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Chemical Geology

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Invited research article

A multistage origin for Neoarchean layered hematite-magnetite ironformation from the Weld Range, Yilgarn Craton, Western Australia

Andrew D. Czajaa,b,⁎, Martin J. Van Kranendonkc,d, Brian L. Beardb,e, Clark M. Johnsonb,e

a Department of Geology, 500 Geology–Physics Bldg., University of Cincinnati, Cincinnati, OH 45221-0013, USAbNASA Astrobiology Institute, USAc School of Biological, Earth and Environmental Sciences, Australian Centre for Astrobiology, University of New South Wales, Kensington, NSW 2052, Australiad Australian Research Council Centre for Excellence for Core to Crust Fluid Systems, Australiae Department of Geoscience, 1215 W. Dayton Street, University of Wisconsin, Madison, WI 53706, USA

A R T I C L E I N F O

Editor: Michael E. Böttcher

Keywords:ArcheanEarly EarthIron formationIron oxidationIron isotopesHematiteMagnetitePyrite

A B S T R A C T

The origins of iron formations remain somewhat enigmatic despite much progress over the past decades. This isdue in part to continuing research demonstrating that these deposits do not have a single origin but rather can beformed under various and variable conditions.

This study describes a formation pathway for an Algoma-type banded iron formation (BIF) from the 2.75billion-year-old Weld Range of Western Australia, based on petrographic and Fe isotope analyses of drill coresamples. The BIF is composed of alternating layers of jaspilitic and Fe-poor chert, magnetite layers of varyingthickness and abundance, and rare pyrite. Petrographic analysis indicates the minerals formed in the orderhematite, magnetite, and then pyrite, with the latter two clearly replacive of “primary” bedded hematite.

Hematite δ56Fe values from all portions of the drill core are positive and essentially identical(δ56Fe= 0.61 ± 0.05‰). Magnetite δ56Fe values are positive, as well (δ56Fe= 0.51 ± 0.07‰) but vary sys-tematically with total magnetite content of the core samples. Pyrite δ56Fe values are positive and higher thanboth hematite and magnetite (δ56Fe= 1.03 ± 0.05‰).

Combined, these analyses reveal a multistage formational model for the Weld Range BIF. Initial depositionoccurred by the partial oxidation of hydrothermally-sourced aqueous Fe(II) to form a Fe(III)-Si co-precipitate,probably by anoxygenic photosynthetic iron oxidizers. This material would have been converted to hematiteunder equilibrium conditions with excess aqueous Fe(II). Magnetite formed by a later intrusion of reducingfluids, mostly along bedding planes, but also as cross-cutting veins. This produced large quantities of magnetitelayers parallel, and sub-parallel, to the original bedding, but that do not represent a primary depositional feature.Pyrite was produced in some portions of the unit, likely under equilibrium conditions, by later infiltration ofreducing fluids containing sulfide.

The entirely positive δ56Fe values for all Fe phases in the Weld Range BIF stands in stark contrast to the moremassive Superior-type BIFs of the Paleoproterozoic Hamersley and Transvaal basins of Western Australia andSouth Africa, respectively, which have δ56Fe values that average close to 0‰ but vary over nearly the entirerange of δ56Fe values measured in nature (δ56Fe=−2.5 to +2.6‰). The Weld Range BIF δ56Fe values aresimilar to those of older Algoma-type BIFs such as the Paleoarchean Isua BIF of Greenland but are less positive onaverage. This overall trend of decreasing average δ56Fe values through time likely records an overall increasingtrend of oxygenation of the Archean surface ocean from 3.8 to 2.45 billion years ago.

