Constraints from Iron Isotope Variations in Sedimentary Rocks from the Kaapvaal and Pilbara Cratons

Chemical Geology 218 (2005) 135–169 www.elsevier.com/locate/chemgeo

Biogeochemical cycling of iron in the Archean–Paleoproterozoic Earth: Constraints from iron isotope variations in sedimentary rocks from the Kaapvaal and Pilbara Cratons

Kosei E. Yamaguchia,b,c, Clark M. Johnsonb,c,T, Brian L. Beardb,c, Hiroshi Ohmotoc,d

aInstitute for Research on Earth Evolution (IFREE), Japan Agency for Marine-Earth Science and Technology (JAMSTEC), 2-15 Natsushima, Yokosuka, 237-0061, Japan bDepartment of Geology and Geophysics, University of Wisconsin - Madison, 1215 W. Dayton St., Madison, WI 53706, USA cNASA Astrobiology Institute, United States dAstrobiology Research Center and Department of Geosciences, The Pennsylvania State University, 435 Deike Building, University Park, PA 16802, USA Received 4 April 2004; received in revised form 21 October 2004; accepted 26 January 2005

Abstract

Iron isotope compositions of low-metamorphic grade samples of Archean–Paleoproterozoic sedimentary rocks obtained from fresh drill core from the Kaapvaal Craton in South Africa and from the Pilbara Craton in Australia vary by ~3x in 56Fe/54Fe ratios, reflecting a variety of weathering and diagenetic processes. Depositional ages for the 120 samples studied range from 3.3 to 2.2 Ga, and Fe, C, and S contents define several compositional groups, including samples rich in Fe, organic carbon, carbonate, and sulfide. 56 The d Fe values for low-Corg, low-Ccarb, and low-S sedimentary rocks are close to 0x, the average of igneous rocks. This range is essentially the same as that of Corg-poor late Cenozoic loess, aerosol, river loads, and marine sediments and those of 56 Corg-poor Phanerozoic–Proterozoic shales. That these d Fe values are the same as those of igneous rocks suggests that Fe has behaved conservatively in bulk sediments during sedimentary transport, diagenesis, and lithification since the Archean. These observations indicate that, if atmospheric O2 contents rose dramatically between 2.4 and 2.2 Ga, as proposed by many workers, such a rise did not produce a significant change in the bulk Fe budget of the terrestrial sedimentary system. If the Archean atmosphere was anoxic and Fe was lost from bedrock during soil formation, any isotopic fractionation between aqueous ferrous 2+ Fe (Feaq ) and Fe-bearing minerals must have been negligible. In contrast, if the Archean atmosphere was oxic, Fe would have been retained as Fe3+ hydroxides during weathering as it is today, which would produce minimal net isotopic fractionation in bulk detrital sediments. 56 2+ Siderite-rich samples have d Fe values of À0.5F0.5x, and experimentally determined Feaq -siderite fractionation factors 2+ 56 56 2+ suggest that these rocks formed from Feaq that had similar or slightly higher d Fe values. The d Fe values calculated for Feaq

T Corresponding author. Department of Geology and Geophysics, University of Wisconsin - Madison, 1215 W. Dayton St., Madison, WI 53706, USA. Tel.: +1 608 262 1710, +1 608 265 6798; fax: +1 608 262 0693. E-mail address: [email protected] (C.M. Johnson).

2+ 56 overlaps those of modern submarine hydrothermal fluids, but it is also possible that Feaq had d Fe values higher than those of 2+ modern hydrothermal fluids, depending upon the Feaq –Fe carbonate fractionation factor that is used. In contrast, Corg-rich samples and magnetite-rich samples have strongly negative d56Fe values, generally between À2.3x and À1.0x, and available fluid–mineral fractionation factors suggest that the Fe-bearing minerals siderite and magnetite in these rocks formed in the 2+ 56 3+ presence of Feaq that had very low d Fe values, between À3x and À1x. Reduction of Fe hydroxide by sulfide, precipitation of sulfide minerals, or incongruent dissolution of silicate minerals are considered unlikely means to produce 56 2+ 3+ significant quantities of low-d Fe Feaq . We interpret microbial dissimilatory Fe reduction (DIR) as the best explanation for 56 2+ producing such low d Fe values for Feaq , and our results suggest that DIR was a significant form of respiration since at least 2.9 Ga. D 2005 Elsevier B.V. All rights reserved.

Keywords: Iron; Isotope; Archean; Proterozoic; Biology

1. Introduction by a variety of inorganic and organic redox reactions, many of which produce significant isotopic fractiona- Evolution of the Archean biosphere and its tions. In the case of Fe, biological and nonbiological influence on redox conditions remains a first-order isotopic fractionations are well demonstrated to occur problem in studies on the evolution of the Earth. in nature and in experiments (e.g., Beard and Johnson, Evidence from molecular phylogeny suggests that 1999, 2004a,b; Beard et al., 1999, 2003a,b; Anbar, photosynthesis evolved quite early in Earth’s history 2004; Anbar et al., 2000; Zhu et al., 2000, 2001, 2002; (Xiong et al., 2000). Biomarker evidence also suggests Brantley et al., 2001, 2004; Bullen et al., 2001; that organisms related to cyanobacteria (oxygenic Matthews et al., 2001, 2004; Sharma et al., 2001; photosynthesizers) evolved by at least 2.7 Ga (Brocks Johnson et al., 2002, 2003, 2004, 2005; Severmann et et al., 1999, 2003). It has also been proposed, on the al., 2004a; Levasseur et al., 2004; Skulan et al., 2002; basis of molecular phylogeny, that dissimilatory Fe3+ Welch et al., 2003; Kehm et al., 2003; Roe et al., 2003; reduction (DIR) may have been one of the earliest Rouxel et al., 2003, 2004; Icopini et al., 2004). forms of respiration on Earth (Vargas et al., 1998), Because some of the largest isotopic fractionations preceding other important respiratory processes, such occur during redox changes (Polyakov and Mineev, as sulfate reduction, nitrate reduction, and oxygen 2000; Schauble et al., 2001; Johnson et al., 2002; reduction (Lovley, 1993). Although such a view is Welch et al., 2003; Anbar et al., 2005), Fe isotopes consistent with known phylogenies of extant organ- hold particular promise in tracing biologically Fe2+ isms, it is impossible to validate these hypotheses oxidation and Fe3+ reduction (Johnson et al., 2004). without corroborating geochemical evidence that Despite great interests in Fe isotope geochemistry, little provides a timeline for the process (Benner et al., work has been done on Archean sedimentary rocks. 2002). The possibility that DIR occurred early in In this contribution, we examine the evidence for Earth’s history is particularly important because of its redox cycling of Fe in the Archean. The dataset implications for the chemical and redox evolution of include low organic-C (Corg) clastic rocks that likely the early Earth. If little Fe3+ existed on the early Earth, reflect terrestrial weathering processes, as well as an Fe3+-reductase would not have functioned because rocks that probably formed in relatively anoxic marine there would be no advantage in assimilatory and environments. A major theme we develop centers on dissimilatory metabolism of Fe (or other metals) to the evidence that dissimilatory Fe3+ reduction (DIR) sustain life (Nealson and Saffarini, 1994). may have been an active metabolic pathway by at Specific metabolisms may produce unique isotopic least 2.9 Ga. One hundred and twenty samples of fractionations of essential elements such as C, H, N, carbonaceous shales, non-carbonaceous shales, mag- and S, which may be detected in the rock record (e.g., netite-rich shales, siderite-rich shales, and greywackes Hayes, 2001; Canfield, 2001). The biogeochemical were selected for this study using drill cores from 10 cycles of Fe, like those of C, H, N, and S, are mediated drill holes representing 14 geologic formations recov- K.E. Yamaguchi et al. / Chemical Geology 218 (2005) 135–169 137 ered in South Africa and Australia. They are 3.25 to significant metamorphism, generally less than greens- 2.20 Ga in age, have undergone minimal metamor- chist facies. South African samples are from the ~3.3 phism, and are free from the effects of modern Ga Fig Tree Group of the Swaziland Supergroup, the 3+ 2+ weathering. We report data for Fe ,Fe ,Corg, and ~3.0 Ga West Rand Group of the Witwatersrand Ccarb (carbonate carbon), and S contents and Fe Supergroup, the ~2.7 Ga Platberg Group of the isotope compositions. Ventersdorp Supergroup, the ~2.6 Ga Wolkberg Group, the ~2.6 Ga Chuniespoort Group, and the ~2.2 Ga Pretoria Group of the Transvaal Supergroup 2. Geological settings and samples (Fig. 3). Australian samples are from the ~2.7 Ga Fortescue Group and the ~2.6 Ga Hamersley Group Archean and Paleoproterozoic strata are exception- of the Mt. Bruce Supergroup (Fig. 4). ally well preserved and exposed on the Kaapvaal Craton of southern Africa (Fig. 1) and on the Pilbara 2.1. Swaziland Supergroup Craton in the Pilbara–Hamersley regions of NW Australia (Fig. 2). Sedimentary successions uncom- The Swaziland Supergroup is exposed in the 3.4 to formably overlie the cratonic basements. We have 3.1 Ga Barberton Greenstone Belt located in the analyzed 103 laminated and nonlaminated black and eastern part of the Kaapvaal Craton (Fig. 1). The greenish shale samples, 9 greywacke samples, and 8 Swaziland Supergroup consists of three groups: from laminated red shale samples (in 14 formations) from lower to upper, the 3.48–3.45 Ga Onverwacht Group, drill cores, which are free of weathering products (120 3.33–3.23 Ga Fig Tree Group, and 3.22–3.10 Ga samples in total). The rocks have not been subject to Moodies Group (Armstrong et al., 1990; Kro¨ner et al.,

Fig. 1. Simplified geological map of South Africa showing the distribution of the Transvaal, Ventersdorp, and the Witwatersrand Supergroups, the Sabie-Pilgrim’s Rest region, the Griqualand West region, the Witwatersrand Basin, and the Barberton Greenstone Belt, which includes the Swaziland Supergroup. Modified after SACS (1980). 138 K.E. Yamaguchi et al. / Chemical Geology 218 (2005) 135–169

Fig. 2. Simplified geological map of the Pilbara–Hamersley district, western Australia, showing the distribution of the Turee Creek, Hamersley, and Fortescue Groups of the Mt. Bruce Supergroup. Drill core samples analyzed in this study were taken from near Wittenoom in the central Hamersley Basin and near Ripon Hill in the eastern Hamersley region. Modified after Geological Survey of Western Australia (1990).

1991; Kamo and Davies, 1994; Lowe and Byerly, The black shale samples of this study are partly 1999). The Onverwacht Group is predominantly from the 3.25 Ga Sheba Formation of the Fig Tree composed of mafic and ultramafic rocks that contain Group (Fig. 3). Sixteen shale samples were collected minor sedimentary and felsic volcanic units. The Fig from a ~90 m long drill core, PU1308, at the Agnes Tree Group is composed of a succession of grey- Mine (Wagener, 1986) in the Barberton Greenstone wacke, shale, chert, dacitic flow, and fragmented Belt (Fig. 3). Nine greywacke samples were also volcanic rocks. The lowermost Sheba Formation of collected from the 3.25 Ga Sheba Formation, using the Fig Tree Group is composed mainly of coarse, drill core MRE10 at the Sheba Mine (Wagener and immature turbiditic sandstone and thin interbedded Wiegand, 1986) in the same district. The sedimentary units of siltstone and shale (Lowe and Byerly, 1999). textures are well preserved except in extremely The Moodies Group is composed of feldspathic and sheared rocks; in general, these rocks were subjected quartzose sandstone, chert-clast, and siltstone. Depo- to very minor shearing and only relatively low-grade sitional settings of the sedimentary rocks of the Fig regional metamorphism (greenschist facies) (Viljoen Tree and Moodies Groups are interpreted to have been and Viljoen, 1969). a foreland or foredeep basin (Jackson et al., 1987; de Wit et al., 1992; Heubeck and Lowe, 1994), or an 2.2. Witwatersrand Supergroup evolving back-arc or passive continental margin that formed adjacent to early rifts or continental shelf and The Witwatersrand Basin extends about 200 km shelf-rise environments (Windley, 1995). SW from the Barberton area and has an areal extent of K.E. Yamaguchi et al. / Chemical Geology 218 (2005) 135–169 139

Fig. 3. Lithostratigraphic column of the Transvaal, Ventersdorp, Witwatersrand, and Swaziland Supergroups in South Africa. Vertical thickness is not to scale. Samples were taken from shaded Groups. Modified after Lowe and Byerly (1999) (geochronological data and references therein).

~80,000 km2 (Fig. 1). The Witwatersrand Supergroup Group: 2914F8 Ma; Armstrong et al., 1991) and a conformably overlies volcanic rocks of the Dominion predominantly arenaceous upper phase (Central Rand Group and uncomformably overlies the Archean Group). The arenaceous rocks include several hori- basement rocks (greenstone-granitoid complexes) of zons of quartz-pebble conglomerate units that host Au the Kaapvaal Craton. The Witwatersrand Supergroup and U mineralization (dreefsT). Shallow marine or tidal was deposited between 2970 and 2714 Ma (Robb et marine with minor alluvial depositional environments al., 1990). Recent reviews on the geology of the have been suggested as the depositional setting of the Witwatersrand Basin have been provided by Coward West Rand Group (Tankard et al., 1982). et al. (1995) and Robb and Meyer (1995). The samples of this study are from the 2.96 Ga The Witwatersrand Supergroup is dominated by Parktown Formation of the Hospital Hill Subgroup of siliclastic deposits that are divided into a mixed the West Rand Group and the ~2.9 Ga Rietkuil argillaceous–arenaceous lower phase (West Rand Formation of the West Rand Group (Fig. 3). Fifteen 140 K.E. Yamaguchi et al. / Chemical Geology 218 (2005) 135–169

