites have d 56 Fe values that cluster about the igneous average, and are only slightly more variable than the variations seen in igneous rocks (Fig. 2; Beard et al. 2002b). The only significant exception to this is seen in some continental shales (Fig. 2). Significant Fe isotope variations are only seen in chemically precipitated sediments, including Pliocene to Recent Fe Mn crusts from the North Atlantic Ocean, and the Late Archean to Early Proterozoic BIFs of this study (Fig. 2). In the modern oceans, because eolian, aerosol, and riverine inputs have d 56 Fe values of zero, MOR hydrothermal fluids remain the only source for negative d 56 Fe values (Fig. 2; Sharma et al. 2001; Beard et al. 2002b). The relatively constant Fe isotope composition of most lithologic sources of Fe, at least in the modern (oxygenated) Earth, may reflect little Fe isotope fractionation between source Fe and mobilized Fe during weathering, transport, deposition and diagenesis (Beard et al. 2002b). In addition, because Fe is a conservative element in an oxygenated environment, with minimal net loss of Fe during weathering, the Fe isotope compositions of detrital material should reflect those of the source regions (Beard et al. 2002b). Distinction between these two scenarios will require development of an experimental database on equilibrium Fe isotope fractionation during weathering processes, in both oxic and anoxic environments. For the moment, we will take the variations illustrated in Fig. 2 to indicate that the source of detrital Fe for ancient sediments, in an oxygenated atmosphere, will have d 56 Fe values near zero, and that

6 528 b Fig. 2 Summary of Fe isotope variations in the Earth. d 56 Fe values are defined as the deviation in 56 Fe/ 54 Fe from the average for terrestrial igneous rocks, in units of per mil (&). Oceanic and continental igneous rocks have homogeneous d 56 Fe values within analytical error (±0.05&, 1r), and weathering products in the modern Earth (loess, aerosols, continental sediments, suspended river loads, and marine sediments) have d 56 F values that cluster closely about the igneous Fe average (vertical gray bar; width is ±0.05&). Some continental sediments, which include Proterozoic and Phanerozoic shales, have slightly more variable d 56 F values (H). In contrast, Fe Mn crusts from the modern Atlantic Ocean (Zhu et al. 2000) have negative d 56 F values, as do high-temperature Mid-Ocean Ridge (MOR) hydrothermal fluids (E; Beard et al. 2002b). Still greater ranges in d 56 F values are seen in the Late Archean to Early Proterozoic BIFs of this study (A D), which, in fact, span nearly the entire range yet observed for natural samples. Data for BIFs reflect calculated end-member mineral compositions (Table 2). Because Fe isotope variations are on the order of only a few per mil (&), only the highest precision MC-ICP-MS data are included in this summary, where d 56 Fe values can be determined to ±0.05& (1r; e.g., Zhu et al. 2001; Beard et al. 2002a). Data sources are Beard et al. (2002a, 2002b) E J, Zhu et al. (2000) E, and this study A D the only significant source for isotopically variable Fe would be Fe(II) from MOR hydrothermal fluids, which should have negative d 56 Fe values (Fig. 2). Isotopic variations between various minerals in chemically precipitated sediments such as BIFs (Fig. 2A D) may reflect isotopic variability in the source(s) of Fe and/or isotopic fractionation during precipitation or diagenetic reactions, either inorganically or related to biologic processing of Fe. Iron isotope fractionations on the order of many per mil are predicted at low temperatures between various Fe oxides and carbonates and dissolved Fe in solution, based on calculation of reduced partition function ratios (b factors) using spectroscopic data (Fig. 3; Polyakov and Mineev 2000; Schauble et al. 2001). In two studies, equilibrium Fe isotope fractionation factors can be inferred based on experimental determination of Fe(III) Fe(II) and Fe(III) hematite Fig. 3 Comparison of equilibrium Fe isotope fractionations that are predicted for inorganic mineral-fluid systems with those determined by experiment. Isotopic fractionations are defined as D 56/54Fe A B=d 56 Fe A d 56 Fe B, in units of per mil (&), following standard notation used for light stable isotopes. Predicted curves calculated from b factors given by Polyakov and Mineev (2000) and Schauble et al. (2001); b factors from Polyakov and Mineev (2000) adjusted to equivalent values for 56 Fe/ 54 Fe ratios. Very large mineral fluid Fe isotope fractionations are predicted at low temperatures, spanning a range of 10&. Experimental determination of equilibrium Fe isotope fractionations for Fe(III) hematite and Fe(III) Fe(II) do not support such large predicted fractionations. Measured Fe(III) hematite fractionation is 0.10& at 98 C (Skulan et al. 2002), which contrasts with the +2.8& fractionation that is predicted. Similarly, measured Fe(III) Fe(II) fractionation at 22 C is +2.7& (Johnson et al. 2002a), which contrasts with the +5.4& fractionation that is predicted (this can be seen in the figure by comparing predicted curves for Fe(II) Mt and Fe(III) Mt)