1. Introduction

Iron formations (IFs) are important geologic features of thePrecambrian Earth, in large part because they represent iron depositionand cycling on a scale that does not occur on the surface of the modernEarth, but also because they have played a major role in inferring

chemical pathways in the ancient atmosphere–biosphere–hydrospheresystem. All models for IF genesis infer deposition in a marine settingand a role for Fe-rich fluids from anoxic bottom waters, magmaticfluids, or hydrothermal fluids (Klein, 2005; Bekker et al., 2010;Konhauser et al., 2017). The depositional and tectonic settings involvedin IF deposition have varied, as does the dominant Fe-bearing

https://doi.org/10.1016/j.chemgeo.2018.04.019Received 9 October 2017; Received in revised form 14 April 2018; Accepted 16 April 2018

⁎ Corresponding author at: Department of Geology, 500 Geology–Physics Bldg., University of Cincinnati, Cincinnati, OH 45221-0013, USA.E-mail address: andrew.czaja@uc.edu (A.D. Czaja).

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mineralogy (James, 1954). Gross (1965) highlighted several types ofIFs, including volcanogenic IFs (Algoma-type), and IFs deposited instable continental margins (Superior-type), where the latter is re-presented by the well-known banded iron formations (BIFs) of the2.45 Ga Hamersley Basin in the Pilbara Craton of Western Australia,and the 2.45 Ga Griqualand West Basin of the Kaapvaal Craton in SouthAfrica. Konhauser et al. (2017) provide a recent thorough review of ironformation depositional pathways.

Iron isotope geochemistry potentially provides a direct measure-ment of iron pathways involved in IF genesis (e.g., Johnson et al., 2003;Dauphas et al., 2004; Dauphas et al., 2007a, 2007b; Frost et al., 2007;Whitehouse and Fedo, 2007; Johnson et al., 2008b; Hyslop et al., 2008;Planavsky et al., 2009; Rouxel et al., 2005; Heimann et al., 2010;Craddock and Dauphas, 2011; Czaja et al., 2013; Li et al., 2015;Steinhoefel et al., 2009b; Steinhoefel et al., 2010). The> 3.7 Ga Isuaand Nuvvuagituq IFs of Greenland and Quebec, respectively, are in-terpreted as Algoma-type volcanogenic IFs and have δ56Fe values thatare generally positive and vary over a relatively narrow range. Simi-larly, the jaspilitic chert of the c. 3.46 Ga Marble Bar Chert (MBC)Member of the Pilbara Craton, Australia, have δ56Fe values that varyover a narrow positive range (Li et al., 2013a), although the MBC is notinterpreted to be a volcanogenic BIF. The Fe in these Paleoarchean orolder units is interpreted to have formed by partial oxidation of hy-drothermally-sourced aqueous Fe(II), possibly by oxygenic or anoxy-genic photosynthesis (Dauphas et al., 2004; Dauphas et al., 2007a,2007b; Whitehouse and Fedo, 2007; Czaja et al., 2013; Li et al., 2013a).Early Neoarchean Algoma-type IFs from Zimbabwe tend to also havepositive δ56Fe values for Fe-oxides that vary over a narrow range andare interpreted to have formed via oxide precipitation, likely by an-oxygenic photosynthesis followed by diagenetic alteration during earlymetamorphism (Steinhoefel et al., 2009b). Strikingly, the ca. 2.45 GaSuperior-type BIFs from the Hamersley (e.g., Dales Gorge Member ofthe Brockman Iron Formation) and Griqualand West basins (e.g.,Kuruman Iron Formation) contain a broad range of δ56Fe values,

including highly negative values (Johnson et al., 2003; Rouxel et al.,2005; Johnson et al., 2008b; Heimann et al., 2010; Steinhoefel et al.,2010; Planavsky et al., 2012; Li et al., 2013b; Li et al., 2015) similar tothose of contemporaneous marine sedimentary materials, includingblack shales, carbonates, and sedimentary sulfides (Rouxel et al., 2005;Yamaguchi et al., 2005; Archer and Vance, 2006; Czaja et al., 2010;Czaja et al., 2012). The appearance of negative δ56Fe values could ei-ther record the emergence of Fe cycling by dissimilatory iron-reducingbacteria (Johnson et al., 2003; Johnson et al., 2008a, 2008b;Steinhoefel et al., 2010; Heimann et al., 2010; Craddock and Dauphas,2011), or they could reflect extensive oxide precipitation (Rouxel et al.,2005; Anbar and Rouxel, 2007; McCoy et al., 2017). In contrast, 1.9 GaSuperior-type IFs in the Lake Superior region of North America (e.g.,Gunflint and Biwabik formations) tend to have positive δ56Fe values,also similar to contemporaneous marine sedimentary rocks. These havebeen broadly interpreted to record partial oxidation of aqueous Fe(II),possibly linked to photosynthesis (Frost et al., 2007; Hyslop et al., 2008;Planavsky et al., 2009). The debate between a biological or abiologicalorigin for the stable Fe isotope signals recorded in IFs and relatedmarine sedimentary rocks has been hampered by the lack of detailedstudies of IFs for which there is no likely role for biology, including insitu analysis that can determine, on a mineral scale, which mineralsmay have been in Fe isotope equilibrium in an abiological system.