Fig. 4. Lithostratigraphic column of the lower Hamersley and upper Fortescue Groups of the Mt. Bruce Supergroup, western Australia. Scale bar is 500 m. Samples were taken from shaded Formations (Wittenoom Dolomite, Marra Mamba Iron, Carawine Dolomite, and Lewin Shale Formations). Correlation of Carawine Dolomite and underlying Lewin Shale in the eastern region to the Hamersley/Fortescue Groups in the central region has not been firmly established; it is suggested that the Carawine Dolomite and Lewin Formations are equivalent to the Wittenoom Dolomite and Jeerinah Formations, respectively (e.g., Simonson and Hassler, 1997). Modified after Geological Survey of Western Australia (1990) and Simonson and Hassler (1997). shale samples from the 2.96 Ga Parktown Formation Ma (Armstrong et al., 1991), conformably overlies the were collected from a ~120 m section of a N500 m Witwatersrand Supergroup and the Dominion Group long drill core, DRH13, recovered in the NW margin of the Kaapvaal Craton. The Ventersdorp Supergroup of the Witwatersrand Basin (Fig. 1). Another set of 15 is mainly composed of continental flood basalt and shale samples from the ~2.9 Ga Rietkuil Formation minor sediments (Winter, 1976), and divided into, were collected from the drill core MGM5. A number from lower to upper, the Klipriviersberg, Platberg, and of samples from the ~2.9 Ga Rietkuil Formation have Pniel Groups. The Klipriviersberg and Platberg sufficiently high abundances of magnetite, which are Groups are largely composed of volcanic rocks, and detectable with a hand magnet, and locally the rocks the Pniel Group is composed of clastic sedimentary are referred to as bmagnetic shalesQ. These rocks were and volcanic rocks. The interpreted depositional only subjected to low-grade regional metamorphism setting for the sediments of the Pniel Group is an (greenschist facies; e.g., Pretorius, 1981; Phillips et alluvial plain (Tankard et al., 1982). al., 1989; Wronkiewicz and Condie, 1990). Seven shale samples of the ~2.6 Ga Rietgat Formation of the Platberg Group were collected from 2.3. Ventersdorp Supergroup a ~800 m section of the ~4.4 km long drill core MSF6 recovered in the NW margin of the Witwatersrand The Ventersdorp Supergroup (e.g., Button, 1981a), Basin (Fig. 1). The drill core MSF6 spans from lower which have age constraints of 2709F4 and 2714F8 to upper, the Florida Quartzite, Livingstone Conglom- K.E. Yamaguchi et al. / Chemical Geology 218 (2005) 135–169 141 erate, Kimberley Conglomerate, Elsburg Quartzite, collected from a 200 m section of the ~4.4 km long Kameelsdoorn, Goedgenoeg, and Rietgat Formations drill core MSF6. Both of the drill cores (JPBR and from the Ventersdorp Supergroup, and the Black Reef MSF6) were recovered in the NW margin of the Quartzite, Oak Tree, Monte Christo, Lyttelton, Eccles, Witwatersrand Basin (Fig. 1). Twelve samples of the Rooihoget, Timeball Hill, and lower Hekpoort Ande- 2.35 Ga Timeball Hill Formation were collected from site Formations from the Transvaal Supergroup (see a 100 m section of the ~400 m long drill core PTB3 in section below). These rocks were not subjected to the Pilgrim’s Rest-Sabie region of the eastern Trans- deformation, but were metamorphosed to greenschist vaal (Fig. 1). The samples have not been deformed facies (e.g., Pretorius, 1981; Phillips et al., 1989; and were only subject to low-grade regional meta- Wronkiewicz and Condie, 1990). morphism (greenschist facies; e.g., Pretorius, 1981; Phillips et al., 1989; Wronkiewicz and Condie, 1990). 2.4. Transvaal Supergroup Several black and red shales were analyzed in this study. The drill core SA1677 was recovered from the The Transvaal Supergroup (e.g., Button, 1981b), Sishen iron mine, in the Griqualand West region/ mainly composed of sedimentary rocks, uncomform- Northern Cape Province of South Africa (Fig. 1), and ably overlies the Ventersdorp Supergroup. The this encompasses red shales of the Gamagara For- Transvaal Group is divided into, from lower to mation (SACS, 1980) or the Mapedi Shale Formation upper, the Wolkberg, Chuniespoort, and Pretoria of the Olifantshoek Group (Beukes and Smit, 1987), Groups (Fig. 3). although some uncertainty exist in correlation to The Wolkberg Group (2630–2560 Ma: Jahn et al., exposed stratigraphic sequences. The ~250 m long 1990; Armstrong et al., 1991) is composed of fluvial drill core SA1677 was sampled at depths between 200 and shallow marine feldspathic quartzite, arkoses, and 245 m, and this section contains fine-grained red siltstones, pelites, carbonate, and conglomerates, as shales and minor red sandstones, as well as conglom- well as subaerially to subaqueously extruded basalt, erate and breccia in the lowermost part of the and probably was deposited in a rift basin (Tyler, examined section. Some of the red shales alternate 1978; So¨hnge, 1986; Button, 1986). The Chunies- with minor white (bleached) layers of variable thick- poort Group (2557F49 Ma: Jahn et al., 1990)is ness (millimeter–centimeter). Eight red shale samples mostly composed of stromatolitic carbonates and iron- were analyzed. As with the other rocks analyzed, formations, and probably deposited in shallow marine these samples were subjected to low-grade regional or intertidal environments based on the presence of metamorphism (greenschist facies). stromatolitic carbonates (Truswell and Eriksson, 1975; Eriksson, 1983; Button, 1986; So¨hnge, 1986; 2.5. Mt. Bruce Supergroup Cledenin et al., 1988). The Pretoria Group (2.3–2.0 Ga) is mainly composed of quartzites and shales with Archean and Paleoproterozoic strata are also minor volcanic rocks and carbonates, and was exceptionally well preserved and exposed on the probably deposited in shallow marine cratonic envi- Pilbara Craton in the Pilbara–Hamersley regions of ronments (Button, 1981b; Tankard et al., 1982). It has NW Australia (Fig. 2). In the southern part of the been suggested that the Kaapvaal Craton was tectoni- Pilbara Craton, the Neoarchean Mt. Bruce Super- cally stable during deposition of the Chuniespoort group uncomformably overlies the cratonic base- Group (So¨hnge, 1986) and the lower section of the ment (e.g., Hickman, 1983). The Mt. Bruce Pretoria Group (Schreiber et al., 1992). Supergroup is the infill of the Hamersley Basin Samples were taken from the 2.64 Ga Black Reef that extends over an area of 106 km2 (Fig. 2). The Formation and the 2.56 Ga Oak Tree Formation of the Mt. Bruce Supergroup is divided into, from lower to Chuniespoort Group, and the 2.35 Ga Timeball Hill upper, the Fortescue, Hamersley, and Turee Creek Formation of the Pretoria Group (Fig. 3). Four Groups. The Fortescue Group (e.g., Thorne and samples of the Black Reef Formation were collected Blake, 1990; Arndt et al., 1991) is mainly com- from a 110 m section of a N2500 m long drill core, posed of mafic volcanic rocks (basalt and komatiite) JPBR. Three samples of the Oak Tree Formation were and volcaniclastic rocks, subordinate silicic volcanic 142 K.E. Yamaguchi et al. / Chemical Geology 218 (2005) 135–169 rocks (rhyolite), conglomerate, sandstone, shale, and sulfur isotope studies (Bottomley et al., 1992). All of carbonates (Fig. 4). The Hamersley Group (e.g., the Australian samples of this study are undeformed Thorne, 1990) conformably overlies the Fortescue and were subjected to relatively low-grade regional Group and is well known for its extensive iron metamorphism (greenschist facies; Hickman, 1983; formations. Other than iron formations, the Hamers- Bottomley et al., 1992; Simonson et al., 1993, 1998; ley Group contains dolomite, carbonaceous shale, Simonson and Hassler, 1997; Woodhead et al., 1998; and rhyolite. The Hamersley Group is divided into, Brocks et al., 1999; Lindsay and Braiser, 2002). from lower to upper, the Marra Mamba Iron Formation, Wittenoom Dolomite Formation, Mt. Sylvia Formation, Mt. McRae Formation, Brockman 3. Analytical methods Iron Formation, Weeli Wolli Formation, Woongarra Volcanics Formation, and Boolgeeda Iron Formation Detailed analytical methods other than Fe isotope (Fig. 4). The Hamersley Group in turn is overlain analysis are given in Yamaguchi (2002) and are only by the Turee Creek Group, and is largely composed briefly described here. All drill core samples were of greywacke, sandstone, siltstone, quartzite, and carefully washed with distilled water, dried, and dolomite. examined visually. Samples that did not contain The samples used in this study are from the 2.72 veins or mineralization were crushed into coarse Ga Pillingini Tuff Formation of the middle Fortescue fragments (5–10 mm) using a jaw-crusher, which in Group, the 2.69 Ga Jeerinah Formation of the upper- turn were picked to remove fragments that contained most Fortescue Group and the N2.60 Ga Marra the original drill core surface. Picked fragments were Mamba Iron Formation and 2.60 Ga Wittenoom further washed with distilled deionized water (18 Dolomite Formation of the Hamersley Group (Fig. MV) in an ultrasonic bath several times to remove 4). Three samples from the Pillingini Tuff Formation, attached dust, dried, and powdered by agate mortar six samples from the Jeerinah Formation, four and pestle at Pennsylvania State University (PSU) samples from the Marra Mamba Iron Formation, and and Ocean Research Institute (ORI) at the University three samples from the Wittenoom Dolomite For- of Tokyo. Between each sample, quartz sand was mation were collected from 200 m, 50 m, 70 m, and 5 processed to minimize cross-contamination between m sections, respectively, of a N1500 m long drill core, samples, and duplicate control samples were also WRL1, recovered near the town of Wittenoom, processed. Mineralogical and textural characteristics western Australia (Fig. 2). Drill core WRL1 is the of the samples were examined under petrographic same as that used to successfully extract Archean microscope using thin sections in both transmitted cyanobacterial biomarkers from the Jeerinah Forma- and reflected light, as well as XRD (X-ray diffrac- tion (Brocks et al., 1999). The carbonate rocks in this tion) at PSU (Rigaku Geigerflex) using powdered core have been studied for C and O isotope variations samples. (Lindsay and Braiser, 2002). Total Fe contents were measured by ICP-AES The Lewin Shale Formation and Carawine Dolo- (inductively coupled plasma-atomic emission spectro- mite Formation occur in the eastern region of the scopy; Leeman Labs PS 3000UV) after decomposi- Hamersley Basin, and are thought to be stratigraph- tion of the powdered samples in solution by standard ically equivalent to the Jeerinah and Wittenoom alkali fusion methods using lithium metaborate at Dolomite, respectively, which occur in the central- PSU and ACTLAB, Inc., and by XRF (X-ray western region of the Hamersley Basin (Fig. 2; fluorescence) at ORI using standard methods. Fe2+ Simonson et al., 1993, 1998; Simonson and Hassler, contents were measured by titration using a modifi- 1997; Woodhead et al., 1998). Eight samples from the cation to Wilson’s method (Wilson, 1955; Yang and Lewin Shale and six samples from the Carawine Holland, 2003). Fe3+ contents were determined by Dolomite Formation were collected from a 500 m difference between total Fe and Fe2+ contents. long drill core, RHDH2A, recovered near the Ripon Procedural blanks and several SRM (standard refer- Hills in the eastern Hamersley region (Fig. 2). The ence materials; SCO1, MAG1, BCR2, W2, JB3, and drill core RHDH2A has been previously used for JG1A) were used to monitor data quality. The data K.E. Yamaguchi et al. / Chemical Geology 218 (2005) 135–169 143

Table 1 Geochemical data for shales and greywackes obtained in this study P Sample ID Individual analyses Grand average Fe2O3 Fe2O3 FeO Corg Ccarb S d56Fe 2S.E. d57Fe 2S.E. d56Fe 1rd57Fe 1r (wt.%) (wt.%) (wt.%) (wt.%) (wt.%) (wt.%) Kaapvaal Craton Transvaal Supergroup Pretoria Group ~2.2 Ga Gamagara Formation SA-07 0.25 0.08 0.42 0.04 0.25 0.08 0.42 0.04 10.91 8.04 2.58 0.10 0.00 b.d. SA-08 À0.03 0.08 0.03 0.04 À0.03 0.08 0.03 0.04 10.76 5.97 4.31 0.08 0.09 b.d. SA-10 0.14 0.04 0.18 0.03 0.14 0.04 0.18 0.03 9.52 7.98 1.39 0.09 0.00 b.d. SA-11 0.14 0.05 0.23 0.03 0.18 0.06 0.24 0.02 10.24 9.25 0.89 0.07 0.00 b.d. 0.22 0.06 0.26 0.03 SA-12 0.18 0.06 0.31 0.04 0.13 0.07 0.23 0.10 13.26 12.42 0.76 0.07 0.00 b.d. 0.09 0.05 0.16 0.04 SA-15 0.09 0.05 0.11 0.02 0.10 0.01 0.13 0.04 18.53 17.69 0.76 0.06 0.20 b.d. 0.10 0.06 0.16 0.04 SA-16R 0.03 0.05 0.06 0.03 À0.01 0.06 0.02 0.06 14.75 14.19 0.50 0.06 0.72 b.d. À0.06 0.04 À0.03 0.03 SA-17 À0.16 0.07 À0.22 0.03 À0.16 0.07 À0.22 0.03 18.42 17.49 0.84 0.06 0.00 b.d. 2.35 Ga Timeball Hill Formation PTB-07 0.05 0.06 0.07 0.03 0.06 0.01 0.05 0.03 7.44 1.26 5.56 0.41 0.00 b.d. 0.06 0.06 0.03 0.04 PTB-08 0.25 0.06 0.33 0.03 0.25 0.06 0.33 0.03 4.85 0.63 3.80 0.35 0.88 0.32 PTB-11 À0.01 0.09 À0.04 0.04 0.01 0.03 À0.04 0.01 8.85 0.14 7.84 0.46 0.00 b.d. 0.03 0.06 À0.03 0.03 PTB-13 0.01 0.08 0.10 0.04 0.01 0.08 0.10 0.04 8.74 0.36 7.54 0.39 0.00 b.d. PTB-16 0.10 0.06 0.25 0.03 0.10 0.06 0.25 0.03 10.18 0.57 8.65 0.39 0.00 b.d. PTB-18 0.03 0.06 À0.01 0.03 0.03 0.01 0.05 0.08 9.42 1.34 7.27 0.39 0.40 0.06 0.02 0.10 0.10 0.04 PTB-19 À0.49 0.06 À0.64 0.05 À0.49 0.06 À0.64 0.05 9.96 1.39 7.71 0.36 0.28 b.d. PTB-20 0.11 0.05 0.25 0.03 0.11 0.05 0.25 0.03 2.87 0.04 2.55 0.50 0.38 0.42 PTB-22 0.10 0.06 0.10 0.04 0.10 0.06 0.10 0.04 9.03 0.71 7.49 0.43 0.01 b.d. PTB-23 0.06 0.05 0.15 0.03 0.15 0.08 0.22 0.09 10.63 1.53 8.19 0.47 0.45 b.d. 0.21 0.05 0.32 0.03 0.18 0.06 0.19 0.04 PTB-25 0.00 0.09 0.02 0.04 0.00 0.09 0.02 0.04 8.61 1.09 6.77 0.43 0.00 b.d. PTB-26 À0.01 0.09 0.06 0.04 À0.01 0.09 0.06 0.04 9.15 1.54 6.85 0.30 0.45 b.d. Chuniespoort Group 2.56 Ga Oak Tree Formation MSF-115 0.27 0.06 0.41 0.03 0.27 0.06 0.41 0.03 1.99 0.10 1.70 2.30 0.15 0.27 MSF-117 0.64 0.07 1.08 0.04 0.64 0.07 1.08 0.04 2.60 0.08 2.27 1.64 0.05 0.94 MSF-119 0.57 0.06 0.87 0.03 0.57 0.06 0.87 0.03 1.49 0.09 1.26 1.66 0.00 b.d. Ventersdorp Supergroup Wolkberg Group 2.64 Ga Black Reef Formation JPBR-1 À0.40 0.08 À0.61 0.05 À0.40 0.08 À0.61 0.05 0.76 0.65 0.10 0.09 12.02 b.d. JPBR-4 À0.47 0.06 À0.67 0.04 À0.47 0.06 À0.67 0.04 1.18 0.09 0.98 0.09 11.71 0.64 JPBR-9 À0.64 0.04 À0.93 0.04 À0.64 0.04 À0.93 0.04 5.12 0.18 4.45 0.33 11.83 0.04 JPBR-14 0.27 0.06 0.37 0.03 0.27 0.06 0.37 0.03 5.16 1.37 3.41 1.11 0.07 b.d. Platberg Group 2.71 Ga Rietgat Formation MSF-146 0.06 0.05 0.10 0.03 0.02 0.04 0.05 0.07 4.61 0.17 4.00 0.32 0.24 2.48 À0.01 0.07 0.00 0.03 MSF-152 À0.06 0.05 À0.13 0.03 À0.06 0.05 À0.13 0.03 6.48 0.32 5.54 0.27 0.30 b.d. (continued on next page) 144 K.E. Yamaguchi et al. / Chemical Geology 218 (2005) 135–169