7 529 fractionation (Fig. 3; Johnson et al. 2002a; Skulan et al. 2002), and in both cases, the experimentally determined fractionation factors are significantly different than those predicted from spectroscopic data. For example, Fe isotope fractionation between hexaquo Fe(III) and hematite at 98 C is near zero, in contrast to a predicted fractionation of +2.8& (Fig. 3). In addition, isotopic fractionation between hexaquo Fe(III) and Fe(II) at 22 C has been measured to be +2.7&, which is approximately half that predicted from spectroscopic data (Fig. 3). Experimentally determined equilibrium Fe isotope fractionation factors are not yet available for magnetite or Fe-bearing carbonates in inorganic systems, although significant fractionations are predicted based on spectroscopic data (Fig. 3). Significant Fe isotope fractionations have been measured in additional experimental systems, including ion exchange columns (Anbar et al. 2000; Matthews et al. 2001), flow-through reactor systems (Bullen et al. 2001), and dissolution of silicates in the presence of organic ligands (Brantley et al. 2001). The dissolution experiments by Brantley et al. (2001) suggest that Fe dissolved from silicate minerals may have low 56 Fe/ 54 Fe ratios relative to the silicate component, when complexed with certain organic ligands, although a missing high 56 Fe/ 54 Fe component must have been produced in these experiments that is not yet identified. The experiments using ion exchange columns or flow-through reactor systems provide some useful information on the possible range of Fe isotope variations, although demonstration of attainment of isotopic equilibrium was not made for several of these experiments (Anbar et al. 2000; Bullen et al. 2001) and, in some cases, the data were interpreted to be dominated by kinetic isotope fractionations (Matthews et al. 2001). Experimental determination of Fe isotope fractionation during biological processing of Fe has been made in several studies. Ferrous Fe produced by reduction of ferric oxides and oxyhydroxides by Fe-reducing bacteria has 56 Fe/ 54 Fe ratios that are 1.3 per mil lower than that of the starting material (Beard et al. 1999, 2002a). In addition, ferric oxyhydroxide precipitates produced by anoxygenic, photosynthetic Fe-oxidizing bacteria are 1.3 to 1.5 per mil higher in 56 Fe/ 54 Fe than the starting Fe(II) solution (Croal et al. 2002). The similar isotopic fractionation between reduced and oxidized Fe in the biologic Fe-reduction and -oxidation experiments has been interpreted to reflect equilibrium Fe isotope exchange between soluble pools of ferric and ferrous Fe (Johnson et al. 2002b). The results of long-term (1.5 years) experiments involving dissimilatory Fe reduction of hydrous ferric oxide and production of Fe carbonates suggests that Fe(II) has 56 Fe/ 54 Fe ratios that are 1.3 to 2.2& higher than biogenically-produced Fe carbonates (Johnson et al. 2002b). In contrast, little isotopic fractionation is measured between Fe(II) and biogenic magnetite produced by dissimilatory reduction of hydrous ferrous oxide (Johnson et al. 2002b), and similar results were observed for Fe(III)- and Fe(II)-bearing systems during formation of intracellular magnetite in magnetotactic bacteria (Mandernack et al. 1999); reconciliation of these studies, which involved different Fe species in solution, remains an important issue. So far, it appears that Fe isotope fractionation is most significant for biologically processed Fe where redox changes occur, and when compared with equivalent inorganic systems, the isotopic fractionation factors are distinct in biologic systems; both Fe-reducing and Fe-oxidizing bacteria produce +1.3& fractionation between Fe(II) and ferric oxide and oxyhydroxides, as compared with +2.8& in the equivalent inorganic system (Table 1). This contrast between biologic and inorganic systems may be thought of as the Fe isotope vital effect for metabolically processed Fe (Johnson et al. 2002b). Sample selection and petrographic relations Because surface exposures of iron formations are commonly deeply weathered, samples for this study were selected from fresh drill core that are representative of the range in lithofacies reflected in the Kuruman and Griquatown iron formations that have been previously described and studied. Cores DW-19A (Beukes 1980) and AD-5 (Klein and Beukes 1989; Beukes et al. 1990; Kaufman 1996) are located near Pomfret, whereas core CN-109 (Beukes and Klein 1990) is located 150 km south, NW of the town of Kuruman. The stratigraphically lowest sample analyzed is a pyrite shale from the Tsineng member of the Gamohaan Formation (AD ), which is composed of interbedded shale and siderite-rich iron formation. Siderite-rich BIF was sampled in the Kliphuis member of the Kuruman Iron Formation, whereas oxide-rich BIF was sampled in the Groenwater and Riries members, as well as the Danielskuil member of the Griquatown Iron Formation. Samples in proximity to a diabase sill at 87.5 to 92.5 m depth in the AD-5 core were avoided due to the effects of contact metamorphism (Kaufman 1996). Most samples have whole-rock chemical data, mineral compositions, and C and O isotope data available from the previous studies of Klein and Beukes (1989), Beukes and Klein (1990), Beukes et al. (1990), and Kaufman (1996). Siderite-rich BIF samples consist of alternating bands of Fe-carbonate and chert, with variable amounts of organic carbon or kerogen that commonly contributes to the microbanded appearance (Klein and Beukes 1989). In general, oxide minerals are rare in the sideriterich BIF samples. In most samples, siderite consists of sub-mesoscopic microsparite, but coarse sparry Fe-carbonates are also found, ranging from ankerite to ferroan dolomite in composition. In a number of samples, coarse euhedral ankerite ferroan dolomite cuts across microbanding of fine-grained siderite, suggestive of early diagenetic crystallization (Klein and Beukes 1989). Siderite compositions average Fe 1.45 Mg 0.50 Ca 0.05 (CO 3 ) 2, whereas ankerite and ferroan dolomite average Ca 1.0- Fe 0.5 Mg 0.5 (CO 3 ) 2 (Klein and Beukes 1989).

8 530 Table 1 Summary of experimental determinations of biologic and inorganic equilibrium Fe isotope fractionation Species D 56/54Fe A B Notes and reference Inorganic systems 1. [Fe III (H 2 O) 6 ] 3+ [Fe II (H 2 O) 6 ] ±0.15& Ferric-ferrous equilibration in dilute aqueous solutions at 22 C. Johnson et al. (2002a) 2. [Fe III (H 2 O) 6 ] 3+ Fe 2 O ±0.20& Extrapolated equilibrium value from acid hydrolysis of dilute ferric nitrate solutions at 98 C; experimental runs up to 203 days. Skulan et al. (2002) 3. Fe 2 O 3 [Fe II (H 2 O) 6 ] ±0.20& Calculated at 22 C from above experiments Biologic systems 4. Fe II Fe 2 O 3 and Fe II HFO 1.3±0.1& Fractionation between ferrous Fe produced during dissimilatory Fe reduction of HFO by Fe-reducing bacteria and substrate of hematite and hydrous ferric oxide (HFO/ferrihydrite), in a wide variety of growth media 5. HFO Fe II +1.3±0.2& to +1.5±0.2& and over a wide range in reaction rates. Beard et al. (1999, 2002a) Fractionation between HFO and ferrous Fe, where ferric precipitates were produced by photosynthetic Fe-oxidizing bacteria in minimal salts growth media, over wide range in reaction rates. Croal et al. (2002) 6. Fe II Fe 3 O & Estimated equilibrium fractionation between ferrous Fe and magnetite produced during dissimilatory Fe reduction of HFO by Fe-reducing bacteria. Johnson et al. (2002b) 7. Fe 3 O 4 Fe III Cl 3 and Fe 3 O 4 Fe II SO 4 0.3& Fractionation between ferric or ferrous Fe in solution during formation of intracellular magnetite in magnetosomes in magnetotactic bacteria. Mandernack et al. (1999) 8. Fe II FeCO & Long-term (1.5 years) fractionation between ferrous Fe and siderite produced during dissimilatory Fe reduction of HFO by Fe-reducing bacteria. Johnson et al. (2002b). 9. Fe II (Fe,Ca)(CO 3 ) & Long-term (1.5 years) fractionation between ferrous Fe and Ca-bearing siderite produced during dissimilatory Fe reduction of HFO by Fe-reducing bacteria. Johnson et al. (2002b) Inorganic experimental studies only listed for those that demonstrated attainment of, or close approach to, isotopic equilibrium. The Fe isotope vital effect may be defined as the difference between equilibrium Fe isotope fractionation in an inorganic system, such as no. 3, and that measured in an equivalent biologic system, such as nos. 4 and 5 (Johnson et al. 2002b) Banding and sedimentary textures of the oxide-rich BIF are similar to those of the siderite-rich BIF. In the AD-5 and CN-109 cores, magnetite is the major oxide mineral (Klein and Beukes 1989; Beukes and Klein 1990), and most commonly occurs as coarse euhedral grains in both chert- and Fe-carbonate-bearing bands, and often cuts across sedimentary structures. Some magnetite also occurs as very fine-grained anhedral to subhedral grains that are intergrown with chert. Hematite is a minor component in most oxide-rich BIF layers in the AD-5 and CN-109 cores, although it is a significant oxide component in micro- and meso-bands in several samples from the DW-19A core. Hematite in the AD-5 and CN-109 cores occurs largely as fine dust in siderite or chert matrix, whereas hematite occurs both as fine dust and fine- to medium-grained anhedral crystals in oxide bands in the DW-19A core. In the AD-5 core above m depth, oxide-rich BIF contains abundant minnesotaite, where its abundance increases with proximity to the sill at 92.5-m depth (Klein and Beukes 1989; Kaufman 1996); only one minnesotaite-rich sample was studied here (sample in the AD-5 core). Determining the genesis of oxide and carbonate minerals in the BIFs samples is critical for interpreting the Fe isotope data. In the absence of secondary (supergene) alternation, or significant metamorphism, the most common oxide mineral in BIFs is generally magnetite, which many workers have argued is either a primary precipitate or early diagenetic re-crystallization product of fine-grained magnetite or ferric and ferrous hydroxide precursors (e.g., Klein 1974, 1983; Walker et al. 1983). Although primary hematite is generally rare in BIFs that have not undergone secondary alteration or metamorphism, hematite that occurs as fine-grained dust in carbonate and chert layers in the AD-5 and CN- 109 cores is interpreted to be one of the earliest oxide minerals (Beukes et al. 1990). In addition, we interpret the hematite-rich bands in the DW-19A core to reflect primary precipitation of hematite, or re-crystallization of ferric hydroxide precursors, as has been suggested for other iron formations (e.g., Klein and Bricker 1977; Ewers and Morris 1981). Fine-grained siderite that is disseminated throughout silica-rich samples is interpreted to be a primary phase, although euhedral ankerite and ferroan dolomite have been interpreted to reflect recrystallization of similar, though finer-grained precursors (e.g., Klein 1983). Coarse-grained, sparry ankerite, and ferroan dolomite are interpreted to reflect diagenetic re-crystallization. Petrographic relations in the hematite-rich BIF from the DW-19A core (Fig. 4) illustrate several possible genetic relations among Fe-bearing minerals in the samples studied here. The coarse-grained, euhedral magnetite layers, as well as fine-grained hematite layers, are interpreted to reflect primary or near-primary oxide precipitates. We would not expect the hematite and magnetite layers to be in Fe isotope equilibrium in such cases, reflecting oxidation/precipitation processes under