Here, we report Fe isotope data from three different Fe-bearingminerals in a 2.75 Ga Algoma-type BIF from the Weld Range ofAustralia that were collected from monomineralic layers or by in situlaser-ablation analysis of individual grains. These data, along with adetailed understanding of the paragenetic sequence of mineral forma-tion, allow development of a robust IF depositional model that is dis-tinct from models proposed for other Precambrian IFs. It provides abaseline with which to evaluate proposals of biological pathways forIFs. Because the Weld Range BIF is of the Algoma-type, this study offersan opportunity to study Fe isotope variations in a BIF system that waslikely dominated by volcanic hydrothermal processes and demonstrates

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Fig. 1. Geologic map and stratigraphic scheme for the Murchison Domain of the Yilgarn Craton, Western Australia. The left panel shows the location of the WeldRange iron formation studied here and the location of the borehole WRRD1128 from which the drill core samples studied here were collected (26°59′01″S,117°34′37″E). The right panel shows the stratigraphic relationship of the supracrustal units of the Murchison Supergroup (including the Wilgie Mia Formation thatcontains the Weld Range BIF) and the granitic and mafic-ultramafic rocks. Patterns are the same as used in the left panel.This figure was modified from Van Kranendonk et al. (2013).

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Fig. 2. Field, hand sample, and thin section photos of the c. 2.75 Ga Weld Range BIF from samples of drill core WRRD1128. A) Field photo of weathered Weld RangeBIF, showing mesoscale layering typical of this unit. The white bar near the middle of the image is the same scale that appears in panel B. B) Field photo of a freshsurface of Weld Range BIF, showing alternating layers of jaspilite (red) and magnetite (gray). Note the highly irregular nature of the magnetite layers. C) Drill core ofthe Weld Range BIF showing alternating layers of jaspilite (pink/white) and magnetite (gray). Camera lens cap is 4 cm diameter. Location of panel D shown in whiterectangle. D) Close-up of part of panel C, showing the slightly discordant nature of a broad magnetite layer (top and base of arrow), indicating the secondary nature ofmagnetite precipitation and possible inflation of primary jaspilite bedding by infiltrating fluids. E) Transmitted light photograph of Weld Range BIF thin sectionshowing alternating parallel layers of jaspilite and magnetite (sample WR5). F) and G) Transmitted light thin section images showing evidence that the magnetiterepresents a secondary replacement feature of bedding (panel F – sample WR4) and in cross-cutting veins (panel G – sample WR6). H) Transmitted light thin sectionimage of a region composed principally of “primary” jaspilite. Note the two unconformities between packages of jaspilite that contain distinct textures, including thedarker red portion in upper image, and the paler jaspilite at image bottom. Note also the delicate nature of the “primary” jaspilitic bedding (shown in close-up inpanel I) and the presence of magnetite along the lower unconformity surface and within the upper, dark red package. I) Close-up of part of panel H, showing mm-scale “primary” jaspilite bedding characterized by “thundercloud-like” tops and relatively flatter bottoms, and finer-scale internal layering.

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the complexity of depositional and post-depositional Fe cycling in theArchean.