Table 1 (continued) P Sample ID Individual analyses Grand average Fe2O3 Fe2O3 FeO Corg Ccarb S d56Fe 2S.E. d57Fe 2S.E. d56Fe 1rd57Fe 1r (wt.%) (wt.%) (wt.%) (wt.%) (wt.%) (wt.%) Kaapvaal Craton Ventersdorp Supergroup Platberg Group 2.71 Ga Rietgat Formation MSF-154 À0.04 0.06 À0.01 0.04 À0.01 0.03 0.01 0.03 6.46 0.58 5.29 0.59 0.50 b.d. 0.01 0.05 0.00 0.04 0.01 0.08 0.05 0.04 MSF-158 À0.71 0.07 À0.95 0.06 À0.71 0.07 À0.95 0.06 6.78 0.53 5.63 0.95 0.00 0.64 MSF-159 0.01 0.06 0.01 0.04 0.01 0.06 0.01 0.04 7.94 0.79 6.43 1.30 0.03 0.38 MSF-162 0.00 0.06 0.03 0.03 0.00 0.06 0.03 0.03 8.18 0.72 6.71 1.59 0.00 0.67 MSF-167 À0.08 0.07 À0.11 0.03 À0.08 0.00 À0.10 0.02 9.61 0.72 8.00 1.10 0.04 0.53 À0.09 0.07 À0.09 0.05 Witwatersrand Supergroup West Rand Group ~2.9 Ga Rietkuil Formation MGM-26 À1.36 0.05 À1.95 0.03 À1.36 0.05 À1.95 0.03 28.65 6.21 20.20 0.17 0.02 0.80 MGM-27 À1.30 0.05 À1.92 0.03 À1.31 0.03 À1.92 0.00 29.44 5.20 21.81 0.10 0.55 0.19 À1.33 0.07 À1.93 0.04 MGM-28 À1.10 0.06 À1.70 0.04 À1.10 0.06 À1.70 0.04 26.16 2.88 20.96 0.16 0.00 b.d. MGM-29 À1.25 0.04 À1.88 0.02 À1.25 0.04 À1.88 0.02 27.33 8.44 17.00 0.19 2.07 0.11 MGM-30 À1.21 0.06 À1.78 0.04 À1.21 0.06 À1.78 0.04 33.24 7.72 22.97 0.21 0.00 b.d. MGM-31 À1.31 0.06 À1.95 0.04 À1.31 0.01 À1.93 0.02 23.77 3.50 18.24 0.15 1.14 b.d. À1.30 0.05 À1.91 0.04 MGM-39 À1.33 0.06 À1.98 0.04 À1.33 0.01 À1.96 0.03 32.58 10.00 20.32 0.19 0.33 b.d. À1.34 0.06 À1.93 0.04 MGM-42 À1.16 0.05 À1.68 0.03 À1.16 0.05 À1.68 0.03 32.71 10.60 19.90 0.18 0.63 0.42 MGM-43 À1.03 0.05 À1.48 0.03 À1.03 0.05 À1.48 0.03 32.83 9.48 21.02 0.18 0.21 b.d. MGM-45 À0.96 0.04 À1.40 0.04 À0.96 0.04 À1.40 0.04 30.02 1.50 25.68 0.17 0.01 b.d. MGM-48 À1.16 0.06 À1.70 0.03 À1.16 0.06 À1.70 0.03 34.09 5.74 25.51 0.19 0.05 b.d. MGM-51 À1.08 0.04 À1.61 0.02 À1.07 0.01 À1.61 0.01 29.45 5.10 21.91 0.07 0.45 0.09 À1.07 0.05 À1.61 0.04 MGM-53 À1.13 0.04 À1.61 0.03 À1.13 0.04 À1.61 0.03 27.65 2.81 22.35 0.19 0.00 b.d. MGM-54 À1.11 0.05 À1.64 0.02 À1.11 0.05 À1.64 0.02 28.91 6.47 20.20 0.17 0.65 0.36 MGM-56 À1.17 0.04 À1.71 0.03 À1.17 0.04 À1.71 0.03 29.76 4.70 22.55 0.16 0.04 0.07 2.96 Ga Parktown Formation DRH-34 À0.29 0.05 À0.50 0.04 À0.29 0.05 À0.50 0.04 29.10 24.68 3.98 0.23 0.04 0.08 DRH-42 À0.34 0.05 À0.47 0.03 À0.35 0.02 À0.51 0.06 30.96 25.55 4.87 0.08 0.15 0.08 À0.36 0.07 À0.55 0.04 DRH-50 À0.45 0.04 À0.70 0.03 À0.45 0.04 À0.70 0.03 18.86 8.96 8.91 0.11 0.28 b.d. DRH-55 À0.39 0.06 À0.66 0.05 À0.43 0.06 À0.67 0.02 29.39 25.47 3.53 0.06 0.02 b.d. À0.47 0.05 À0.68 0.04 DRH-61 À0.50 0.05 À0.72 0.03 À0.48 0.02 À0.68 0.05 29.20 24.59 4.15 0.17 0.07 b.d. À0.47 0.05 À0.64 0.03 DRH-65 À0.36 0.06 À0.50 0.04 À0.36 0.06 À0.50 0.04 27.69 23.92 3.39 0.07 0.02 0.69 DRH-79 À0.49 0.07 À0.63 0.04 À0.49 0.07 À0.63 0.04 27.46 8.15 17.38 0.04 0.02 0.46 DRH-88 0.04 0.09 À0.01 0.04 0.05 0.01 0.02 0.04 41.72 20.22 19.35 0.14 0.08 b.d. 0.06 0.09 0.04 0.04 DRH-93 À0.55 0.04 À0.76 0.02 À0.55 0.04 À0.76 0.02 36.14 19.23 15.22 0.25 0.09 b.d. DRH-96 À0.31 0.05 À0.40 0.04 À0.31 0.05 À0.40 0.04 17.17 2.23 13.45 0.51 0.00 b.d. DRH-99 0.09 0.06 0.11 0.03 0.09 0.06 0.11 0.03 15.82 1.50 12.89 0.07 0.02 b.d. DRH-102 0.02 0.04 À0.04 0.03 0.01 0.02 À0.02 0.02 14.87 0.83 12.64 0.70 0.42 b.d. 0.00 0.04 À0.01 0.04 DRH-107 0.10 0.06 0.19 0.03 0.10 0.06 0.19 0.03 12.02 0.58 10.30 0.18 0.05 b.d. K.E. Yamaguchi et al. / Chemical Geology 218 (2005) 135–169 145

Table 1 (continued) P Sample ID Individual analyses Grand average Fe2O3 Fe2O3 FeO Corg Ccarb S d56Fe 2S.E. d57Fe 2S.E. d56Fe 1rd57Fe 1r (wt.%) (wt.%) (wt.%) (wt.%) (wt.%) (wt.%) Kaapvaal Craton Witwatersrand Supergroup West Rand Group 2.96 Ga Parktown Formation DRH-112 À0.04 0.04 0.03 0.03 À0.04 0.04 0.03 0.03 7.99 1.28 6.04 0.27 0.01 b.d. DRH-121 0.01 0.05 0.01 0.03 0.01 0.05 0.01 0.03 7.84 1.42 5.78 0.22 0.00 b.d. Swaziland Supergroup Fig Tree Group 3.25 Ga Sheba Formation (greywackes) MRE-01 0.08 0.07 0.12 0.03 0.08 0.07 0.12 0.03 6.08 0.58 4.95 0.16 1.77 b.d. MRE-07 À0.09 0.06 À0.07 0.04 À0.09 0.06 À0.07 0.04 9.13 0.46 7.80 0.21 1.21 b.d. MRE-13 À0.08 0.06 À0.09 0.04 À0.08 0.06 À0.09 0.04 9.09 0.58 7.66 n.d. 2.17 b.d. MRE-15 À0.04 0.08 À0.13 0.06 À0.04 0.01 À0.10 0.05 6.87 0.60 5.64 0.17 1.15 b.d. À0.03 0.07 À0.06 0.00 MRE-18 0.11 0.07 0.12 0.04 0.11 0.07 0.12 0.04 n.a. n.a. n.a. n.a. n.a. n.a. MRE-21 À0.02 0.06 À0.06 0.04 À0.02 0.06 À0.06 0.04 7.17 0.73 5.8 0.16 1.51 b.d. MRE-22 0.13 0.07 0.14 0.04 0.13 0.07 0.14 0.04 7.43 0.60 6.15 0.17 1.21 b.d. MRE-28 0.13 0.07 0.13 0.04 0.13 0.07 0.13 0.04 n.a. n.a. n.a. n.a. n.a. n.a. MRE-33 0.17 0.08 0.23 0.04 0.17 0.08 0.23 0.04 n.a. n.a. n.a. n.a. n.a. n.a. 3.25 Ga Sheba Formation (black shales) PU-07 À0.35 0.07 À0.45 0.03 À0.32 0.04 À0.40 0.06 43.50 3.17 36.30 0.68 8.07 0.15 À0.29 0.06 À0.36 0.05 PU-09 À0.37 0.04 À0.49 0.03 À0.35 0.03 À0.50 0.01 42.93 0.43 38.25 0.94 8.20 b.d. À0.33 0.05 À0.50 0.03 PU-11 À0.29 0.05 À0.34 0.04 À0.27 0.02 À0.38 0.05 36.49 0.99 31.95 0.43 6.91 0.16 À0.26 0.05 À0.41 0.04 PU-12 À0.07 0.06 À0.04 0.03 À0.07 0.06 À0.04 0.03 37.92 3.92 30.60 0.94 7.27 0.67 PU-13 0.07 0.06 0.05 0.04 0.07 0.06 0.05 0.04 31.61 1.29 27.29 1.43 6.25 b.d. PU-14 À0.15 0.07 À0.25 0.04 À0.11 0.06 À0.22 0.04 43.45 2.39 36.95 1.02 8.38 b.d. À0.07 0.08 À0.19 0.04 PU-16 0.25 0.04 0.37 0.03 0.26 0.02 0.43 0.09 13.25 1.45 10.62 2.97 2.19 b.d. 0.28 0.09 0.49 0.04 PU-18 0.16 0.06 0.20 0.03 0.13 0.05 0.18 0.03 40.21 3.43 33.10 1.29 7.83 b.d. 0.10 0.10 0.16 0.06 PU-19 À0.29 0.10 À0.45 0.06 À0.22 0.07 À0.32 0.11 11.73 1.68 9.04 0.46 3.24 b.d. À0.16 0.06 À0.22 0.05 À0.20 0.05 À0.30 0.04 PU-20B 0.03 0.08 0.13 0.04 0.03 0.08 0.13 0.04 17.55 3.25 12.87 0.22 2.78 b.d. PU-21 À0.01 0.07 0.06 0.04 À0.01 0.07 0.06 0.04 38.35 1.97 32.74 1.31 6.71 2.15 PU-22 0.10 0.05 0.14 0.03 0.12 0.02 0.14 0.00 25.32 6.76 16.70 2.28 3.88 2.70 0.13 0.06 0.14 0.03 PU-24 0.01 0.05 0.04 0.03 0.00 0.01 0.05 0.00 49.30 1.36 43.14 1.08 9.22 b.d. À0.01 0.06 0.05 0.04 PU-26 À0.15 0.06 À0.27 0.03 À0.19 0.04 À0.27 0.02 42.62 0.59 37.83 0.49 8.14 b.d. À0.20 0.05 À0.29 0.03 À0.24 0.06 À0.27 0.04 À0.16 0.05 À0.24 0.04 PU-27 À0.27 0.08 À0.37 0.04 À0.27 0.08 À0.37 0.04 43.11 1.47 37.47 0.93 8.07 b.d. PU-28 À0.51 0.07 À0.73 0.04 À0.41 0.08 À0.58 0.14 41.13 0.24 36.80 0.42 7.80 b.d. À0.35 0.06 À0.48 0.04 À0.37 0.05 À0.53 0.04 (continued on next page) 146 K.E. Yamaguchi et al. / Chemical Geology 218 (2005) 135–169