9 531 Fig. 4 Photomicrographs of sample DW-19A 3/74, showing petrographic relations between hematite, magnetite, and siderite. Hematite exists primarily as very fine-grained layers in a largely chert matrix, suggestive of primary precipitation. Magnetite is commonly coarsegrained and euhedral, and likely reflects recrystallization of precursor oxides in most cases, although it may also reflect reactions of hematite + Fe(II) to magnetite and siderite. A Large-scale features (scale bar is 5 mm total across). Box in A notes area of detail in reaction layer shown in B, which shows fine-grained (primary) hematite surrounded by apparent reaction zones of euhedral magnetite and siderite. Scale bar in B is 1 mm across. Both photographs taken in plane transmitted light different conditions. In contrast, layers that contain mixed oxide phases in the DW-19A core are characterized by clumps of euhedral magnetite, surrounded by an Fe-depleted (siderite-rich) zone (Fig. 4) that may indicate breakdown of primary hematite (in the presence of Fe(II) and carbonate ion) to secondary magnetite and siderite. It is possible that Fe isotope equilibrium was attained between magnetite and siderite during such diagenetic reactions. Analytical methods Core samples for Fe isotope analysis were micro-drilled or scratched with a W-C bit, generally sampling less than 1 mg of material. Digital photomicrographs were taken of all core and thin section samples, carefully correlated with the sampled area, and mineral abundances were estimated using image analysis methods and reflected and transmitted light microscopy. Samples were processed through ion-exchange chromatography to ensure that all potential isobars such as Ca and Cr were removed. The matrix was eluted using 7.0 M HCl, followed by collection of Fe in 0.5 M HCl. Yields are quantitative so that possible mass fractionation during ion-exchange chemistry has no effect on the data. Iron isotope compositions were measured using a Micromass IsoProbe, a multicollector inductively-coupled plasma mass spectrometer (MC-ICP- MS) with a magnetic sector. The hexapole collision cell in the IsoProbe completely eliminates 40 Ar 14 N and 40 Ar 16 O isobars, and nearly completely eliminates 40 Ar 16 OH, which is corrected using an on-peak-zero approach. Instrumental mass bias is done using a standard-sample-standard approach, and the integrity of the data is shown by the fact that the isotope ratios of 54 Fe, 56 Fe, and 57 Fe plot along mass-dependent fractionation lines (see Electronic Supplement). Between 100 and 200 ng of Fe are used for isotopic analysis with the IsoProbe, which permits a long-term precision (1- SD) of ±0.05& and ±0.07&, respectively, for d 56 Fe and d 57 Fe values, as determined by replicate analysis of samples and ultrapure standards. See Beard et al. (2002a) and Skulan et al. (2002) for additional details. We routinely measure three ultra-pure Fe standards; two are internal laboratory standards purchased from Johnson Mathey (UW J-M Fe) and High Purity Standards (UW HPS Fe), and the third standard is a certified isotope reference material, IRMM-014, available from the Institute for Reference Materials and Measurements (Taylor et al. 1992, 1993). Inter-laboratory comparison can be made by analysis of IRMM-014. The measured Fe isotope compositions of these three standards during the course of this study were UW J-M Fe d 56 Fe=+0.25±0.05& and d 57 Fe=+ 0.39±0.07& (1- SD n=47); UW HPS Fe d 56 Fe=+0.49±0.05& and d 57 Fe=+0.74±0.07& (1-SD n=52); and IRMM-014 d 56 Fe= 0.09 ±0.05& and d 57 Fe= 0.11±0.07& (1-SD n=54; see Electronic