2. Geologic setting, materials, and methods

Iron formation samples were collected from diamond drill core(WRRD1128) recovered from the 2.75 Ga Wilgie Mia Formation of theMurchison Supergroup from the Weld Range, western YouanmiTerrane, Yilgarn Craton, Western Australia (Fig. 1; Van Kranendonket al., 2013). Samples came from three stratigraphic positions: sampleWR4, 227.57–227.87m (GSWA sample 197433); sample WR6,231.65–231.95m (GSWA sample 197435); and sample WR5,236.4–236.7m (GSWA sample 197434). Iron formations (IFs) in thisunit conformably overlie subaerially deposited volcaniclastic rocks andare interpreted to be part of a collapsed felsic volcanic caldera system(Van Kranendonk et al., 2013). These units were regionally metamor-phosed to ~320 °C (Gole, 1980; Watkins and Hickman, 1990). In out-crop, the Weld Range BIFs consist of centimetric-layered hematite-magnetite rocks, with a siliceous matrix (Fig. 2A and B). Hand samplesare composed of layers of finely bedded, fine-grained, jaspilitic chertand medium-grained magnetite (Fig. 2), with the local occurrence ofpyrite in one of the three samples (Fig. 3). In the drillcore, pyrite be-comes more common at depth. The samples studied here correspond tothe least altered Beebyn (Weld Range BIF) samples described byDuuring et al. (2018).

In core samples, despite the overall similar appearance of the al-ternating layers of hematite and magnetite to those of previously re-ported BIFs (e.g., Fig. 2B), the Weld Range BIF records evidence of adifferent formation pathway. The jaspilite shows clear evidence forprimary bedding features in many parts of the studied core (i.e., in-ternal unconformities, cyclical jaspilite-chert couplets in planar beddedmaterial, and asymmetrical way-up features from flat bases and bilious“thundercloud” top contacts of jaspilitic laminae; Fig. 2G–I). On theother hand, the magnetite does not ever show these primary features,but rather, the magnetite layers cut fine-scale lamination in the jaspi-litic layers, at high angles to very low angles, and are always visiblysecondary (Fig. 2C–G). Therefore, magnetite replaced fine-grained he-matite layers in discordant veins and in bedding parallel layers thatinflate the original bedding.

Samples of hematite, magnetite, and pyrite were collected with atungsten carbide hand scribe from individual layers of jaspilite ormagnetite, or from mm-scale grains of pyrite (Fig. 3). Magnetite occursin both thin (sub-millimeter) and thick (centimeter to sub-centimeter)layers, both types of which were sampled. Hematite- and magnetite-bearing samples are estimated to be> 90% pure (in terms of Fe mi-neralogy) by this sampling method, but the pyrites have variableamounts of magnetite contamination (e.g., Fig. 3I–K). Thus, laser-ab-lation analyses of pyrites were performed to measure their Fe isotopecomposition, without contamination by microscopic magnetite. Com-parison of these values with those obtained by micro-sample analysisprovides a measure of the level of contamination in each of the wholegrain samples. The powders collected by hand scribe were dissolvedand the Fe purified using ion-exchange chromatography. Samples forlaser ablation were prepared as polished slabs embedded in 1-inchepoxy rounds (cf., Czaja et al., 2013).

All iron isotope measurements were made using a MicromassIsoprobeMC-ICP-MS at the University of Wisconsin, Madison with eitheran Aridus desolvating nebulizer for solution-based isotope analyses(Beard et al., 2003; Albarède and Beard, 2004), or Photon MachinesAnalyte-fs for laser-ablation analyses (e.g., Czaja et al., 2013). In thisstudy, Fe isotope compositions are defined asδ56Fe= [(56Fe/54Fe)sample / (56Fe/54Fe)standard)− 1]×1000 andδ57Fe= [(57Fe/54Fe)sample / (57Fe/54Fe)standard)− 1]×1000, both inunits of per mil (‰), and are reported relative to the average of igneousrocks. The international standard IRMM-014 has a δ56Fe value of−0.09‰ on this scale. For solution-based isotope analyses, samples