Table 1 (continued) P Sample ID Individual analyses Grand average Fe2O3 Fe2O3 FeO Corg Ccarb S d56Fe 2S.E. d57Fe 2S.E. d56Fe 1rd57Fe 1r (wt.%) (wt.%) (wt.%) (wt.%) (wt.%) (wt.%) Pilbara Craton Mt. Bruce Supergroup Hamersley Group 2.60 Ga Wittenoom Dolomite Formation WRL-104 À0.49 0.04 À0.74 0.02 À0.53 0.06 À0.80 0.09 10.62 0.00 9.56 0.74 10.59 0.08 À0.57 0.06 À0.86 0.03 WRL-105 À0.61 0.10 À0.88 0.05 À0.63 0.02 À0.90 0.02 13.01 2.61 9.36 2.78 0.13 1.64 À0.65 0.08 À0.91 0.04 WRL-106 À0.74 0.05 À1.12 0.04 À0.76 0.04 À1.14 0.03 17.08 3.47 12.25 2.67 0.04 0.14 À0.79 0.06 À1.16 0.03 2.60 Ga Carawine Dolomite Formation RHDH-101 À1.87 0.07 À2.81 0.04 À1.87 0.07 À2.81 0.04 15.19 1.31 12.49 1.41 7.21 0.44 RHDH-103 À1.98 0.05 À2.84 0.03 À1.98 0.05 À2.84 0.03 4.07 1.67 2.16 1.02 9.22 0.08 RHDH-105 À2.28 0.08 À3.43 0.05 À2.28 0.08 À3.43 0.05 7.73 4.32 3.07 2.91 6.10 0.73 RHDH-107 À1.44 0.07 À2.11 0.04 À1.39 0.07 À2.04 0.10 5.44 2.57 2.58 1.91 6.68 0.14 À1.34 0.05 À1.97 0.03 RHDH-109 À0.80 0.08 À1.27 0.05 À0.80 0.08 À1.27 0.05 6.74 3.37 3.03 4.24 2.15 0.36 RHDH-110 À0.42 0.06 À0.71 0.04 À0.42 0.06 À0.71 0.04 4.20 1.98 2.00 4.45 7.29 1.23 N2.60 Ga Marra Mamba Iron Formation WRL-110 À0.65 0.06 À1.00 0.03 À0.63 0.03 À0.92 0.10 28.95 6.02 20.64 0.19 0.01 b.d. À0.62 0.05 À0.85 0.04 WRL-112 À0.17 0.05 À0.21 0.03 À0.14 0.04 À0.18 0.04 34.33 14.20 18.12 0.37 1.51 0.37 À0.11 0.05 À0.15 0.03 WRL-113 0.12 0.09 0.23 0.05 0.15 0.05 0.28 0.07 32.38 16.59 14.21 0.93 1.40 3.16 0.19 0.05 0.33 0.03 WRL-114 0.31 0.05 0.46 0.03 0.33 0.02 0.46 0.00 34.70 12.87 19.65 0.74 3.80 0.14 0.35 0.05 0.45 0.03 Fortescue Group 2.69 Ga Jeerinah Formation WRL-118 À1.54 0.05 À2.25 0.04 À1.54 0.05 À2.25 0.04 4.40 2.63 1.59 5.52 0.14 1.55 WRL-120 À0.73 0.08 À1.14 0.05 À0.71 0.03 À1.11 0.05 3.45 2.87 0.52 7.39 2.32 1.84 À0.69 0.07 À1.07 0.03 WRL-122 À0.79 0.09 À1.16 0.05 À0.79 0.09 À1.16 0.05 4.71 3.72 0.89 12.04 0.38 2.98 WRL-123 À1.00 0.06 À1.42 0.03 À1.00 0.06 À1.42 0.03 5.79 3.72 1.86 8.19 0.34 3.03 WRL-127 À0.71 0.06 À1.01 0.03 À0.67 0.05 À0.99 0.02 13.34 3.05 9.26 2.73 0.31 1.95 À0.64 0.09 À0.97 0.04 WRL-129 À0.42 0.04 À0.62 0.04 À 0.42 0.04 À0.62 0.04 5.40 1.16 3.82 3.87 0.10 0.74 2.69 Ga Lewin Shale Formation RHDH-111 0.01 0.07 0.00 0.03 0.05 0.05 0.04 0.06 4.25 1.94 2.08 2.67 0.34 0.21 0.08 0.05 0.08 0.03 RHDH-112 À0.24 0.07 À0.43 0.05 À0.24 0.07 À0.43 0.05 4.12 2.35 1.59 6.83 0.32 1.44 RHDH-114 0.04 0.04 0.06 0.03 0.04 0.04 0.06 0.03 4.76 2.32 2.20 1.56 0.19 0.07 RHDH-115 0.36 0.11 0.53 0.06 0.36 0.11 0.53 0.06 5.74 0.88 4.37 2.66 0.03 0.07 RHDH-117 0.02 0.08 0.10 0.04 0.09 0.10 0.12 0.03 7.52 1.91 5.05 1.83 0.04 0.59 0.16 0.09 0.14 0.05 RHDH-119 0.04 0.05 0.06 0.03 0.04 0.05 0.06 0.03 8.63 2.06 5.91 2.50 0.44 0.58 RHDH-120 0.38 0.06 0.55 0.03 0.38 0.06 0.55 0.03 10.61 0.19 9.38 2.05 0.10 4.04 RHDH-121 0.28 0.05 0.40 0.04 0.28 0.05 0.40 0.04 4.77 1.36 3.07 2.18 0.85 0.28 2.72 Ga Pillingini Tuff Formation WRL-133 0.35 0.07 0.53 0.05 0.30 0.07 0.46 0.10 9.78 1.05 7.86 n.a. n.a. n.a. 0.25 0.06 0.39 0.04 K.E. Yamaguchi et al. / Chemical Geology 218 (2005) 135–169 147

Table 1 (continued) P Sample ID Individual analyses Grand average Fe2O3 Fe2O3 FeO Corg Ccarb S d56Fe 2S.E. d57Fe 2S.E. d56Fe 1rd57Fe 1r (wt.%) (wt.%) (wt.%) (wt.%) (wt.%) (wt.%) Pilbara Craton Mt. Bruce Supergroup Fortescue Group 2.72 Ga Pillingini Tuff Formation WRL-134 0.29 0.05 0.40 0.04 0.29 0.05 0.40 0.04 16.96 1.75 13.69 0.21 0.01 0.09 WRL-135 0.30 0.04 0.40 0.03 0.30 0.04 0.40 0.03 8.34 1.84 5.85 0.26 0.12 0.07 WRL-136 À0.12 0.09 À0.08 0.05 À0.09 0.04 À0.06 0.03 7.48 0.66 6.14 0.14 0.42 0.13 À0.06 0.07 À0.05 0.04 56 56 54 Iron isotopic compositions reported using standard delta notation: d Fe=[(Rsample/Rstandard)À1]*1000, where R = Fe/ Fe, standard=average igneous rocks (Beard et al., 2003a,b). S.E.: standard error. 1r: 1 standard deviation. Corg: organic carbon. Ccarb: carbonate carbon. n.a.: not analyzed. b.d.: below detection. produced at PSU and ACTLAB, Inc. for both of the of igneous rocks (Beard et al., 2003b). On this scale, studied samples and the SRM samples agreed very the IRMM-014 standard has a d56Fe value of well. Repeated analyses of the same samples yield a À0.09F0.05x. reproducibility for total Fe and Fe2+ contents of better than 5%. Total C, Corg, and S contents were measured on 4. Results powdered samples using an elemental analyzer (EA) at PSU (CE Instruments NA2500) and also at Tohoku The whole-rock samples analyzed in this study 3+ 2+ University (Carlo Erba EA1108). For Corg content have wide ranges in Fe ,Fe , total Fe, Corg,Ccarb, determinations, the powdered samples were treated and S contents (Table 1), commensurate with the with 2 N HCl at room temperature for N24 h to wide variety of lithologies involved. In some cases, completely decompose carbonate. Ccarb contents were petrographic observations in transmitted and calculated from the difference in C contents deter- reflected light, augmented by XRD data, can identify mined for bulk powders and those measured on the dominant Fe-bearing mineral assemblages, and decarbonated powders. The reproducibility of Corg these observations document that some suites are and S contents was better than 0.2 wt.%. enriched in Fe carbonates (bsideriteQ), magnetite, and Isotopic compositions of Fe (Table 1)were chlorite, and some contain variable amounts of measured using MC-ICP-MS (multi collector, induc- sulfides. The fine-grained nature of most of the tively coupled plasma mass spectrometry; GV Instru- samples, however, requires classification based on ments IsoProbe) at the University of Wisconsin- bulk chemical compositions. Corg contents range up Madison. Detailed analytical methods are given in to 12 wt.%, and we define those samples that contain Beard et al. (2003b) and Albare`de and Beard (2004), N0.5 wt.% Corg as bCorg-richQ, and those that have and are only briefly described here. The powdered b0.5 wt.% as bCorg-poorQ. Carbonate-C contents samples were decomposed by mixed acid digestion (Ccarb) as high as 12 wt.% are measured in the (HF-HNO3) and purified for Fe using anion-exchange sample suite, and we define those that have N2.0 chromatography. Forty-one of the 120 samples used in wt.% Ccarb as bCcarb-richQ. Iron contents vary greatly, this study were analyzed at least two times, and these where total Fe is as high as 35 wt.% in the sample repeat measurements demonstrate an external repro- Psuite, and we define those samples that have ducibility of better than F0.06x (1r), which is Fe2O3 N10 wt.% as bFe-richQ. S and Corg contents essentially identical to that obtained on pure stand- are generally well correlated in the samples, where ards. Isotopic compositions are reported using stand- significant S contents are largely found in the Corg- ard delta notation in units of per mil (parts per 1000 or rich suite (Table 1). Using these definitions, many 56 3 x); d Fe=[(Rsample/RBulk Earth) À1]Â10 ,where sample suites fall into more than one group. Average 56 54 R = Fe/ Fe and RBulk Earth is taken as the average chemical and isotopic compositions are given in 148 K.E. Yamaguchi et al. / Chemical Geology 218 (2005) 135–169

Table 2, and a summary of the samples and the low in total Fe contents (Tables 2 and 3), and these chemical groups that they fall in is given in Table 3. units are most closely analogous to the rocks used in Because our chemical classifications are based on the the post-Archean shale composite of Taylor and dominant characteristics for specific geologic units, McLennan (1985), as well as the Phanerozoic low- individual samples within a unit will sometimes have Corg clastic rocks and marine sediments analyzed for compositions that fall outside our definitions. Fe isotope compositions by Beard et al. (2003a). The relatively high Ccarb but low FeO contents of the 4.1. Fe, Corg,Ccarb, and S contents Black Reef Formation indicate that the carbonate phase is dolomite and not an Fe-rich carbonate such Six units fall in the blow-CorgQ group, greywackes as siderite or ankerite (Fig. 5). In contrast, the of the 3.25 Ga Sheba Formation, the 2.96 Ga moderately high Ccarb and FeO contents of the Parktown Formation, the ~2.9 Ga Rietkuil Forma- greywackes of the Sheba Formation suggest that tion, the 2.72 Ga Pillingini Tuff Formation, the 2.64 Fe-bearing dolomite occurs in these rocks (Fig. 6), as Ga Black Reef Formation, the 2.35 Ga Timeball Hill well as silicate minerals, basedP on petrographic Formation, and the red shales of the ~2.2 Ga observations. The very high Fe3+/ Fe contents of Gamagara Formation (Figs. 5 and 6). Of this suite, the red shales of the Gamagara Formation, in all but the Black Reef Formation are also classified addition to their color, indicates that ferric oxides as bCcarb-poorQ (Table 3). The greywackes of the such as hematite are the primary Fe-bearing phases Sheba Formation, as well as the Pillingini Tuff, in these rocks. Other bFe-richQ units that belong to Black Reef, and Timeball Hill formations are also the blow-CorgQ suite include the Parktown and

Table 2 Summary of geochemical data for Archean shales and greywackes of this study P (3+) (3+) (2+) 56 a Formation Approx. Number of Fe2 O3 Fe2 O3 Fe O d Fe Corg Ccarb S depositional samples Average Average Average Average Average Average Average of age (Ga) F1r F1r F1r F1r F1r F1r F1r (wt.%) (wt.%) (wt.%) (wt.%) (wt.%) (wt.%) (wt.%) Kaapvaal Craton Gamagara Fm. ~2.2 8 13.30F3.62 11.63F4.51 1.50F1.31 +0.07F0.13 0.1F0.0 0.1F0.3 0.0F0.0 (red shales) Timeball Hill Fm. 2.35 12 8.31F2.27 0.88F0.54 6.69F1.83 +0.03F0.18 0.4F0.1 0.2F0.3 0.1F0.1 Oak Tree Fm. 2.56 3 2.03F0.56 0.09F0.01 1.74F0.51 +0.49F0.20 1.9F0.4 0.1F0.1 0.6F0.5 Black Reef Fm. 2.64 4 3.06F2.41 0.57F0.59 2.24F2.03 À0.31F0.40 0.4F0.5 8.9F5.9 0.3F0.4 Rietgat Fm. 2.71 7 7.15F1.59 0.55F0.23 5.94F1.26 À0.12F0.2 0.9F0.5 0.2F0.2 0.7F0.8 Rietkuil Fm. ~2.9 15 29.77F2.91 6.02F2.78 21.37F2.33 À1.18F0.12 0.2F0.0 0.4F0.6 0.3F0.3 Parktown Fm. 2.96 15 23.08F10.31 12.57F10.87 9.46F5.41 À0.23F0.24 0.2F0.2 0.1F0.1 0.1F0.2 Sheba Fm. 3.25 9b 7.63F1.23 0.59F0.09 6.33F1.15 +0.04F0.10 0.2F0.0 1.5F0.4 0.0F0.0 (greywackes) Sheba Fm. 3.25 16 34.90F11.68 2.15F1.66 29.48F10.96 À0.10F0.20 1.1F0.7 6.6F2.2 0.4F0.8

Pilbara Craton Wittenoom 2.60 3 13.57F3.27 2.03F1.81 10.39F1.61 À0.64F0.12 2.1F1.1 3.6F6.1 0.6F0.9 Dolomite Fm. Carawine 2.60 6 7.23F4.15 2.54F1.41 4.22F4.07 À1.46F0.73 2.7F1.5 6.4F2.4 0.5F0.4 Dolomite Fm. Marra Mamba N2.60 4 32.59F2.63 12.42F4.54 18.16F2.83 À0.07F0.42 0.8F0.3 1.7F1.6 0.9F1.5 Iron Fm. Jeerinah Fm. 2.69 6 6.18F3.60 2.86F0.94 2.99F3.28 À0.86F0.38 6.6F3.4 0.6F0.9 2.0F0.9 Lewin Shale Fm. 2.69 8 6.30F2.37 1.63F0.76 4.21F2.59 +0.12F0.21 2.8F1.7 0.3F0.3 0.9F1.3 Pillingini Tuff Fm. 2.72 4 10.64F4.32 1.33F0.57 8.39F3.65 +0.20F0.19 0.2F0.1 0.2F0.2 0.1F0.0 a d56 Fe values are calculated for 56 Fe/54 Fe ratios relative to Bulk-Earth (see Beard et al., 2003a,b), in units of per mil (parts per 1000). b Nine samples for Fe isotope analysis, and six samples for Fe2O3 and FeO content analysis. K.E. Yamaguchi et al. / Chemical Geology 218 (2005) 135–169 149

Table 3 Classification of rock samples based on the carbon and iron contents

Formation Abbreviation Age Corg-rich Ccarb-rich PFe-rich Major Fe-bearing minerals Rock type (Ga) Corg N0.5 Ccarb N2 Fe2O3 N10 assigned wt.% wt.% wt.% Kaapvaal Craton 3+ Gamagara Fm. GMG ~2.2 No No Yes Fe -oxides Corg-poor, (red shales) Fe-rich 2+ Timeball Hill Fm. TBH 2.35 No No No Fe -silicates Corg-poor 2+ Oak Tree Fm. OAK 2.56 Yes No No Fe -silicates Corg-rich 2+ Black Reef Fm. BLR 2.64 No Yes No Fe-bearing dolomiteFFe -silicates Corg-poor, Ccarb-rich 2+ Rietgat Fm. RGT 2.71 Yes No No Fe -silicates Corg-rich Rietkuil Fm. RKL ~2.9 No No Yes Magnetite+Fe2+-silicates Fe-rich 3+ 2+ Parktown Fm. PKT 2.96 No No Yes Magnetite+Fe -oxides+Fe -silicates Corg-poor, Fe-rich 2+ Sheba Fm. SHW 3.25 No No No Fe -silicates Corg-poor (greywackes)