10 532 Supplement). These values can be used to adjust data from other laboratories that report Fe isotope variations relative to the IRMM- 014 standard (e.g., Belshaw et al. 2000; Zhu et al. 2001), which has a slightly different Fe isotope composition than our whole-earth or igneous value. Also note that Fe isotope variations defined using the parameter, which is defined in parts per 10,000 (e.g., Zhu et al. 2001), will appear to be ten times larger in magnitude, although the analytical uncertainty will be an order of magnitude larger as well as compared to variations defined using the d parameter. Results Iron isotope compositions for individual bands in the Transvaal BIFs vary from d 56 Fe= 2.5 to +1.1&, almost spanning the entire range yet measured for natural Fig. 5 Measured whole-rock Fe isotope compositions for samples ( 1 mm thick) taken from layers in banded iron formations and related rocks, according to their stratigraphic position in cores CN-109, DW-19A, and AD-5. Members for formations shown: Da Danielskuil member; Ri Riries member; Gr Groenwater member; Kl Kliphuis member; Ts Tsineng member. Symbols indicate the dominant mineralogy for specific layers that were sampled. Depth numbers (in meters) given for each sample, and these correspond to those studied by Klein and Beukes (1989), Beukes et al. (1990), and Beukes and Klein (1990). Labels in boxes indicate rock type; Ox oxide-rich facies (primarily magnetite), H-Ox hematite-rich oxide facies; Carb carbonate-rich facies; Ox-Ank ankerite-bearing oxide-rich facies; Sh shale. Distinctive petrographic features are also noted adjacent to specific samples terrestrial samples (Figs. 2 and 5, Table 2). The lowest d 56 Fe values were determined for the pyrite-rich shale from the Gamohaan Formation (sample in the AD-5 core), where both pyrite veins and shale matrix have low d 56 Fe values. Iron carbonates from the sideriterich BIF samples of the Kliphuis member of the Kuruman Iron Formation, as well as the lower parts of the Groenwater member, consistently have higher d 56 Fe values between 1.0 and +0.1&, where the lowest values are associated with ankerite rhombs (e.g., sample in the AD-5 core). Low d 56 Fe values for Fe carbonates also tend to be associated with samples rich in organic carbon, such as sample in the AD-5 core. In many samples, oxide-rich layers have d 56 Fe values that are higher than those of Fe carbonates from the same hand sample, such as in samples (AD-5 core), 3/71 (DW-19A core), (CN-109 core), and (CN-109 core). This relation is particularly well illustrated for sample (AD-5 core), which was sampled in detail due to its excellent examples of siderite, ankerite, and magnetite. However, in two of the hematite-rich samples from the DW-19A core, the oxide-rich layers have d 56 Fe values that overlap those of Fe carbonates. Overall, hematite-rich layers have d 56 Fe values between 0.7 and +0.6&, and magnetite-rich layers have a similar range in d 56 Fe values, between 0.6 and +1.1&.

11 533 Table 2 Fe isotope compositions for samples from Transvaal banded iron formations Mineral modes (%) Core and depth Sample d 56 Fe MEAS d 57 Fe MEAS d 56 Fe MIN Mineral Mt Hem Si Sid Ank rhomb Fe-dol spar Pyrite Shale Griquatown Iron Fm. CN ± ± Sid <1 tr F ± ±0.04 CN CN Kuruman Iron Fm. CN J-51 Avg: F ± ± ± ± Sid <1 tr Avg: J ± ± ± ± Sid Avg: J ± ± Mt ± ±0.04 Avg: J ± ± Sid < J-47 a 0.21 ± ± Fe-Dol spar 3 tr 97 F-816 a 0.13 ± ± ± ± Fe-dol spar 3 tr ± ±0.04 Avg: ± ± Mt ± ±0.03 J-50 Avg: ± ± Mt ± ± ± ±0.03 Avg: J-52 Avg: DW-19A 3/74 R-11 (118 m) b Avg: R-12 Avg: R-14 Avg: R-15 Avg: AD ± ± Sid 11 tr ± ± ± ± Sid 3 tr ± ± ± ± Mt 97 tr ± ± ± ± Mt 87 tr ± ± ± ± Hem ± ± ± ± Fe-dol spar 1 tr R ± ±0.03 Avg: R ± ± Sid ± ±0.03 Avg: R ± ± Mt 87 tr 6 7 F ± ± ± ± Hem Avg: F ± ± Hem ± ±0.03 Avg: ± ± Sid ± ±0.03 J-28 Avg: J-29 Avg: DW-19A 3/71 R-7 (128 m) b Avg: ± ± Sid ± ± ± ± Mt ± ±0.03

12 534 Table 2 (Contd.) Mineral modes (%) Core and depth Sample d 56 Fe MEAS d 57 Fe MEAS d 56 Fe MIN Mineral Mt Hem Si Sid Ank rhomb Fe-dol spar Pyrite Shale DW-19A 3/38 (145 m) b R-1 AD R ±0.06 ± ±0.03 ± Sid <1 tr Avg: R ± ± Sid < ± ±0.02 Avg: R ± ± Mt 78 tr R ± ± ± ± Hem < Avg: F ± ± Hem ± ±0.03 Avg: ± ± Sid 11 tr ± ± ± ± ± ±0.05 Avg: R ± ± Mt 86 tr ± ± ± ± ± ±0.05 Avg: R-3 Avg: R-4 Avg: R ± ± Hem ± ± ± ± Mt 92 tr ± ± ± ± Hem ± ± ± ±0.06 Avg: R ± ± ± ± Sid < Avg: R ± ± Mt ± ±0.03 Avg: F ± ± Hem J ± ± ± ± Mt Avg: J ± ± Ank e < ± ±0.02 Avg: J ± ± Mt 95 <1 5 J ± ± ± ± Mt 99 <1 Avg: J ± ± Sid ± ±0.03 Avg: J ± ± Sid ± ± ± ±0.03 Avg: J-19 Avg: J ± ± Mt ± ± ± ± Mt ± ± ± ±0.03 Avg:

13 535 Table 2 (Contd.) Mineral modes (%) Core and depth Sample d 56 Fe MEAS d 57 Fe MEAS d 56 Fe MIN Mineral Mt Hem Si Sid Ank rhomb Fe-dol spar Pyrite Shale AD AD AD ± ± Sid J ± ± ± ±0.04 Avg: J ± ± Sid ± ± Sid < ± ±0.04 J-33 Avg: J-34 Avg: J ± ± Sid ± ± ± ± Sid ± ± ± ±0.03 Avg: c 0.16 ± ± Sid J ± ±0.03 Avg: J ± ± Sid < ± ±0.03 Avg: ± ± spar Fe-dol <1 7 tr 93 J ± ±0.03 Avg: c 0.23 ± ± Sid J ± ±0.04 Avg: d 0.29 ± ± Sid J ± ±0.04 Avg: J ± ± Sid < ± ± ± ±0.04 Avg: ± ± Sid < ± ±0.04 J-43 Avg: F-818 Avg: Gamohaan Fm. AD J ± ± Fe-dol spar ± ± ± ± Py ± ±0.04 Avg: F-817 Avg: F-875 Avg: J-24 Avg: ± ± Py ± ± ± ± Py ± ± ± ± Py ± ±0.04 a Samples J-47 and F-816 are duplicate samples (separate chemical processing) from same spot on core sample CN b For DW-19A core, equivalent depth in nearby AD-5 core noted c Samples J-38 and J-41 are duplicate samples (separate chemical processing) from same spot on core sample AD d Re-processed aliquot from original sample through separate chemistry e Calculated d 56 Fe MIN for ankerite is 2.95&, which appears too low; arbitrarily set to 2.00& d values in units of per mil (&), referenced to average value of terrestrial igneous rocks (see text). All duplicates are repeat mass analyses of same solutions on different days (under different running conditions), unless otherwise noted. Uncertainties for individual mass spectrometry runs noted (2-SE, internal statistics). d 56 Fe MIN is calculated for end-member composition for mineral noted in column to right, based on image analysis of mineral modes and Fe isotope composition of contaminant from nearest layer. Mt Magnetite; Hem hematite; Si silica (chert); Sid siderite; Ank ankerite; Ank Rhomb ankerite rhombs; Fe-Dol Spar sparry ankerite/ferroan dolomite; Py pyrite