were analyzed on multiple days and accuracy was checked using mul-tiple standards (IRMM-014 and two lab standards, J-M Fe and HPS Fe)and nine test solutions. Seven of the test solutions were made by col-lecting the acid used to separate the rest of the elements from Fe on theanion-exchange columns (cf., Czaja et al., 2012), and the other twowere artificial solutions made from high-purity elemental standards. Allsolutions were amended with a measured quantity of Fe of knownisotopic composition (HPS Fe, δ56Fe=0.49‰) and processed throughthe entire purification protocol. These nine test solutions have anaverage δ56Fe value of 0.49 ± 0.03‰ (2 standard deviations). Theoverall external precision of the Fe isotope analyses was 0.07‰ (2standard deviations) based on multiple analyses of standards and testsolutions over the course of the analytical sessions. For laser ablationanalyses, samples were bracketed with analyses of two distinct in-housepyrite standards (Bal-4-13B-1 and Pyr-49-186; Supplementary Table 2).Bal-4-13B-1 has an average δ56Fe value of −1.39 ± 0.09‰ (2-σ) asmeasured by conventional MC-ICP-MS and −1.42 ± 0.20‰ (2-σ) asmeasured by laser ablation analyses over many sessions. Pyr-49-186 hasan average δ56Fe value of 0.00 ± 0.16‰ (2-σ) as measured by con-ventional analyses and −0.04 ± 0.46‰ (2-σ) as measured by laserablation analyses over many sessions.

The amount of magnetite in each sample was estimated usingtransmitted light images of thin sections of each core piece and thesoftware ImageJ (v. 1.50). The jaspilite containing fine-grained hema-tite was translucent in these images (see Fig. 3), so only the magnetiteand pyrite grains appeared opaque. Pyrite only occurred in sample WR6and made up an insignificant portion of the opaque minerals in terms ofareal coverage. Thus, after applying a threshold filter, the percentage ofblack pixels was measured and recorded as the percentage by area ofmagnetite in each core piece.

Analyses of magnetite compositions of grains from the same layersas those that were sampled for isotope analyses were performed using aHitachi S-3400 scanning electron microscope (SEM) in backscattermode, and a Cameca SX51 electron microprobe, both at the Universityof Wisconsin, Madison. The electron probe was used to measure theconcentrations of Fe, Si, Al, Mg, Mn, Ti, and Cr in individual magnetitegrains.

3. Results

Petrographic and geochemical analyses of the Weld Range BIF drillcore indicate that the Fe minerals formed in the order: hematite thenmagnetite then pyrite. Hematite appears as the “most primary” pre-cipitate, displaying delicate bedding textures such as flat bottoms andthundercloud-like tops to mm-scale “beds”, micron-size hematitegrains, and small-scale unconformities between sets of bedded jaspilite-chert couplets (Figs. 2 and 3C) (Trendall and Blockley, 1970; Beukesand Gutzmer, 2008). Previous studies on other IFs have shown thathematite would likely not have been the primary precipitate, but wouldhave formed from dewatering and diagenesis of a poorly-crystalline Fe(III) gel precursor (e.g., Klein, 2005; Sun et al., 2015) and most likely anFe(III)-Si gel precursor (e.g., L. Wu et al., 2011). Low-angle to bedding-parallel cross-cutting relations indicate a secondary origin for themagnetite layers relative to the “primary” jaspilite layers (Fig. 2D; VanKranendonk et al., 2013). Pyrite grains in the Weld Range samples arelarge and angular and are exclusively found to have grown aroundmagnetite grains or layers (Fig. 3I–K) indicating a replacement origin.The near-primary nature of hematite in the jaspilite, and the secondarynature of magnetite and pyrite, is also consistent with the low abun-dance of Fe in the jaspilite (0.9–7.2 wt%; Supplementary Table 1) andhigh abundances of Fe in the magnetite layers (54.5–70.9 wt%; Sup-plementary Table 1), having in some cases an almost stoichiometricconcentration based on the ideal formula of Fe3O4 (72.4 wt%).