Sheba Fm. SHS 3.25 Yes Yes Yes Mg-siderite Ccarb-rich, Fe-rich

Pilbara Craton 2+ Wittenoom WTN 2.6 Yes Yes Yes Fe-oxides+Fe -silicatesFFe-bearing Corg-rich, Dolomite Fm. dolomite Fe-rich

Carawine CRW 2.6 Yes Yes No Magnetite+Fe-bearing dolomite Corg-rich, Dolomite Fm. Ccarb-rich Marra Mamba MMB N2.60 Yes No Yes Fe-bearing dolomite+magnetite Corg-rich, Iron Fm. +Fe2+-silicates Fe-rich 2+ Jeerinah Fm. JRN 2.69 Yes No No Magnetite+Fe -silicates+sulfides Corg-rich 2+ Lewin Shale Fm. LEW 2.69 Yes No No Fe -silicatesFmagnetiteFsulfides Corg-rich 2+ Pillingini Tuff Fm. PLT 2.72 No No No Fe -silicates Corg-poor

Rietkuil formations. Ferric oxides are a minor Formation and Marra Mamba Iron Formation) are also component in these rocks based on color and classified as bFe-richQ (Table 3); the shales of the b Q petrographicP observations. The moderately high Sheba Formation also belong to the Ccarb-rich suite 3+ Fe / Fe ratios (although not as high as the (Table 3). The Corg-rich carbonates of the Wittenoom Gamagara rocks) reflect the high abundance of Dolomite Formation are additionally classified as magnetite in these rocks. Magnetite is particularly bCcarb-richQ and bFe-richQ (Table 3). Of the remaining prominent in the Rietkuil Formation based on tests units in the bCorg-richQ suite, the Carawine Dolomite using a hand magnet and reflected light microscopy, Formation is also classified as bCcarb-richQ, although it although a ferrous Fe-bearing phase is also present is not considered part of the bFe-richQ suite (Fig. 7). based on Fe2O3–FeO variations (Fig. 6), which is The mineralogy of the major Fe-bearing phases in most likely chlorite based on transmitted light the bCorg-richQ suite is apparently quite variable. Iron microscopy. in the Carawine Dolomite seems likely to lie in Fe- The majority of units in this study fall in the bCorg- bearing dolomites and magnetite, given the observed richQ group, including the black shales of the 3.25 Ga Fe2O3–FeO and FeO–Ccarb variations (Fig. 7), Sheba Formation, the ~2.7 Ga Rietgat, Lewin Shale, whereas the major Fe-bearing phase in the black and Jeerinah formations, black shales from the ~2.60 shales of the Sheba Formation is clearly Mg-bearing Ga Marra Mamba Iron Formation, as well as the siderite (Fig. 8). One of the three samples from the Carawine and Wittenoom Dolomite formations, and Wittenoom Dolomite is high in carbonate (Fig. 7), but the 2.56 Ga Oak Tree Formation (Figs. 7 and 8). Two this is a Ca–Mg carbonate, and it seem most likely of the black shale units in the bCorg-richQ suite (Sheba that the Fe budget for all three samples are dominated 150 K.E. Yamaguchi et al. / Chemical Geology 218 (2005) 135–169

Corg-poor group

1.0 15 a c 0.8 10 0.6 Fe ∑

/ [wt.%] 3+ 0.4 org Fe GMG 5 C 0.2 TBH BLR

0.0 0 0.5 %

40 15 b d 30 10

20 [wt.%] Fe [wt.%] carb ∑

5 ∑ Fe = 7 % or siderite 10 C ∑Fe2O3 = 10 % 2.0 % 0 0 -3 -2 -1 0 1 56 30 δ Fe [‰] e e

20 magnetit [wt.%] 3 O 2 10 ∑ Fe2O3 = 10 % Fe

0 0102030 FeO [wt.%] P P 3+ 56 56 Fig. 5. Inter-relations for Corg-poor group samples among contents and/or isotopeP compositions ofP (a) Fe / Fe vs. d Fe, (b) Fe vs. d Fe, (c) FeO vs. Corg, (d) FeO vs. CcarbPand (e) FeO vs. Fe2O3. Horizontal lines at Fe2O3 =10 wt.% or Fe=7 wt.% in (b), at Corg =0.5 wt.% in (c), at Ccarb =2.0 wt.% in (d), and at Fe2O3 =10 wt.% in (e) represents bthresholdQ contents for grouping samples into Corg-rich/-poor, Ccarb-rich, and Fe-rich suites. Lines for stoichiometric composition of siderite (FeCO3,Ccarb/FeO=1:6) and of magnetite (Fe3O4,Fe2O3/FeO=20:9) are drawn in (d) and (e), respectively. Filled circles: ~2.2 Ga Gamagara Formation; open squares: 2.35 Ga Timeball Hill Formation; filled diamonds: 2.64 Ga Black Reef Formation. See Table 3 for classification of sample types and key to abbreviations.

by oxides and silicates, although a contribution from Fe2O3–FeO variations; it seems unlikely that the high Fe-bearing dolomite cannot be excluded. Iron in the Fe2O3 contents reflect abundant hematite in these Marra Mamba shales seems likely to reflect a mixture black shales given the high Corg contents (Fig. 7), and of carbonate, silicate, and oxides based on Ccarb– the oxide phase is probably magnetite. The low total K.E. Yamaguchi et al. / Chemical Geology 218 (2005) 135–169 151

Corg-poor group

1 15 PLT a c 0.8 RKL PKT SHW 10 0.6 Fe

∑ [wt.%] / 3+

0.4 org C

Fe 5 0.2

0 0 0.5 %

40 15 b d 30 10

20 [wt.%] Fe [wt.%]

∑

carb 5 ∑ Fe = 7 % or siderite 10 ∑ C Fe2O3 = 10 % 2.0 % 0 0 -3 -2 -1 0 1 δ56 30 30 Fe [‰] te e

20 magneti 20

[wt.%] ∑ Fe O = 10 %

3 2 3 O

2 10 10 Fe

0 0 0102030 FeO [wt.%] P 3+ Fig. 6. Inter-relationsP for additional samples from the Corg-poor group samples among contents and/or isotopeP compositions of (a)P Fe / Fe vs. 56 56 d Fe, (b) Fe vs. d Fe, (c) FeO vs. Corg, (d) FeO vs. CcarbPand (e) FeO vs. Fe2O3. Horizontal lines at Fe2O3 =10 wt.% or Fe=7 wt.% in (b), at Corg =0.5 wt.% in (c), at Ccarb =2.0 wt.% in (d), and at Fe2O3 =10 wt.% in (e) represents bthresholdQ contents for grouping samples into Corg-rich/-poor, Ccarb-rich, and Fe-rich suites. Lines for stoichiometric composition of siderite (FeCO3,Ccarb/FeO=1:6) and of magnetite (Fe3O4, Fe2O3/FeO=20:9) are drawn in (d) and (e), respectively. Filled circles: 2.72 Ga Pillingini Tuff Formation; open squares: ~2.9 Ga Rietkuil Formation; filled diamonds: 2.96 Ga Parktown Formation; open triangles: 3.25 Ga Sheba Formation (greywackes). See Table 3 for classification of sample types and key to abbreviations.

P 3+ Fe and Fe2O3 contents of the Oak Tree Formation silicate minerals. The high Fe / Fe ratios of the (Fig. 7), as well as the lack of significant carbonate, Jeerinah Formation, coupled with the high Corg suggests that Fe in these shales is dominated by contents (Fig. 8) and black color indicates that Fe in 152 K.E. Yamaguchi et al. / Chemical Geology 218 (2005) 135–169

Corg-rich group

1.0 15 OAK a c 0.8 WTN CRW MMB 10

Fe 0.6 ∑

/ [wt.%] 3+ 0.4 org Fe 5 C 0.2

0.0 0 0.5 %

40 15 b d 30 10

20 ∑ Fe = 7 % or [wt.%] ∑Fe O = 10 % Fe [wt.%] 2 3 carb

∑ 5 siderite 10 C 2.0 % 0 0 -3 -2 -1 0 1 30 δ56Fe [‰] te e

20 magneti [wt.%] 3 O

2 10 ∑Fe2O3 = 10 % Fe

0 010 20 30 FeO [wt.%] P 3+ Fig. 7. Inter-relationsP for the additional samples from Corg-rich group samples among contents and/or isotopeP compositions of (a)P Fe / Fe vs. 56 56 d Fe, (b) Fe vs. d Fe, (c) FeO vs. Corg, (d) FeO vs. CcarbPand (e) FeO vs. Fe2O3. Horizontal lines at Fe2O3 =10 wt.% or Fe=7 wt.% in (b), at Corg =0.5 wt.% in (c), at Ccarb =2.0 wt.% in (d), and at Fe2O3 =10 wt.% in (e) represents bthresholdQ contents for grouping samples into Corg-rich/-poor, Ccarb-rich, and Fe-rich suites. Lines for stoichiometric composition of siderite (FeCO3,Ccarb/FeO=1:6) and of magnetite (Fe3O4, Fe2O3/FeO=20:9) are drawn in (d) and (e), respectively. Filled circles: 2.56 Ga Oak Tree Formation; open squares: 2.60 Ga Wittenoom Dolomite Formation; filled diamonds: 2.60 Ga Carawine Dolomite Formation; open triangles: N2.6 Ga Marra Mamba Iron Formation. See Table 3 for classification of sample types and key to abbreviations. P these rocks is dominated by magnetite, although Fe- lower Fe3+/ Fe ratios than the temporally correlative bearing silicates are also likely to be present. In Jeerinah Formation may indicate a lower Fe3+ oxide contrast, that the Lewin Formation has significantly content for the Lewin Formation, and Fe is likely to K.E. Yamaguchi et al. / Chemical Geology 218 (2005) 135–169 153

Corg-rich group

1.0 15 a c 0.8 10 0.6 Fe ∑

/ [wt.%] 3+ 0.4 JRN org Fe 5

LEW C 0.2 RGT SHS 0.0 0 0.5 %

40 15 b d 30 10

20 [wt.%]

Fe [wt.%] siderite carb

∑ ∑ Fe = 7 % or 5 10 ∑ C Fe2O3 = 10 % 2.0 % 0 0 -3 -2 -1 0 1 56 50 δ Fe [‰] e e 40

magnetit 30 [wt.%] 3 20 O 2 ∑ Fe2O3 = 10 %

Fe 10

0 01020304050 FeO [wt.%] P 3+ Fig. 8. Inter-relationsP for additional samples from the Corg-rich group samples among contents and/or isotopeP compositions of (a)P Fe / Fe vs. 56 56 d Fe, (b) Fe vs. d Fe, (c) FeO vs. Corg, (d) FeO vs. CcarbPand (e) FeO vs. Fe2O3. Horizontal lines at Fe2O3 =10 wt.% or Fe=7 wt.% in (b), at Corg =0.5 wt.% in (c), at Ccarb =2.0 wt.% in (d), and at Fe2O3 =10 wt.% in (e) represents bthresholdQ contents for grouping samples into Corg-rich/-poor, Ccarb-rich, and Fe-rich suites. Lines for stoichiometric composition of siderite (FeCO3,Ccarb/FeO=1:6) and of magnetite (Fe3O4, Fe2O3/FeO=20:9) are drawn in (d) and (e), respectively. Filled circles: 2.69 Ga Jeerinah Formation; open squares: 2.69 Ga Lewin Shale Formation; filled diamonds: 2.71 Ga Rietgat Formation; open triangles: 3.25 Ga Sheba Formation (sideritic shale). See Table 3 for classification of sample types and key to abbreviations. resideP primarily in silicate minerals. The very low reflected light microscopy, indicate that the major 3+ Fe / Fe ratios and low Ccarb contents of the Rietgat repository of Fe in this unit includes silicate and Formation (Fig. 8), in addition to observations by sulfide minerals. 154 K.E. Yamaguchi et al. / Chemical Geology 218 (2005) 135–169

Observations using reflected light microscopy 1 indicates that where sulfides are present, they are 0.5 % dominated by pyrite. Although S contents tend to be highest in samples that are rich in C (Table 1), in org 0 0.00±0.05 ‰ most cases, Fe2+/S ratios are low, indicating that pyrite

is a minor repository for Fe. Of the 120 samples ] –0.5 ‰ analyzed, only 19 havez10% of their Fe budget in ‰ pyrite as calculated from S contents and only 5 -1 Fe [ samples have z50% of their Fe budget in pyrite; 4 of 56 δ these samples are from the Jeerinah Formation and the fifth is one sample from the Lewin Formation. -2 4.2. Fe isotope compositions Corg-poor groups

Corg-rich groups The d56Fe values of the whole-rock samples in this -3 study vary from +0.64x to À2.28x, spanning a 0.01 0.1 110 100 large part of the total range that has been measured to C [wt.%] date for terrestrial samples. The grand average of the org 56 x 56 dataset has a d Fe value of À0.30 , suggesting that Fig. 9. d Fe–Corg variations for Archean–Paleoproterozoic sedi- the sample suite is biased to lithologies that have low mentary rocks of this study. Threshold average Corg contents of a d56Fe values relative to those expected for bulk formation (0.5 wt.%, shown by a vertical line) is used to group continental crust, based on the near-zero d56Fe values samples into Corg-rich suite (filled symbols) and Corg-poor suite (open symbols). Open symbols enclosed by large circle indicate the of igneous rocks (Beard et al., 2003b). With the magnetite-rich samples from the ~2.9 Ga Rietkuil Formation. Note 56 exception of one unit, the d Fe values for Corg-poor log scale used for x-axis. Horizontal grey bar at 0.00F0.05x and are restricted to values that lie in between À0.6x and horizontal line at À0.5x indicate d56Fe values for igneous rocks (Beard et al., 2003b) and average composition for mid-ocean ridge +0.2x (Fig. 9); the exceptions are the Corg-poor and magnetite-rich samples of the ~2.9 Ga Rietkuil hydrothermal fluids (Sharma et al., 2001; Beard et al., 2003a), respectively. Formation (Fig. 10). The Corg-rich suite has the largest range in d56Fe values, extending to values as For example, the average d56Fe value for the Timeball low as À2.5x (Figs. 9 and 10). There are strong Hill and Pillingini Tuff Formations and greywackes of correlations between Fe isotope compositions of the the Sheba Formation is +0.06F0.16x (Figs. 5 and 6). samples and chemical groups, particularly when Similarly, the Corg-poor Gamagara Formation, whose considered in light of the major Fe-bearing minerals Fe budget is dominated by Fe3+-oxides (hematite) and that are likely to be present in these fine-grained rocks is also considered part of the bFe-richQ suite, has a 56 (Fig. 10). Although silicate minerals are abundant in relatively homogenous d Fe value of +0.06F0.13Px many samples, we expect detrital and authigenic that is independent of total Fe contents or Fe3+/ Fe 56 silicates to have d Fe values near zero, given the ratios (Fig. 5). In contrast, the Corg-poor, but Ccarb-rich Fe isotope compositions of at least modern weathering Black Reef Formation has quite variable d56Fe values products (Beard and Johnson, 2004a). It seems most (Fig. 5). It is important to note, however, that the three likely that the minerals that will produce bulk-rock carbonate-rich samples of the Black Reef Formation d56Fe values that deviate from zero will include Fe3+- have relatively constant d56Fe values (average: oxides (hematite; Fe2O3), magnetite (Fe3O4), and À0.51F0.12x), whereas the non-carbonate sample carbonate (Fig. 10). In only a few samples, primarily in this unit has a distinctly different d56Fe value of from the Jeerinah Formation, do sulfides control the +0.27x. 56 56 bulk-rock d Fe values. The most unusual d Fe values for the Corg-poor Geologic units that belong to the Corg-poor suite suite are found in the Rietkuil Formation, which also where Fe2+-bearing silicate minerals are the major Fe- belongs to the Fe-rich group. The average d56Fe value bearing phases have d56Fe values very close to zero. for the Rietkuil Formation is À1.18F0.12x (Fig. 6), K.E. Yamaguchi et al. / Chemical Geology 218 (2005) 135–169 155

tions) tend to have lower d56Fe values. The very high Fe3+-oxides (GMG) Fe2O3 contents of several samples, relative to their FeO contents, suggests the presence of Fe3+-oxide, Fe2+-silicates (TBH, SHW, PLT) which, if it had near-zero or positive d56Fe values, Magnetite + Fe3+-oxide might explain the more modestly negative d56Fe + Fe2+-silicates (PKT)