14 536 Although sampling of individual layers attempted to isolate layers that were as close to mono-mineralogic as possible, oxide minerals occur in many of the carbonaterich layers and vice versa. Using image analysis on the same exact areas that were sampled to calculate mineral modes, we have estimated d 56 Fe values for pure mineral compositions in specific layers, using the Fe contents of the minerals determined by electron microprobe (Klein and Beukes 1989; Beukes and Klein 1990). This assumes that the contaminant component has the same Fe isotope compositions as adjacent layers that are enriched in that component (Table 1). No correction is made if the impurity is solely chert. For the vast majority of samples, this correction is 0.2&. Calculated end-member d 56 Fe values for minerals highlight the mineral-specific groupings in Fe isotope compositions for BIFs (Fig. 6). In all cases but one, d 56 Fe values calculated for magnetite are significantly higher than those calculated for Fe carbonate from the same sample. There is no correlation between d 56 Fe values for hematite and magnetite from the same sample, although the range in d 56 Fe values for hematite tends to be larger than that of magnetite in the same sample. Moreover, the d 56 Fe values for Fe carbonate generally decrease to more negative values when their end-member compositions are calculated, and the lowest values are estimated for ankerite rhombs. d 56 Fe values calculated for siderite in layers that also contain sparry ankerite/ferroan dolomite tend to be higher than those calculated for siderite where siderite is the only carbonate. The average d 56 Fe values for individual samples do not all cluster about zero, which might be expected if the source of Fe for the BIFs had the same isotopic composition as igneous rocks or modern detrital sediments, loess, and aerosols (Fig. 2). The shale sample appears to be uniformly low in its d 56 Fe value, and the siderite-rich BIFs have d 56 Fe values that largely lie between 0.0 and 0.5 (Fig. 6). Some of the oxide-rich BIFs appear to have average d 56 Fe values that lie near zero (i.e., AD , DW-19A 3/38, and DW-19A 3/71), but others do not (DW-19A 3/74 and several samples from the CN-109 core). There is some suggestion of an overall decrease in d 56 Fe values with increasing stratigraphic height for oxide minerals in the AD-5 and DW-19A cores, which lie within 5 km of each other, although this weak trend is not continued in the CN-109 core, which lies 150 km to the south. We infer from these observations that the isotopic composition of the source of Fe to individual BIF horizons was variable, for both carbonate and oxide components. In terms of late diagenesis or early metamorphic changes that may additionally obscure interpretations, oxide production accompanied by CO 2 loss [Reactions (1), (2), (3) above] are only evident in the C and O isotope compositions of the Kuruman BIF samples that occur in proximity of the diabase sill in the upper part of the Groenwater member in the AD-5 core (Fig. 1). We do not consider Reactions (1), (2), and (3) Fig. 6 d 56 Fe values calculated for pure mineral compositions based on modal abundances as determined by digital image analysis (Table 2). Samples from oxide-rich facies in stratigraphic order; carbonate-facies samples combined for clarity. Symbols and notations as in Fig. 5. Several consistent groupings become clear for the calculated end-member compositions, where, for a given layer, magnetite generally has d 56 Fe values that are higher than those of carbonate from the same core sample, and pyrite and pyrite-rich shale have the lowest d 56 Fe values. Inferred compositions for ankerite-rich layers have very low d 56 Fe values, lower than those of siderite. d 56 Fe values for hematite are quite variable. Carbonate samples that are rich in organic carbon tend to have lower d 56 Fe values as compared with carbonates from C-poor layers in the same core sample. d 56 Fe values calculated for ankerite are among the lowest measured in the section. The overall decrease in d 56 Fe values for magnetite with increasing stratigraphic height for core AD-5 and DW-19A may reflect an overall decrease in d 56 Fe values with time for the Fe source(s); d 56 Fe values for magnetite from stratigraphically higher samples in the CN-109 core are higher and do not continue this trend to have played a significant role in modifying the Fe isotope compositions of the samples studied here, because only one sample (121.3 in the AD-5 core) lies

15 537 within the zone of sill-related metamorphic effects, and this sample is a siderite-facies BIF that contains no oxide minerals. Inter-mineral Fe isotope fractionations Detailed O isotope studies of iron formations demonstrate that although isotopic fractionations between minerals broadly correlate with metamorphic temperatures (e.g., Perry 1983), isotopic fractionations may vary by several per mil for non- or weakly-metamorphosed iron formations (e.g., Becker and Clayton 1976; Kaufman et al. 1990; Yapp 1990; Winter and Knauth 1992; Kaufman 1996). For example, D 18/16O quartz siderite varies from 0 to+8& in iron formations that have only undergone diagenesis or weak metamorphism, suggesting that O isotope equilibrium was often not attained. Possible explanations may include changes in fluid composition during diachronous precipitation of chert and siderite, or during early diagenesis and fluid loss. We do not expect Fe isotope fractionation to follow the same isotopic disequilibrium recorded in O isotopes. The genesis of chert, of course, has no bearing on Fe isotope fractionations in iron formations. Similarly, post-depositional infiltration of dilute waters, although a potentially important process for O isotope variations, will have little impact on Fe isotope compositions in Ferich minerals. The most likely factor that will affect attainment of Fe isotope equilibrium among various minerals is whether or not Fe-bearing minerals from adjacent bands precipitated from a fluid that had the same Fe isotope composition. Many samples contain siderite and magnetite in varying proportions, and if we assume that adjacent siderite- and magnetite-rich bands precipitated from a fluid of similar Fe isotope composition (although distinct in, for example, carbonate-ion contents), the Fe isotope compositions of siderite and magnetite may reflect nearequilibrium fractionations. Sparry ankerite and ferroan dolomite also appear to be diagenetic in origin, although these late-stage minerals may not have formed in equilibrium with siderite and magnetite. Following this approach, we calculated D 56/54Fe magnetite ankerite 0.8& (Fig. 7), and obtained similar results for D 56/54Fe magnetite siderite, assuming that these minerals generally formed in isotopic equilibrium. Using the maximum regional metamorphic temperatures of 110 to 170 C (Miyano and Beukes 1984), we can assume that the d 56 Fe values for magnetite and carbonate may reflect Fe isotope fractionation between temperatures of early diagenesis and low-grade metamorphism, say between 25 and 170 C. These results strongly contrast with the calculated Fe isotope fractionation factors of Polyakov and Mineev (2000), which suggest that temperatures between 25 and 170 C produce D 56/54Fe magnetite ankerite and D 56/ 54Fe magnetite siderite between +7.7 to +3.6 and +6.2 to +2.9&, respectively. This is far greater than the range measured in our study (Fig. 7). Although the relative Fig. 7 Comparison of Fe isotope fractionations between magnetite and hematite, ankerite, and siderite calculated for pure minerals from BIF samples studied here (histograms in lower part, B) with those calculated based on Mo ssbauer shifts (Polyakov and Mineev 2000; upper part, A). Note the change in scale for D 56/54Fe Mt X values between the upper and lower parts of the figures. In general, D 56/54Fe Mt X values measured in the BIF samples follow those predicted by Polyakov and Mineev (2000), in the order magnetite ankerite >magnetite siderite >magnetite hematite, although the measured isotopic differences are significantly less than those predicted if the BIF samples reflect attainment of isotopic equilibrium at temperatures of C or less (Miyano and Beukes 1984). The D 56/54Fe Mt X values less than 0.6 all belong to pairs that involve magnetite that has d 56 Fe<0, which may reflect equilibration with an unusual fluid. Note, however, that pyrite is predicted to have d 56 Fe values similar to those of magnetite, which is not observed in our samples. Magnetite is predicted to have d 56 Fe values, which are significantly higher than those of hematite, and is also not observed order of decreasing d 56 Fe values we measure of magnetite >siderite >ankerite/ferroan dolomite is the same as that predicted by Polyakov and Mineev (2000), the magnitude of the predicted fractionations appears to be