Each Fe phase studied from the Weld Range drillcore samples haspositive δ56Fe values that occur over a relatively narrow range(Supplementary Tables 1 and 2, and Fig. 4). Hematite in the bedded

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(caption on next page)

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jaspilite has δ56Fe values that range from 0.54 to 0.72‰ (average of0.61 ± 0.05‰ 1 s.d., n=14; Supplementary Table 1). The magnetitelayers have slightly lower δ56Fe values that range from 0.37 to 0.60‰(average of 0.51 ± 0.07‰ 1 s.d., n=16; Supplementary Table 1). Thesampled pyrite grains have δ56Fe values that range from 0.81 to 0.94‰(solution-based analyses: average of 0.87 ± 0.05‰ 1 s.d., n=5;Supplementary Table 1), but extend to higher δ56Fe values for laserablation analyses of 0.97 to 1.12‰ (average of 1.03 ± 0.05‰ 1 s.d.,n=13 grains; Supplementary Table 2).

The three core samples (WR4, 5, and 6) contain significantly dif-ferent amounts of magnetite as measured by image analysis (Table 1and Fig. 5) ranging from 4.6 vol% in section WR4 to 40 vol% in sectionWR6. The average magnetite δ56Fe values of each core sample varysystematically with the percentage of magnetite in the sample (Fig. 6),whereas the average hematite δ56Fe values for each core sample do not.

All magnetite grains examined are zoned with respect to elementalcomposition. In backscatter SEM images, these zones appear as light(electron dense) and dark (less electron dense) bands when high con-trast is used (Fig. 3D, H, and N). Electron probe analyses revealedsystematic differences in elemental compositions for the dark and lightbands in terms of their Fe and Si contents, and to a lesser extent in termsof their Al, and Mg concentrations (Supplementary Table 1 and Fig. 5).The dark bands had relatively less Fe (average of ~1.0 wt% less asFe3O4) and greater concentrations of Si (average of ~0.8 wt% more asSiO2) than the light bands. There are, however, only very minor to nodifferences between the dark bands or light bands of the three coresections. Concentrations of Cr, Ti, and Mn are all below detection limits.

4. Discussion

The Fe isotope, petrographic, and mineralogical data from the2.75 Ga Weld Range BIF presented here reveal a formation pathwaythat is distinct from other reported Archean IFs, both Algoma- andSuperior-type. Although magnetite in many IFs is inferred to haveformed after hematite (e.g., Bekker et al., 2010), this is generallythought to have occurred during early sediment compaction and burialdiagenesis, before lithification. However, petrographic analyses of thedrill core indicate that the Weld Range magnetite layering is a sec-ondary feature formed after lithification of the primary jaspilitic chertand not via direct precipitation from seawater, as evidenced by theinflated magnetite bedding (Fig. 2D) and the ubiquitous secondarymagnetite replacement of well-bedded jaspilitic layers (Fig. 2F and G).

The model for the formation of the Weld Range BIF proposed hereinvolves multiple stages including production of an Fe(III)-Si co-pre-cipitate during a single stage of oxidation of aqueous Fe(II) to form theprecursor to the hematite-bearing jasper, followed by (minor)

Fig. 3. Spatial relations of the Fe mineral phases in the Weld Range BIF samples from drill core WRRD1128. A) Photograph of a thin section of core sample WR4showing the fine-scale alternating hematite-poor (light colored) and relatively hematite-rich (darker colored) jaspilite layering. The white box indicates the areaimaged in panel B. B) Transmitted light photomicrograph of jaspilite layers. White boxes indicate the areas pictured in panel C and D. C) Higher magnificationtransmitted light photomicrograph of the jaspilite layer in B. The inset shows a backscatter electron (BSE) image of a portion of the jaspilite layer illustrating the veryfine-grained hematite (Hem) and surrounding quartz (Qtz) of which it is composed. D) BSE image of a large magnetite grain with relatively silica-rich and silica-poorzones. E) Photograph of a thin section of core sample WR6B showing the fine-scale alternating hematite-poor and relatively hematite-rich jaspilite layering (light anddark pink layers, respectively), thick and thin magnetite layers (black layers), and occasional opaque euhedral to subhedral pyrite grains (e.g., boxed areas labeled I,J, and K). The white boxes indicate the areas pictured in parts F through K. F) Reflected light photomicrograph of a thin magnetite layer (Mag) surrounded by jaspilite(Jasp) and with a pyrite grain (Pyr). G) Reflected light photomicrograph of a massive magnetite layer (Mag) with minor jasper throughout (dark gray to black). H)BSE image of a magnetite layer (black-boxed area in panel G) with relatively silica-rich and silica-poor zones. I–K) Reflected light photomicrographs of a dis-continuous magnetite layers (medium gray grains) surrounded by jaspilite (dark gray to black) with large pyrite grains (yellow). L) Photograph of a thin section ofcore sample WR6A in transmitted light showing the fine-scale alternating hematite-poor and relatively hematite-rich jaspilite layering (light and dark pink layers,respectively) as well as relatively thin magnetite layers (black layers). M) Reflected light photomicrograph of a thin magnetite layer (Mag) surrounded by jaspilite(Jasp) (white-boxed area in panel L). N) BSE image of a magnetite layer (white-boxed area in panel M) with relatively silica-rich and silica-poor zones.