Fe-dolomite + Fe2+-silicates (BLR) Table 4 Summary of Fe isotope compositions of Fe sources and fractiona- Magnetite + Fe2+-silicates tions between Fe species (RKL) 56 Corg-poor group Fe sources and Fe species d Fe (x) References Major Fe sources— Corg-rich group continental: Mg-siderite (SHS) Bulk igneous rocks 0.00F0.05 Beard et al. (2003b) Magnetite + Fe-dolomite F + Fe2+-silicates (WTN, MMB) Phanerozoic weathering 0.0 0.1 Beard products et al. (2003a) Fe2+-silicates ± magnetite ± sulfides (OAK, RGT, LEW) Major Fe sources—marine: Hydrothermal MOR Fe À0.6 to À0.3 Sharma Magnetite + Fe2+-silicates + sulfides (JRN) et al. (2001) Beard Magnetite + Fe-bearing dolomite (CRW) et al. (2003a) Fluid–mineral 3+ -3 -2 -1 0 1 2 fractionations—Fe reduction: δ56Fe [‰] 2+ 3+ Feaq –Fe oxide during À1.3F0.1 Beard et al. DIR (equilibrium) (1999, 2003b) Fig. 10. Summary of d56Fe values for Archean–Paleoproterozoic Fe 2+–Fe3+ oxide during À2.6 to À1.3 Johnson et al. sedimentary rocks of this study relative to geochemical and aq DIR (kinetic) (2005) mineralogical groups. C -poor samples shown in open symbols org Fe 2+–Fe O (magnetite) À1.3F0.1 Johnson et al. and C -rich samples shown in filled symbols. See Table 3 for aq 3 4 org during DIR (equilibrium) (2005) additional details on geochemical groups. 2+ Feaq –Fe CO3 (siderite) 0.0F0.1 Johnson et al. during DIR (equilibrium) (2005) 2+ Feaq –Fe CO3 (siderite) +0.4 to +0.6 Wiesli et al. and this is thought to reflect a mixture between during abiotic formation (2004) 2+ magnetite and Fe -bearing silicate minerals based on Mineral–fluid 2+ petrographic observations and Fe2O3–FeO variations fractionations—Fe (Fig. 6). Given the fact that other units in the C - oxidation: org 3+ 2+ F poor group that are dominated by Fe2+-bearing silicate Fe oxide–Feaq during +0.9 0.2 Bullen et al. 56 abiotic oxidation (2001) minerals have d Fe values near zero (Timeball Hill, 3+ 2+ Fe oxide–Feaq during +1.5F0.2 Croal et al. Pillingini Tuff, and Sheba greywackes; Fig. 10), we APIO (2004) infer that the bulk-rock d56Fe values for the Rietkuil Aqueous Fe3+–Fe2+ Formation reflect a mixture between a very low-d56Fe fractionations at 22 8C: 3+ 2+ F magnetite component and a near-zero Fe2+-bearing Feaq –Feaq +2.9 0.2 Johnson et al. (2002) Welch silicate component. These inferences are supported by et al. (2003) 56 the d Fe values for the Parktown Formation, which d56 Fe values are calculated for 56 Fe/54 Fe ratios relative to Bulk- also belongs to the Corg-poor and Fe-rich suites. Earth (see Beard et al., 2003a,b), in units of per mil (parts per 1000). Samples from the Parktown Formation that have near- For major Fe source, d56 Fe values reported as measured. For zero d56Fe values also have the lowest total Fe isotopic fractionations between phases A and B (A–B), d56 Fe values calculated using isotopic fractionation factors (DA–B) from contents (Fig. 6) and are likely to be dominated by 56 2+ sources noted, setting d FeB equal to zero. For isotope fractiona- Fe -bearing silicate minerals. In contrast, those tions produced by biological cycling of Fe, bDIRQ denotes samples that have higher Fe contents and contain dissimilatory Fe3+ reduction and bAPIOQ denotes anaerobic photo- abundant magnetite (based on petrographic observa- synthetic Fe2+ oxidation. 156 K.E. Yamaguchi et al. / Chemical Geology 218 (2005) 135–169

values of magnetite-bearing Parktown Formation (Fig. 7). The Corg-rich shale units from the Marra samples as compared to the Rietkuil Formation. Mamba Iron Formation contain significant carbonate 56 Three units of the Corg-rich suite do not contain but little sulfide and highly variable d Fe values, high carbonate or total Fe contents and have Fe2+- ranging from À0.63x to +0.33x (Fig. 7). Finally, bearing silicates or Fe sulfides as the primary the most carbonate-rich black shales analyzed in the repository of Fe in the bulk samples, including the Corg-rich suite are those from the Sheba Formation, Oak Tree, Lewin Shale, and Rietgat Formations. where strong correlations between Ccarb and FeO Although the d56Fe values for these units are contents suggest a uniform composition for Mg- significantly more variable than units from the bearing siderite [Fe0.75Mg0.25(CO3)] in these rocks, 2+ Corg-poor suite that also have Fe -bearing silicates but variable abundance (Fig. 8); despite large ranges as the primary Fe-bearing phase, the average d56Fe in total Fe contents and Mg-siderite abundance in the value for these three units is +0.09F0.31x. Based Sheba Formation black shales, their d56Fe values are on reflected light microscopy and S contents, the relatively restricted at À0.10F0.20x (Fig. 8). samples in this suite that have d56Fe values furthest The ranges in Fe isotope compositions as a from zero (both positive and negative) also tend to function of bulk-rock composition are summarized have the greatest proportion of sulfide minerals. The in Fig. 10. d56Fe values are lowest in dolomite- and fourth unit analyzed that belongs to the Corg-rich magnetite-rich rocks. The few rocks that are rich in suite that does not contain abundant carbonate or pyrite also have low d56Fe values. Rocks whose Fe 3+ high total Fe contents is theP Jeerinah Formation, budgets are dominated by Fe oxides or silicates 3+ 56 which, based on the high Fe / Fe ratios, Fe2O3– tend to have near-zero to slightly positive d Fe FeO variations (Fig. 8), and high S contents, as well values. Siderite-rich rocks have d56Fe values that are as petrographic observations, appears to have higher than those of dolomite-rich rocks. These magnetite and sulfide as the major controlling Fe observations generally follow those found in mono- phases. The d56Fe values of the Jeerinah Formation mineralic layers in banded iron formations (Johnson are uniformly negative, with an average of et al., 2003). À0.86F0.38x. Two units, the Carawine Dolomite and Wittenoom Dolomite, in the Corg-rich suite also belong to the 5. Discussion Ccarb-rich suite, where the carbonate phase is 56 dolomite, based on FeO–Ccarb variations (Fig. 7), Wide ranges in d Fe values are observed for and these units have consistently negative d56Fe sedimentary rocks over the last 3.3 Ga (Fig. 11). values (Fig. 7). In all cases, low Fe2+/S ratios Although it is striking that the largest range in indicate that sulfide is not the major repository of d56Fe values for whole-rock samples seems to occur Fe. Carbonate abundance in the samples from in the Archean, we caution that our Archean sample Carawine Dolomite is quite variable, and their suite is biased toward samples that are Corg-rich, d56Fe values span a large range, from À0.42x to and many contain high Fe and carbonate contents, À2.28x (Fig. 7), and it is not clear from the bulk whereas samples of similar composition from the compositions or petrographic observations why the Mesoproterozoic, Neoproterozoic, and early Phaner- range in Fe isotope compositions is so large. One ozoic have not been analyzed. Nevertheless, sedi- possibility is that magnetite is also a controlling mentary rocks that are Corg-poor, Ccarb-poor, and Fe- 56 phase, whose presence is inferred from Fe2O3–FeO poor appear to have relatively constant d Fe values variations (Fig. 7). Although the Wittenoom Dolo- over much of Earth’s history, deviating little from mite units analyzed are assigned to the Ccarb-rich the average for igneous rocks, or the compositions suite, carbonate contents are variable (Fig. 7), and of modern river sediments, aerosols, loess, and these samples also belong to the Fe-rich suite, where modern clastic marine rocks (Fig. 11). The largest 56 the Fe-bearing phases may include Fe-bearing range in d Fe values are found in Corg-rich rocks dolomite, Fe2+-bearing silicates, and oxides; d56Fe or those that are rich in carbonate, magnetite, or values for these samples average À0.64F0.12x pyrite, although, so far, Fe isotope variations for K.E. Yamaguchi et al. / Chemical Geology 218 (2005) 135–169 157

Fig. 11. Iron isotope compositions of the Archean–Paleoproterozoic sedimentary rocks of this study versus their depositional ages. Literature data are also included for comparison: magnetite, hematite, siderite, and pyrite separates from BIF of the 2.5 Ga Kuruman Iron Formation (Johnson et al., 2003); late Jurassic Kimmeridge Clay Formation from UK (Matthews et al., 2004), modern aerosol, suspended river load, loess, and marine sediments (Beard et al., 2003a). Also included are data for supergene Fe ore from the Griqualand West, South Africa (K. Yamaguchi et al., unpub. data), and for the Cretaceous (Cenomanian–Turonian) black shales and their adjacent carbonate rocks from Italy (K. Yamaguchi et al., unpub. data). such rocks that are Archean in age seem to be to samples that probably formed under marine larger than those of Mesozoic rocks (Fig. 11). An diagenetic conditions, we discuss samples that belong important exception is the large range measured to the Ccarb-rich suite, comparing dolomite- and for Fe2+-rich pore fluids and pyrite in relatively siderite-rich samples, looking to insights these rocks anoxic, modern marine sediments (Fig. 11). may provide into the Fe isotope compositions of the The large range in bulk-rock d56Fe values mea- ancient oceans. Finally, we discuss rocks that belong sured for the Archean–Paleoproterozoic sedimentary to the Corg-rich or Fe-rich suites in the context of rocks in this study, which correlate with groupings possible biogenic mineral formation. The approaches 2+ based on chemical compositions and mineralogy, by which the Fe isotope composition of Feaq in suggest a variety of pathways and environments were ancient marine environments may be calculated is a involved in Fe cycling prior to final lithification. We major focus in the discussion that follows. A first consider samples that most likely reflect terres- summary of Fe isotope compositions of major Fe trial weathering and transport processes, using sam- sources and mineral-fluid fractionation factors that ples from the Corg-poor and Ccarb-poor suites. Turning bear on this discussion is found in Table 4. 158 K.E. Yamaguchi et al. / Chemical Geology 218 (2005) 135–169

5.1. Iron isotope fractionations during terrestrial range in Fe isotope compositions of igneous rocks weathering (d56Fe=0.00F0.05x) is preserved in modern terrestrial weathering products, including suspended river The Corg-poor and Ccarb-poor sedimentary rocks loads, loess, aerosols, and late Cenozoic Corg-poor probably record the least amount of marine dia- marine sedimentary rocks, despite the fact that genesis, and therefore provide insight into Fe cycling chemical weathering generally results in formation in terrestrial (surficial) weathering processes. In Fig. of extensive secondary oxide minerals and clays. For 12, we compare the d56Fe values samples from the 2.2 example, the suspended river loads plotted in Fig. 12 Ga Gamagara, 2.35 Ga Timeball Hill, and 2.72 Ga contain 20% to 60% oxide minerals by mode, and yet Pillingini Tuff Formations, as well as greywackes have the same restricted d56Fe values as igneous from the 3.25 Ga Sheba Formation, to the Fe isotope rocks, suggesting that despite large changes in Fe3+/ compositions of modern weathering products and Fe2+ ratios during weathering, bulk sedimentary their source materials. As discussed in Beard et al. detritus remains unchanged in its d56Fe values in (2003a,b) and Beard and Johnson (2004a), the narrow modern (oxygenated) environments. Because of the very low solubility of Fe3+-oxide minerals that are formed during modern terrestrial weathering, Fe acts 1 as a conservative element and is not lost through fluid–mineral interactions, resulting in no change in Fe isotope compositions (Beard et al., 2003a; Beard 0.8 and Johnson, 2004a). Iron isotope compositions for the 3.25 to 2.2 Ga 0.6 Corg-poor and Ccarb-poor rocks from this study appear Fe ∑