16 538 too great by a factor of 4 to 10; this discrepancy would seem too large to be explained by disequilibrium Fe isotope fractionation. The relative d 56 Fe values of magnetite, hematite, and pyrite may be loosely constrained from our results, although there is no geologic evidence, for example, that magnetite and pyrite formed in equilibrium with each other, and previous work has interpreted the genesis of hematite and magnetite to be quite different. Nevertheless, the calculations of Polyakov and Mineev (2000) predict significant Fe isotope fractionation between magnetite and hematite, where D 56/54Fe magnetite hematite= +3.2 to +1.5& between 25 and 170 C (Fig. 7), and our results do not support such large differences in Fe isotope compositions. Furthermore, Polyakov and Mineev (2000) predict that pyrite and magnetite should have very similar d 56 Fe values, and yet in the BIF samples studied here, magnetite has the highest positive d 56 Fe values and pyrite and pyrite-bearing samples have the most negative d 56 Fe values measured. Our results highlight the importance of experimentally calibrating Fe isotope fractionations among common low-temperature minerals if Fe isotope geochemistry is to become a useful avenue of research. We can refine the b factors predicted by Polyakov and Mineev (2000) and Schauble et al. (2001) for hexaquo Fe(III), hematite, magnetite, siderite, and ankerite through new experimental constraints in inorganic systems, and assuming that magnetite, siderite, and ankerite formed diagenetically in Fe isotope equilibrium. We take the b factor for hexaquo Fe(II) reported by Schauble et al. (2001) to be correct because it is consistent with experimental data (Skulan et al. 2002). Although the uncertainties in revised b factors for Fe carbonates remain relatively large, due to the spread in isotope compositions measured for the minerals in the BIF samples, we believe they are closer to the true values than those predicted by Polyakov and Mineev (2000), despite the uncertainty that the isotopic contrasts reflect equilibrium isotope fractionations. Comparison of our preferred set of b factors and those calculated by Polyakov and Mineev (2000) and Schauble et al. (2001) is given in Table 3. Graphs of the two sets of b factors, as well as a discussion of the process by which these were calculated, is presented in the Electronic Supplementary Material. Fig. 8 Calculated Fe isotope compositions of fluids in equilibrium with minerals from BIF samples of this study. Ranges shown encompass the range in d 56 Fe values measured for specific samples, as well as the effect of a range in temperatures from 25 to 100 C. d 56 Fe values for fluids are for Fe(II) species (hexaquo complexes), except for hematite, where d 56 Fe values are calculated for Fe(III) species (hexaquo complexes). In A, d 56 Fe values for fluid compositions are calculated using Fe isotope fractionation factors from Polyakov and Mineev (2000) and Schauble et al. (2001). The large disagreement, by many per mil, in calculated d 56 Fe values for fluids suggests that either the predicted Fe isotope fractionations are not correct by many per mil, precipitation of carbonate and oxide minerals occurred from distinct fluids, or isotopic equilibrium was not attained. In B, d 56 Fe values for fluid compositions are calculated using our preferred set of Fe isotope fractionation factors (Table 3). Ranges in compositions calculated for B additionally include estimated uncertainties in the preferred fractionation factors. The relative self-consistency of the calculated d 56 Fe values for fluids is imposed by incorporation of the BIF data in developing the preferred fractionation factors, although the absolute d 56 Fe values are not; the absolute d 56 Fe values for Fe(II) are fundamentally constrained by the b factor for Fe(II) from Schauble et al. (2001). d 56 Fe values calculated for paleo fluids are best constrained for Fe(III) using hematite (Skulan et al. 2002) Fe isotope variations in ancient fluids The key to inferring the Fe isotope compositions of ancient fluids from the rock record lies in accurate mineral-fluid isotope fractionation factors. Using the calculated reduced partition function ratios, or b factors, for minerals (Polyakov and Mineev 2000) and hexaquo species (Schauble et al. 2001), we can calculate the d 56 Fe values of Fe(II) or Fe(III) that might have been in equilibrium with specific minerals, assuming a temperature range, for example, between 25 and 100 C (Fig. 8). If minerals that might have formed in isotopic equilibrium, such as magnetite and siderite, do not produce the same calculated fluid d 56 Fe value, then either the assumption of isotopic equilibrium is incorrect or the isotope fractionation factors are in error. The widely variant d 56 Fe values calculated for dissolved Fe(II) based on magnetite, siderite, and ankerite (Fig. 8A) suggests that significant errors exist in the predicted Fe isotope fractionations, which in turn indicate that the predicted b factors are imprecise. It is particularly striking that use of the b factors proposed by Polyakov and Mineev (2000) and Schauble et al. (2001) produces a c Table 3 Summary of predicted and preferred Fe isotope fractionation factors Temp. Fe II Fe III Fe II siderite Fe II ankerite Fe II magnetite Fe III magnetite Fe III hematite Source ( C) & +2.0& +3.5& 4.2& +1.3& +4.4& Polyakov and Mineev (2000); Schauble et al. (2001) 2.6& 1.7 to +0.3& +0.3 to +1.8& 2.5 to 2.4& +0.1 to +0.2& 0.2& Preferred set (this study) & +1.3& +2.3& 2.7& +0.8& +2.9& Polyakov and Mineev (2000); Schauble et al. (2001) 1.7& 1.1 to +0.2& +0.2 to +1.2& 1.6 to 1.5& 0.0 to +0.1& 0.1& Preferred set (this study) Fe isotope fractionation given as D 56/54Fe A B, in per mil. All Fe species in solution are hexaquo complexes. See Data repository for detailed discussion of how preferred set of fractionation factors were derived