to show conservative behavior that is similar to /

3+ modern weathering products (Fig. 12), although the

Fe 0.4 range in Fe isotope compositions is slightly larger than that of modern Corg-poor and Ccarb-poor sediments. The relative extent of weathering is noted in 0.2 P terms of Fe3+/ Fe ratios (Fig. 12), and the Archean Pillingini samples and greywackesP from the Sheba 3+ 0 Formation, have a range of Fe / Fe ratios from 0.05 to 0.22 and d56Fe values that cluster (with some -1 -0.5 0 0.5 1 scatter) about zero. The PaleoproterozoicP Timeball δ56 Fe [‰] Hill Formation has similar Fe3+/ Fe ratios, ranging from 0.02 to 0.17, and most samples have d56Fe Modern samples x Igneous rocks values that scatter about zero to +0.1 (Fig. 12). The Suspended river loads red beds of theP ~2.2 Ga Gamagara Formation have 3+ 56 2.2 Ga Gamagara Fm. very high Fe / Fe ratios and have d Fe values that 2.35 Ga Timeball Hill Fm. scatter about zero to +0.1x (Fig. 12). Despite some 2.72 Ga Pillingini Tuff Fm. scatter in the Fe isotope compositions for the 3.25 Ga Sheba Fm. Archean–Paleoproterozoic C -poor and C -poor (greywacke) org carb sedimentary rocks that exceeds analytical errors (Fig. 56 Fig. 12. d56Fe values for igneous rocks (46 samples) and suspended 12), as a group, they have much more restricted d Fe river loads (Beard et al., 2003a,b) and the Corg-poor suites of the values than the other rocks in this study (Fig. 11). Archean–Paleoproterozoic sedimentary rocks studied here, relative If surficial weathering processes have produced no to their extent of secondary oxidation through weathering. Fe apparent Fe isotope fractionation for bulk sedimentary isotope data for igneous rocks and suspended river loads plotted relative to their Fe fraction that is oxide minerals (adapted form detritus over the last 3.25 Ga, does such an observation provide constraints on atmospheric O2 Beard and Johnson, 2004a). Fe isotopeP data for sedimentary rocks from this study plotted relative to Fe3+/ Fe ratios. contents? There is a long-lasting controversy over the K.E. Yamaguchi et al. / Chemical Geology 218 (2005) 135–169 159 timing and mechanism for the rise of atmospheric (Fig. 13; Beard and Johnson, 2004a). Zhu et al. oxygen (e.g., Dimroth and Kimberley, 1976; Clemm- (2002) and Beard and Johnson (2004b) report Fe ery and Badham, 1982; Towe, 1994; Ohmoto, 1997; isotope differences (up to ~0.5x) among some high- Holland, 1999; Phillips et al., 2001), although the temperature minerals in mantle-derived nodules and majority of scientists believe that the atmosphere and eclogites, but these isotopic variations are thought to oceans in the Archean and Paleoproterozoic had been be homogenized during magma generation, and essentially globally anoxic (e.g., Holland, 1984, 1994, therefore are unlikely to be found in high-temperature 1999). Congruent dissolution of minerals, without continental rocks (Beard and Johnson, 2004a). redox changes or formation of a new phase, will not Reductive dissolution of igneous and metamorphic produce Fe isotope fractionations (Johnson et al., rocks under an anoxic atmosphere, which has been 2004). Preferential, but complete, dissolution of traditionally thought to be important for the soil- different minerals in high-temperature rocks is forming processes on biota-free continents that are unlikely to produce significant Fe isotope fractiona- overlain by an anoxic atmosphere prior to ~2.2 Ga tions because inter-mineral isotopic variations in such (e.g., Holland, 1984, 1994; Rye and Holland, 1998), is rocks are quite small. Although predicted Fe isotope unlikely to produce significant Fe isotope variability. fractionations between various silicate minerals and Fe3+-bearing phases are rare in igneous and meta- oxides is on the order of 0.2x to 0.6x in 56Fe/54Fe morphic rocks, constituting only a few modal percent. ratio at 600–800 8C(Fig. 13; Polyakov and Mineev, The data so far indicate that Fe-oxides have Fe isotope 2000), such large fractionation in samples that compositions that are identical to those of Fe silicates represent igneous magmas have yet to be found (Fig. 13; Beard and Johnson, 2004b). Moreover, experimental data indicate that inorganic reduction 3+ 2+ of Fe -oxides to Feaq , utilizing H2 or other reductant, is extremely slow at temperatures below Hornblende ~250 8C(Kishima and Sakai, 1984). Biotite Non-stoichiometric, incongruent dissolution of 2+ Magnetite Fe -bearing silicate minerals such as hornblende, in Olivine the presence of organic ligands such as acetic acid, oxalic acid, citric acid, and siderophores do produce aqueous Fe that has d56Fe values up to 1x lower in 56Fe/54Fe ratios (Brantley et al., 2001). However, these relatively large isotopic fractionations are produced only at very small extents of dissolution (b0.1%), and virtually all aqueous Fe was Fe3+ in these experiments, reflecting the fact that ligands with a high affinity for Fe3+ were used. Similarly large Olivine Diopside Magnetite fractionations were observed for soil extractions, where less than 0.1% extraction of total Fe occurred. -0.2 0 0.2 0.4 0.6 0.8 Although the results of Brantley et al. (2001, 2004) δ56Fe [‰] have important implications for the sources and isotopic compositions of dissolved riverine Fe in Fig. 13. Iron isotope compositions of coexisting high-temperature modern (oxic) environments, it remains unclear if they rock-forming minerals. Data measured for olivine, magnetite, may serve as a guide for the isotopic effects expected biotite, and hornblende from four volcanic rocks show no under terrestrial weathering conditions in an O2-poor significant fractionations (data from Beard and Johnson, 2004b), atmosphere. If riverine Fe fluxes were small in the suggesting that differential weathering will not produce significant Archean, they would have had little effect on the Fe isotope fractionation, assuming weathering occurs through congruent dissolution of Fe-bearing phases. Black bars at bottom isotopic compositions of an Fe-rich ocean, such as are predicted isotopic compositions (calculated for 600–800 8C) may have existed in the Archean (Canfield et al., from Polyakov and Mineev (2000). 2000). 160 K.E. Yamaguchi et al. / Chemical Geology 218 (2005) 135–169

2+ It is possible, however, that riverine Feaq fluxes show that altered basalt sections generally have would have been greater in the Archean than today, positive d56Fe values, up to +1.3x (Rouxel et al., assuming the Archean atmosphere was relatively 2003), and these rocks may represent the isotopic 2+ 56 anoxic. Assuming dissolution and transport of Feaq complement to the low d Fe values that have been 2+ and Mgaq was similar under an anoxic atmosphere measured for hydrothermal fluids. Sulfide minerals at 2+ 56 (Holland, 1984), a minimum riverine Feaq flux in active hydrothermal systems can have d Fe values an anoxic Archean would have been 1.8Â1014 g/ that are both higher and lower than those of hydro- year (Holland, 1978; Holland and Petersen, 1995). thermal fluids (Rouxel et al., 2004), and the role that 2+ Holland (1984) suggests an Feaq content for modern sulfide precipitation plays in determining the Fe mid-ocean ridge (MOR) hydrothermal fluids of 100 isotope compositions of hydrothermal Fe remains 2+ ppm, which produces a MOR Feaq flux of unclear (Beard and Johnson, 2004a). Although the 12 2+ 5.0Â10 g/year. The riverine Feaq flux under an details by which Fe isotope fractionations occur in anoxic Archean atmosphere, therefore, may have marine hydrothermal systems are not yet well under- 2+ been ~36 times larger than the modern MOR Feaq stood, it appears that the major controls on the Fe 2+ flux. The MOR Feaq flux may have been higher isotope composition of fluids that are exhaled to the than today if heat flow was higher, but this may have modern deep ocean include leaching of Fe and sulfide 2+ been offset by greater riverine Feaq contents in the precipitation under relatively anoxic and high-temper- Archean if pCO2 was high and weathering rates ature conditions, as well as lower-temperature sulfide were greater (Ohmoto et al., 2004). If there is little precipitation and formation of Fe3+-rich alteration 2+ 2+ Fe isotope fractionation between Feaq and Fe - zones in oceanic crust (Rouxel et al., 2003, 2004). If bearing minerals during large extents of dissolution, the deep oceans were less oxic in the past than they as suggested by extrapolation of the data of Brantley are today (Canfield, 1998; Canfield and Raiswell, et al. (2001, 2004) to high extents of dissolution, 1999; Canfield et al., 2000; Shen et al., 2003; Arnold then we might expect the d56Fe values of the et al., 2004), we might expect the proportion of high- Archean oceans to be closer to zero than modern d56Fe, Fe3+-rich altered oceanic crust to be less, which MOR fluids, assuming the atmosphere was relatively would be expected to shift the Fe isotope composition anoxic. of the hydrothermal Fe flux to the oceans to less negative d56Fe values. 5.2. Isotopic effects of Fe cycling in marine In addition to processes in oceanic hydrothermal environments systems, diagenetic processes in Phanerozoic marine sediments may produce significant variations in Fe The isotopic composition of Fe delivered to the isotope compositions for Corg-rich, relatively anoxic modern oceans varies greatly depending upon the environments. Variations in d56Fe values up to 1x are 3+ 3+ pathway. Fe -oxide or Fe -rich clays produced by observed in the Corg-rich Kimmeridge Clay Formation terrestrial weathering, delivered to the oceans through of Jurassic age (Matthews et al., 2004), which are suspended river loads or aerosols and eolian transport, interpreted to reflect Fe cycling through several have d56Fe values close to zero (Beard et al., 2003a). pathways, including DIR of Fe3+-oxide/hydroxide There is an indication that the d56Fe values of minerals, as well as formation of Fe-rich dolostones dissolved Fe from riverine input are low (e.g., Fantle and pyrite from Fe2+-bearing diagenetic fluids. Sim- et al., 2004; Ingri et al., 2004), consistent with the low ilar ranges in d56Fe values are found in Cretaceous d56Fe values measured for the exchangeable compo- (Cenomanian–Turonian) black shales (K. Yamaguchi nent of soils (Brantley et al., 2001, 2004). The major et al., unpub. data). Larger isotopic variations are flux of Fe to the deep oceans comes from hydro- found in modern marine sediments that are under- 56 2+ thermal systems such as those in mid-oceanic ridge, going active diagenesis, where d Fe values for Feaq back-arc spreading, or ocean island settings, and these and diagenetic minerals such as iron monosulfide and typically have negative d56Fe values, between À0.6x pyrite, vary by up to 4x, and are interpreted to reflect and À0.2x (Sharma et al., 2001; Beard et. al., biological redox cycling of Fe and S (Severmann et 2003a). Studies of Mesozoic sections of oceanic crust al., 2004b). These results demonstrate that significant K.E. Yamaguchi et al. / Chemical Geology 218 (2005) 135–169 161

Fe isotope variations may be found in diagenetic estimated the equilibrium fractionation factor 2+ 2+ environments that are relatively reducing, indicating between [Fe (H2O)6] –FeCO3 at 20 8Ctobe that they may be effectively decoupled from the +0.5x, based on precipitation experiments in abiotic terrestrial weathering processes that occur in con- systems, which is 1.6x lower than the predicted nection to high atmospheric O2 contents. fractionation. Johnson et al. (2004) estimated that the equilibrium fractionation factor between Fe2+ and 5.2.1. Origin of carbonate-rich suite FeCO3 during DIR is ~0x, whereas that of between 2+ Mineral–fluid fractionation factors allow calcula- Fe and Ca-substituted siderite (Ca0.15Fe0.85CO3)is 56 2+ tion of the d Fe values for Feaq that may have estimated to be +0.9x. These experimental results been in equilibrium with the carbonate-rich suites, confirm the sensitive nature of mineral–fluid fractio- assuming that the measured bulk-rock Fe isotope nation factors to Ca substitution in Fe carbonates. 56 2+ compositions are close to those of the end-member Calculated d Fe values for Feaq in equilibrium carbonate. This assumption is reasonable in terms of with the carbonate-rich units vary greatly, from the Fe mass balance for the siderite-rich shales of the À2.5x to +1x, using the experimentally measured 3.25 Ga Sheba Formation, and Fe-dolomite units of mineral–fluid fractionation factors (Fig. 14). It is not the 2.64 Ga Black Reef Formation, 2.60 Ga yet clear if there is a significant difference in the 2+ Wittenoom Dolomite Formation, and 2.60 Ga Car- Feaq –Fe carbonate fractionation factors between awine Dolomite Formation, as well as pure siderite biotic and abiotic systems, given the markedly differ- layers in the 2.5 Ga Kuruman Iron Formations from ent experimental approaches in the studies of Johnson the Kaapvaal Craton that have been analyzed by et al. (2005) and Wiesli et al. (2004). The majority of 56 2+ Johnson et al. (2003). the calculated d Fe values for Feaq lie between 2+ Fractionation factors between Feaq and Fe À0.5x and +0.5x, which overlap the isotopic carbonate may be obtained from the work of compositions of modern hydrothermal fluids from Polyakov and Mineev (2000) and Schauble et al. mid-oceanic ridges (MOR), where d56Fe values (2001), who calculated reduced partition function between À0.6x and À0.2x have been measured to ratios for a variety of Fe-bearing fluids and minerals date (Sharma et al., 2001; Beard et al., 2003a). Given based on spectroscopic data. The predicted fractio- the possibility that the Fe isotope compositions of nation factors vary greatly, and are a strong function hydrothermal fluids in the Archean may have been of carbonate stoichiometry. For example, the fractio- shifted to slightly higher d56Fe values due to a 2+ 2+ 3+ nation between [Fe (H2O)6] and FeCO3 at 20 8C decrease in Fe -bearing alteration zones (see dis- is predicted to be +2.12x for 56Fe/54Fe (Polyakov cussion above), our preferred explanation for the 56 2+ and Mineev, 2000; Schauble et al., 2001). In contrast, calculated d Fe values for Feaq for the majority of 2+ 2+ the [Fe (H2O)6] –CaMg0.5Fe0.5(CO3)2 fractionation the carbonate units is that this most likely represents 2+ 2+ 2+ is predicted to be +3.65x, and the [Fe (H2O)6] – the Fe isotope composition of ambient seawater Fe . Ca1.1Mg0.5Fe0.3Mn0.1(CO3)2 fractionation is pre- An important remaining question is the role that dicted to be +2.65x. It is unknown what the variable carbonate stoichiometry may play in explain- 56 2+ mineral–fluid fractionation is for Fe-bearing dolomite ing the range in calculated d Fe values for Feaq , (bferroan dolomiteQ); however, based on the work by which we suspect will be substantial. Polyakov and Mineev (2000) and Schauble et al. (2001), it seems likely to be distinct from that of Fe- 5.2.2. Origin of magnetite-rich suite 56 2+ rich carbonates such as siderite and ankerite. The d Fe values for Feaq that was in equilibrium Experimental measurement of fractionation fac- with the magnetite-rich suites (samples from the 2.69 2+ tors between Feaq and Fe carbonate at low temper- Ga Jeerinah Formation, ~2.9 Ga Rietkuil Formation, atures (~20 8C) in abiologic and biotic systems and 2.96 Ga Parktown Formation, as well as pure (Johnson et al., 2005; Wiesli et al., 2004), as well as magnetite separated from the 2.5 Ga Kuruman Iron estimates based on natural assemblages (Johnson et Formation) may be calculated using the experimen- 2+ al., 2003), suggests that the predicted fractionations tally determined fractionation factor between Feaq are significantly too large. Wiesli et al. (2004) and magnetite (À1.3x; Johnson et al., 2005). Low 162 K.E. Yamaguchi et al. / Chemical Geology 218 (2005) 135–169

2.50 Ga Kuruman Iron Fm. (Johnson et al., 2004, 2005). Johnson et al. (2004, (siderite separates) 2005) presented arguments that magnetite layers in the 2.5 Ga Kuruman Iron Formation reflected a 56 2+ mixture between a low-d Fe Feaq component that 2.60 Ga 2+ Carawine was produced by DIR and an Feaq component Dol. Fm. that was derived from marine hydrothermal sources that had higher d56Fe values. Magnetite is a common end product in DIR experiments (e.g., 2.60 Ga Wittenoom Dol. Fm. Lovley et al., 1987), and Fe isotopes provide a means to identify such magnetite in the rock record. The results of the current study provide additional evidence of DIR in Archean rocks, beyond that 2.64 Ga Black Reef Fm. suggested for banded iron formations (Johnson et al., 2005). Nevertheless, in the following section we explore possible abiotic processes that may produce 56 2+ 3.25 Ga Sheba Fm. (shales) large quantities of low-d Fe Feaq .