17 539 very large range in d 56 Fe values for aqueous solutions that have not been found in natural Fe isotope variations (Fig. 2). It is also important to note that the very positive d 56 Fe values that are calculated for hexaquo Fe(III) based on hematite (Fig. 8A) also appear to be excessive, based on the restricted range measured in natural samples. Using our preferred set of b factors for hexaquo Fe(II), Fe(III), hematite, magnetite, siderite, and ankerite (Table 3), we calculate d 56 Fe values for Fe(II) and Fe(III) in the solutions from which the BIFs were formed (Fig. 8B) that are significantly different than those calculated using the b factors from Polyakov and Mineev (2000) and Schauble et al. (2001). It is important to note however, that although the relative agreement in d 56 Fe values calculated for hexaquo Fe(II) based on siderite, ankerite, and magnetite in part reflects the fact that the BIF data themselves were used to refine the b factors (see Electronic Supplement), the absolute d 56 Fe values that are calculated do not directly depend on the approach used to refine the b factors, but instead fundamentally reflect our assumption that the hexaquo Fe(II) b factor of Schauble et al. (2001) is correct. Although the uncertainties in the preferred set of b factors remains significant, the calculations are suggestive that the Fe(II) species that were associated with BIF formation had d 56 Fe values that were less than zero, consistent with a MOR-hydrothermal source (Sharma et al. 2001; Beard et al. 2002b). In contrast, the preferred b factor for hematite is well constrained (relative to Fe(III)) by experimental data (Skulan et al. 2002), and this permits a relatively precise estimate for the d 56 Fe values of hexaquo Fe(III) from which primary hematite may have precipitated, which varies between +0.5 and 0.5& (Fig. 8B). Implications for the origin of banded iron formations and ancient geochemical cycling of iron The Fe isotope variations measured in BIFs, combined with experimental and theoretical constraints on isoto-

18 540 pic fractionation in biologic and abiologic systems allow us to place limits on the sources of Fe for BIFs and the potential pathways in which redox cycling of Fe occurs in their formation. These limits will be further refined as additional experimental data on Fe isotope fractionation become available, which will allow us to better address uncertainties that remain in regard to diagenesis and fluid-mineral fractionation factors. Effect of oceanic residence time The large contrast in solubility of ferrous and ferric aqueous species predicts markedly different geochemical cycling in the Earth under anoxic and oxygen-bearing conditions. The isotopic effects of such contrasting behavior can be well illustrated by the contrasting residence time that is predicted for Fe in the oceans under anoxic and relatively oxic conditions and the concomitant ability of Fe isotope compositions to respond to changes in the isotopic compositions of various inputs (Fig. 9). The very low Fe contents in the modern (oxic) Fig. 9 Effect of residence time of Fe on the ability to change Fe isotope compositions of the oceans, in both anoxic and relatively oxic environments. Initial d 56 Fe value of the oceans is assumed to be 0.6&, the lowest value inferred for the Late Cenozoic oceans by Zhu et al. (2000). Changes in isotope compositions reflect an instantaneous change to d 56 Fe=0& for the isotopic composition of the influx. The model is calculated using standard flux models and residence time definitions (e.g., Hodell et al. 1989; Berner and Berner 1996). Very short residence times are calculated for Fe in oxygen-bearing ocean masses, such as is characteristic of the modern oceans, due to the extremely low Fe contents of seawater, from 1.5 (without mid-ocean ridge hydrothermal flux) to 45 years (mid-ocean ridge hydrothermal flux included). In contrast, the very high Fe contents that are estimated for ocean masses that are anoxic and rich in ferrous Fe implies a long residence time of 1.5 million years. Residence times calculated assuming an oxygenated ocean that contains 50 ppt Fe (e.g., Rue and Bruland 1995), and a reduced ocean that contains 50 ppm Fe (e.g., Ewers 1983; Sumner 1997), a dissolved riverine Fe flux of g/year (see compilation by Berner and Berner 1996), and a mid-ocean ridge Fe flux of g/year (Stein and Stein 1995). These calculations indicate the dramatically different isotopic responses that may occur for Fe on the Earth under different redox conditions, presumably reflecting contrasting conditions of nearsurface free oxygen contents oceans produces a very short residence time for Fe, varying between 10 0 and 10 2 years, resulting in an extreme sensitivity in the isotope composition of seawater Fe to changes in the isotopic composition of the major sources for Fe. For example, the 0.6& increase in d 56 Fe values that has been inferred to have occurred in seawater over the last 2 million years from studies of Fe Mn crusts (Zhu et al. 2000) could be accomplished on the order of 10 1 to 10 2 years using the very short residence time of Fe in the modern oceans (Fig. 9). Such changes may reflect the relative balance of hydrothermal Fe input, which appears to have negative d 56 Fe values, and surface sources of Fe, such as riverine, eolian, and aerosol sources, which have d 56 Fe values near zero (Fig. 2; Beard et al. 2002b). A general expectation, therefore, is that under conditions where Fe has a short residence time, it is likely that the Fe isotope composition of minerals precipitated from seawater would be quite variable over short time intervals. We further expect that the d 56 Fe values for chemically precipitated oxide minerals will largely reflect the Fe isotope composition of Fe(III) because chemical reactions will tend to go to completion and, therefore, any equilibrium or kinetic isotope fractionation that may occur during precipitation will have minimal influence on the overall measured d 56 Fe values for oxide minerals. The long residence time for Fe that is inferred for anoxic conditions, possibly on the order of 10 6 years in the Archean if Fe(II) contents were quite high, makes it quite difficult to rapidly change the Fe isotope compositions of ancient, anoxic ocean water masses (Fig. 9). For example, a 0.6& change in the d 56 Fe values of the Fe inputs would take 10 6 to 10 7 years to be fully reflected in the Fe isotope composition of an Fe-rich ocean mass (Fig. 9). We would therefore expect that under anoxic conditions (local or global), the Fe isotope compositions of minerals that precipitated from seawater would be relatively homogenous over time intervals of several million years, assuming that temperature effects on Fe isotope fractionation factors are not large, and that diagenetic influences were minor. Although the Fe isotope composition of Fe-rich, oxygen-poor fluids may be relatively resistant to changes from changing fluxes or removal of Fe through precipitation, small changes in the Fe isotope fractionation factors between fluid and mineral may produce large changes in the d 56 Fe values of chemically precipitated minerals if they comprise a minor fraction of the mass of the system. Moreover, isotopic exchange during diagenesis, when the system may be relatively isolated from the main mass of dissolved Fe, may produce significant isotopic variability under such conditions. We anticipate that calculation of the d 56 Fe values of the ancient oceans using minerals that precipitated under anoxic, ferrous-rich conditions will be relatively difficult due to the present uncertainty in mineral-fluid isotope fractionation factors for Fe, which become significant issues when (1) Fe precipitation reactions may not go to completion under anoxic

19 541 conditions, and (2) the mass balance of the system may lie largely in the fluid component. Nevertheless, once the essential mineral-fluid fractionation factors are known, Fe isotope data will potentially place important constraints on the limits of isotopic variability in ancient environments. This in turn will place constraints on the magnitude and origin of possible changes in the isotopic compositions of the major Fe sources in near-surface environments. The 0.5& overall decrease in d 56 Fe values for BIF samples in the Groenwater member, from sample AD to DW-19A 3/74 (equivalent to 118 m depth in the nearby AD-5 core; Figs. 5 and 6) places some constraints on the residence time of Fe during deposition of the main BIF sequence. Uncompacted sedimentation rates for the Kuruman Iron Formation have been estimated to be between 20 and 60 m/million years (Altermann and Nelson 1998), suggesting that the 34-m interval sampled between AD and DW-19A 3/74 reflects 0.6 to 1.7 million years of sedimentation. Assuming a change in the d 56 Fe values for the Fe flux on the order of 0.5 to 1.0&, the response time for the BIF assemblage would be consistent with an Fe residence time on the order of 10 6 years (Fig. 9), suggesting involvement of an ocean mass that was relatively anoxic and rich in Fe(II). Geochemical pathways for Fe during BIF formation We illustrate several geochemical pathways for Fe that may have occurred in the upper water column, as well as those that may have occurred in the deep basin during primary precipitation or diagenesis (Fig. 10). The BIF facies model of Beukes et al. (1990) envisions the upper water column to be somewhat oxygen-bearing (although significantly less than the modern oceans) and Fe-poor, whereas the deeper parts of the basin are thought to have been anoxic and Fe-rich, and this model provides a useful framework for interpretation of the Fe isotope data. We can constrain the isotope compositions of several Fe fluxes, as discussed above, where d 56 Fe= 0.6 to 0.3& for MOR hydrothermal fluids, and riverine, eolian, and aerosol input had d 56 Fe=0.0±0.07, based on observations in the modern Earth (Fig. 1). If the atmosphere was oxygen-poor at the time of BIF formation, we would expect the d 56 Fe values of the suspended river and atmospheric fluxes to be close to those of modern-day values, but the dissolved riverine and atmospheric fluxes might have had negative d 56 Fe values if significant ferrous Fe existed in surface environments (Beard et al. 2002b). We assume that the d 56 Fe values for the hydrothermal component that is measured today would be similar to that of the Archean because MOR hydrothermal activities in the Late Archean to Early Proterozoic seem likely to have been similar to what they are today, although a significantly reduced Archean ocean may negate this assumption due to the large Fe(III) Fe(II) isotope fractionation (Johnson et al. 2002a). Several pathways for producing primary hematite or ferric oxyhydroxide are possible in the upper water column. One involves abiotic precipitation of hematite (or its hydroxide precursors) in somewhat oxygenated, Fe-poor upper portions of the water column (where soluble Fe is largely ferric), followed by settling into the deeper anoxic parts of the water column (Beukes et al. 1990). If hematite formed in this manner is precipitated abiotically from dilute Fe(III)-bearing waters, then the d 56 Fe values of hematite should reflect those of the Fe(III) in the water (Fig. 3). In case 1 (Fig. 10), the d 56 Fe values of Fe(III) are assumed to be in isotopic equilibrium with deep Fe(II) across the chemocline; because essentially all Fe resides in the Fe(II)-rich zone, the d 56 Fe values of Fe(III) will remain fixed by the abiologic Fe(III) Fe(II) fractionation factor in case 1, and will not be influenced by the riverine and eolian fluxes. Case 1 predicts very high d 56 Fe values for hematite, which are not observed in the Kuruman Griquatown BIF samples studied here. In contrast, if hematite has abiotically precipitated from the upper water column where upwelling reduced Fe is oxidized as it comes in contact with oxygen-rich surface waters (case 2), without maintenance of isotopic equilibrium across the chemocline, the d 56 Fe values of Fe(III) will reflect those of the Fe(II) sources, mitigated by the d 56 Fe values of the riverine and eolian fluxes. In this scenario, d 56 Fe values of hematite will be moderately negative to zero, which is observed in several BIF samples in this study; however, moderately positive d 56 Fe values for hematite cannot be explained by this model, although such values might be produced if partial Fe(III) Fe(II) equilibration occurred. Note that if the d 56 Fe values of surface fluxes were slightly negative, as might be the case in an oxygen-poor atmosphere (Beard et al. 2002b), the negative d 56 Fe values for hematite measured here may still be explained by this model. Another possibility reflects precipitation of hematite in the upper water column by oxidation of Fe(II) by photosynthetic Fe-oxidizing bacteria, as has been proposed by Widdel et al. (1993; case 3). Based on the +1.3 to +1.5& fractionation measured between ferric oxyhydroxide precipitates and Fe(II) during photosynthetic Fe oxidation reported by Croal et al. (2002), this scenario would predict d 56 Fe values for Fe(III) between zero and +1.3, depending upon the relative fluxes of Fe(III) from bacterial processing and those of riverine and eolian inputs. Such moderately positive d 56 Fe values for hematite are observed in the BIF samples studied here. It is important to note that cases 1 and 2 imply greater oxygen contents in the upper water column (although perhaps much less than today), whereas case 3 could occur in a relatively anoxic upper water column. The mixed valance state of Fe in magnetite and the wide variety of Fe pathways in which it may form makes interpretation of the Fe isotope variations we measure for magnetite relatively complex. If, for example, magnetite formed through precipitation of Fe(OH) 2 (s) and

20 542 Fe(OH) 3 (s) precursors (e.g., Klein 1983), possibly from a mixed valance solution, we would expect the ferrous and ferric solid and solution components to have low and high d 56 Fe values, respectively, based on experimental studies (e.g., Bullen et al. 2001; Johnson et al. 2002a, 2002b; Skulan et al. 2002). The final d 56 Fe value for magnetite would then reflect the relative proportions of ferrous and ferric species in the solution and solid components of the system, as well as the associated Fe isotope fractionation factors. In addition, adsorption of ferrous and ferric Fe species onto magnetite (e.g., Haderlein and Pecher 1999) may impart possible isotopic effects. Unfortunately, the detailed pathways by which magnetite may form and react with a fluid are not yet understood in terms of Fe isotope fractionations. It is striking that the d 56 Fe values for magnetite and hematite from the same samples are generally similar, suggesting that, regardless of the differences in Fe pathways during formation of these minerals, both oxides may largely record the isotopic compositions of the major sources of Fe that are eventually sequestered into oxide minerals in BIFs. Mineral formation in the deep portions of the basin may have occurred by a number of processes, including primary precipitation, diagenesis, and microbially mediated precipitation. In case 4 (Fig. 10), all minerals are assumed to have precipitated inorganically; hematite is assumed to have precipitated from the upper portions of the water column, whereas magnetite, Fe-carbonate, and siderite are envisioned to have precipitated in the deep basin, presumably under relatively anoxic conditions; this essentially follows the BIF facies model proposed by Beukes et al. (1990). Using the preferred fractionation factors noted in Table 3, the predicted d 56 Fe values for magnetite, Fe-carbonate, and siderite fall within the range of those measured in this study, although the very high d 56 Fe values that are predicted for siderite, and particularly magnetite, are not observed. Note that although the preferred fractionations in Table 3 are based on an assumed isotopic equilibrium between magnetite and Fe carbonates, the absolute d 56 Fe values have no such assumption and fundamentally rely on the predicted b factor for Fe(II) from Schauble et al. (2001).