5.3. Abiological processes for producing low-d56Fe 2+ Feaq -3 -2 -1 0 1 1.5 3+ 56 δ FeFe2+aq [‰] Other than biological reduction of Fe -(hydr)ox- 2+ ides, processes that could have produced Feaq that 56 Fractionation factors used: had very negative d Fe values in these rocks include:

2+ Biological, Fe aq – Ca-subsituted siderite (+0.9 ‰) 2+ Nonbiological, Fe aq – siderite (+0.5 ‰) 2.50 Ga Kuruman Iron Fm. 2+ Biological, Fe aq – siderite (0 ‰) (magnetite separates)

2+ Fig. 14. Calculated Fe isotope compositions of Feaq that would have been in equilibrium with the Fe carbonates from the carbonate-rich 2.69 Ga Jeerinah Fm. Archean–Paleoproterozoic samples of this study. For each of the 2+ five groups of samples, three Feaq –Fe carbonate fractionation 2+ factors are used: bottom, Feaq –siderite fractionation of 0x, based 2+ on DIR experiments of Johnson et al. (2005); middle, Feaq –siderite fractionation of +0.5x, based on abiotic precipitation experiments ~2.9 Ga Rietkuil Fm. 2+ of Wiesli et al. (2004); top, Feaq –Ca-substituted siderite fractionation of +0.9x, based on DIR experiments of Johnson et al. (2005). Vertical bar represents range in d56Fe values of igneous rocks.

2.96 Ga Parktown Fm.

56 2+ d Fe are calculated for Feaq in equilibrium with the magnetite-rich units (Fig. 15). The majority of the 56 2+ -3 -2 -1 0 1 1.5 calculated d Fe Feaq values are significantly lower 56 than those measured for modern hydrothermal fluids δ FeFe2+aq [‰] from MOR settings (À0.6x to À0.2x; Sharma et al., 2+ 2+ 2001; Beard et al., 2003a). Aqueous Fe that has low Fig. 15. Calculated Fe isotope compositions of Feaq that would have d56Fe values seems most likely to have been produced been in equilibrium with magnetite for the four magnetite-rich groups of this study. Calculations use a Fe2+–magnetite fractiona- by biological reduction of Fe3+(OH) during dia- aq 3 tion of À1.3x, based on the DIR experiments of Johnson et al. genesis of sediments, based on the Fe isotope (2005). Vertical bar represents range in d56Fe values of igneous fractionation factors that are currently available rocks. K.E. Yamaguchi et al. / Chemical Geology 218 (2005) 135–169 163

2+ (i) coupled redox reactions among metals such as Fe measured Feaq –FeS fractionations between +0.3x and Mn; (ii) Rayleigh fractionation involving precip- and +0.9x, suggesting reduction by dissolved sulfide itation of Fe3+ hydroxides; (iii) reductive dissolution that produced FeS as a product might be a means to of Fe3+ hydroxides by dissolved sulfide. It is possible, produce limited quantities of relatively high-d56Fe 3+ 2+ yet unknown, if reduction of Fe -bearing phases by Feaq . In addition, pyrite from modern and ancient 2+ 2+ 56 56 Mn may produce Feaq that has low d Fe values, sedimentary environments generally have low d Fe but from a mass-balance perspective, there is insuffi- values (Johnson et al., 2003; Matthews et al., 2004; cient Mn to reduce a significant proportion of Fe in Severmann et al., 2004b), suggesting that pyrite 2+ the majority of the rocks studied here; Mn contents formation should also produce Feaq that has rela- range from 0.1 to 1 wt.% MnO, which is much less tively high d56Fe values. Regardless of the uncertain- than those of Fe (~2 to 30 wt.% total Fe as Fe2O3). ties in fractionation factors or the relative balance of 3+ 2+ 2+ Precipitation of Fe oxides upon oxidation of Feaq dissolved sulfide and Feaq , the relatively low sulfide during Rayleigh fractionation could produce a resid- contents of many of the rocks in this study, 2+ 56 ual Feaq pool that has low d Fe values, given particularly in the magnetite- and carbonate-rich rocks 3+ 2+ 3+ 56 measured Fe oxide/hydroxide–Feaq and Feaq – that are used to infer the d Fe values for ancient 2+ 2+ Feaq fractionation factors (Bullen et al., 2001; Feaq , suggests that high ambient dissolved sulfide Johnson et al., 2002; Welch et al., 2003). Oxidation contents were not present during formation of most of 2+ of Feaq in the Archean might have occurred through the rocks studied here. interaction with O2-bearing surface waters or UV- photo oxidation (e.g., Braterman and Cairns-Smith, 5.4. Significance of biological Fe reduction in the 1983). In addition, oxidation through anoxygenic Archean oceans photosynthesis (Widdel et al., 1993)mayhave 3+ produced Fe oxides under O2-poor conditions, and If we are correct in our interpretation that the low- 2+ 3+ 56 2+ Feaq –Fe oxide isotopic fractionations for such a d Fe values that are calculated for Feaq from several process are generally similar to those produced by rock suites, a bio-available Fe3+ pool such as 3+ abiotic oxidation by O2 (Croal et al., 2004). Assuming amorphous Fe -oxyhydroxide (Lovley and Phillips, 2+ 3+ aFeaq –Fe oxide fractionation of À0.9x to À1.5x 1987; Lovley, 1991; see Fig. 16) must have been (Bullen et al., 2001; Croal et al., 2004), and an initial present. We were not certain at this stage whether 56 2+ 3+ d Fe value for Feaq of 0x, 50 to 70% precipitation these Fe hydroxides were produced during weath- would be required to produce a d56Fe value of À1.0x ering of rocks on land or chemically precipitated from 2+ 56 2+ for the residual Feaq . To produce the very low d Fe Fe -rich seawater (Fig. 16). We are also not certain 2+ 3+ values inferred for Feaq from some of the samples, whether the formation of Fe -(hydr)oxides was N95% precipitation is required. In the absence of caused by aerobic reactions or anaerobic reactions evidence for the required complementary high-d56Fe (Fig. 16). Finally, it remains unknown if evidence for oxide component in the sedimentary sequences DIR reflects a global or local processes in the Archean studied, an abiotic oxidation–precipitation model oceans. requires physically separate regions of oxidation and This study suggests that DIR might have already deposition. If the Archean oceans were relatively been an important pathway for respiration for life as Fe2+-rich, such a model seems an implausible means early as 2.9 Ga, based on the low d56Fe values of the 56 2+ by which large quantities of low-d Fe Feaq may be magnetite-rich Rietkuil Formation. The early occur- produced. rence of microbes capable of DIR suggested in this Reductive dissolution of Fe3+ oxides by sulfide study is consistent with its relatively deep root in the (e.g., Poulton, 2003) is a possible means that low- btree of lifeQ (Lovley, 1991, 1993; Nealson and 56 2+ d Fe Feaq could be produced, although the very low Saffarini, 1994). At the very least, the large range in 56 solubility of Fe sulfides would limit the amount of d Fe values for magnetite- and Corg-rich Archean 2+ Feaq that would be produced. No experimental rocks reported here provide strong support for redox studies have yet investigated Fe isotope fractionation cycling of Fe, given the fact that some of the largest during such processes, although Butler et al. (2003) Fe isotope fractionations exist between Fe3+ and Fe2+ 164 K.E. Yamaguchi et al. / Chemical Geology 218 (2005) 135–169

Atmospheric O near 2 UV-oxidation surface ocean (?)

Anoxygenic anaerobic Fe from continents Fe-oxidizing Atmosphere photosynthesizers

δ 56Fe detrital Fe3+ Fe2+ = 0 [‰] Ocean If deep oceans Detrital were anoxic

Dissolved O 2 ? Fe2+ Fe3+ δ 56Fe(II) Fe-reducing bacteria MOR ∆ 56Fe = –1.3 [‰] = –0.5 [‰] Burial Upwelling Fe2+ Fe2+

Submarine hydrothermal systems

Fig. 16. Schematic model for atmosphere–ocean environments showing potential pathways for Fe cycling in the Archean. The major source for 2+ 56 Feaq to the oceans is assumed to be from submarine hydrothermal systems, where a Fe value of À0.5x is assumed, based on the 2+ compositions measured in modern systems. Upwelling Feaq may be oxidized through interaction with atmospheric O2 or, under low-O2 conditions, through UV photo oxidation or anoxygenic photosynthesis. In addition, oxidative terrestrial weathering would produce Fe3+ phases, which, in this case, is expected to have d56Fe values equal to those of igneous rocks. Production of the low-d56Fe values that are calculated for 2+ Feaq from the magnetite-rich rocks of this study (Fig. 15) is thought to be produced by DIR in the upper portions of the sediment column in marine basins. species, as predicted from spectroscopic data (Poly- have d56Fe values that are close to zero, similar akov and Mineev, 2000; Schauble et al., 2001; to the Fe isotope compositions of Phanerozoic Corg- Anbar et al., 2005) or measured in experiments poor clastic sedimentary rocks. This observation (Johnson et al., 2002; Welch et al., 2003). If the suggests that sedimentary transport, diagenesis, and Archean oceans were reducing and relatively rich in lithification of clastic Corg-poor sediments since the 2+ Feaq that was derived from submarine hydrothermal Archean has produced little net Fe isotope fractio- sources, as suggested by many researchers (Canfield nation in bulk rock samples, despite possibly large et al., 2000; Beukes and Klein, 1992), a very changes in redox conditions on the surface of the significant level of DIR activity would have been Earth. In contrast, samples that contain significant required to dominate the Fe isotope composition of magnetite and/or siderite have a wide range in 2+ 56 Feaq relative to the probably significantly higher d Fe values, reflecting diagenetic processes in the 56 2+ d Fe values of the hydrothermal sources. Alterna- presence of Feaq that had Fe isotope compositions 56 2+ tively, the low-d Fe Feaq component may reflect that were substantially different that the bulk crustal local conditions in restricted basins, or even pore baseline. water compositions rather than a global composition Iron isotope compositions of Archean siderite-rich for Archean seawater. rocks that have slightly negative d56Fe values (À0.5F0.5x), and the calculated Fe isotope compo- 2+ sitions of Feaq that may have coexisted with the 6. Conclusions diagenetic minerals that were produced in these rocks overlaps those of submarine hydrothermal fluids. The Fe isotope compositions of most Corg-poor, Although their remain uncertainties in the stoichio- 2+ Archean–Paleoproterozoic samples studied here metric effects on Feaq –Fe carbonate fractionation K.E. Yamaguchi et al. / Chemical Geology 218 (2005) 135–169 165 factors, to a first order it appears that the d56Fe values isotope work was done as a post-doctoral researcher at for MOR hydrothermal systems may have been the University of Wisconsin-Madison. [LW] similar in the Archean as compared to today. Iron isotope compositions of the Corg-rich and magnetite-rich samples of 2.9 to 2.6 Ga in age have References d56Fe values that range from À2.5x to À1.0x, 2+ suggesting that the Fe -bearing minerals (siderite, Albare`de, F., Beard, B.L., 2004. Analytical methods for non- magnetite, and pyrite) in these rocks formed by traditional isotopes. In: Johnson, C.M., Beard, B.L., Albare`de, reduction of Fe3+-(oxy)hydroxides, based on the F. (Eds.), Geochemistry of Non-Traditional Stable Isotopes, fact that the d56Fe values calculated for Fe 2+ that Reviews in Mineralogy and Geochemistry, vol. 55, pp. 113–152. aq Anbar, A.D., 2004. 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Because DIR requires a process to revisited: a single zircon microprobe study. Earth Planet. Sci. produce Fe3+ (oxy)hydroxides, our results provide Lett. 101, 90–106. Armstrong, R.A., Compston, W., Retief, E.A., Williams, I.S., strong evidence for extensive biogeochemical redox Welke, H.J., 1991. Zircon ion microprobe studies bearing on cycling of Fe in the Archean. the age and evolution of the Witwatersrand triad. Precambrian Res. 53, 243–266. Arndt, N.T., Nelson, D.R., Compston, W., Trendall, A.F., Thorne, Acknowledgements A.M., 1991. The age of the Fortescue Group, Hamersley Basin, western Australia, from iron microprobe zircon U–Pb results. Aust. J. Earth Sci. 38, 261–281. Comments by Alan Matthews, Nic Beukes, and an Arnold, G.L., Anbar, A.D., Barling, J., Lyons, T.W., 2004. anonymous reviewer are appreciated. We thank the Molybdenum isotope evidence for widespread anoxia in mid- editors of the bisotope biosignaturesQ special issue, A. Proterozoic oceans. Science 304, 87–90. Matthews and J. Horita, for giving us the opportunity Beard, B.L., Johnson, C.M., 1999. High-precision iron isotope measurements of terrestrial and lunar materials. Geochim. to present our study. Access to drill cores was Cosmochim. Acta 63, 1653–1660. provided by W.E.L. Minter, N. Beukes, M. Barley, Beard, B.L., Johnson, C.M., 2004a. Fe isotope variations in the B. Krapez, Anglo American Prospecting Service, modern and ancient Earth and other planetary bodies. In: Sheba Mine, Agnes Mine, and CRA Exploration. Johnson, C.M., Beard, B.L., Albare´de, F. (Eds.), Geochemistry We thank D. Walizer, T. Kakegawa, Y. Watanabe, H. of Non-Traditional Stable Isotopes, Reviews in Mineralogy and Geochemistry, vol. 55, pp. 319–357. Gong, K. Bickle, R. Wilkin, S. Welch, R, Poulson, Beard, B.L., Johnson, C.M., 2004b. Inter-mineral Fe isotope and R. Wiesli for valuable assistance with the variations in mantle derived rocks and implications for the Fe analytical work. This study was supported by geochemical cycle. Geochim. Cosmochim. Acta 68, 825–837. NASA-Ames Research Center and the NASA Astro- Beard, B.L., Johnson, C.M., Cox, L., Sun, H., Nealson, K.H., 1999. biology Institute through Joint Research Initiative JRI Iron isotope biosignatures. Science 285, 1889–1892. Beard, B.L., Johnson, C.M., Von Damm, K.L., Poulson, R.L., NCC 2-5449; additional support was provided by the 2003a. Iron isotope constraints on Fe cycling and mass balance NASA Astrobiology Institute (NCC 2-1057), NASA in oxygenated Earth oceans. Geology 31, 629–632. Exobiology Program (NAG 5-9089), National Sci- Beard, B.L., Johnson, C.M., Skulan, J.L., Nealson, K.H., Cox, L., ence Foundation (EAR 97-06279, EAR 02-29556), Sun, H., 2003b. Application of Fe isotopes to tracing the and the Japanese Ministry of Education, Culture, geochemical and biological cycling of Fe. Chem. Geol. 195, 87–117. Sports, Science and Technology (#100411004). 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