Fig. 4. Iron isotope compositions of each Fe phase from three core samplescollected from different depths (WR4: 227.57–227.87 m; WR6:231.65–231.95 m; WR5: 236.40–236.70m). Pyrite data include both those ac-quired by desolvating nebulizer and laser ablation analyses. The error bars are 1s.d. for replicate analyses and 2 standard errors for single analyses (seeSupplementary Tables 1 and 2).

Table 1Magnetite contents of samples.

Core sample and depth Magnetite (% by area)

WR4 (227.57–227.87 m) 4.6WR6 (231.65–231.95 m)

WR6A 37.6WR6B 35.1WR6C 47.3Avg. 40.0

WR5 (236.40–236.70 m)WR5A 17.9WR5B 15.1Avg. 16.5

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compaction and lithification, and, in later stages, the production ofmagnetite and pyrite by fluid mobility through the rocks. Provided thatmetamorphic/hydrothermal effects may be accounted for, the Fe iso-tope compositions of the hematite may still provide an assessment ofthe extent of oxidation of the initial Fe(III) precipitates. The clear cross-cutting nature of magnetite and replacement textures of pyrite, how-ever, suggests they formed at significantly higher temperatures thanwould be associated with burial diagenesis and arose from mobilefluids, probably during metamorphism, resulting in concentration of

iron from the diffuse (low-grade) hematite-bearing protolith. Below weexplore the possibility that both magnetite and pyrite formed throughlater addition of two pulses of reduced fluids, the latter fluid containingsulfide that produced the pyrite.

The formation of magnetite and pyrite via reduction of hematite byorganic carbon is not favored in this model. If organic carbon had beenthe reductant, a significant amount of Fe carbonate is expected to havebeen produced (Posth et al., 2013), but none was detected in any of thesamples studied, nor in any of the mapped outcrops across the WeldRange. The relative δ56Fe values of hematite, magnetite, and pyritemeasured in the Weld Range BIF generally follow that expected for Feisotope equilibrium (Polyakov and Mineev, 2000; Polyakov et al., 2007;Blanchard et al., 2009; Polyakov and Soultanov, 2011; Frierdich et al.,2014), and the implications of this for mineral paragenesis is exploredbelow.

4.1. Formation of primary Fe(III) precipitates

The jaspilite layers formed during Stage 1 of the Weld Range BIFformation model by slightly varying degrees of incomplete Fe(II)aqoxidation in the water column followed by Fe(III)–Si precipitation.Evidence for this process comes from the low-Fe contents and thenarrow range of hematite δ56Fe values (Fig. 7A), suggesting the oxidantwas limited relative to the source of Fe(II)aq. The source of oxidizingpower is not constrained by the data presented here, but the hydro-thermally-sourced Fe(II)aq in the water column could have been oxi-dized by either free oxygen (presumably produced by oxygenic photo-synthesis), or by anoxygenic Fe-oxidizing microorganisms (e.g., Bekkeret al., 2010; Czaja et al., 2012; Czaja et al., 2013). This is supported byemerging evidence for free oxygen in Meso- to Neoarchean shallowoceans (e.g., Satkoski et al., 2015; Ossa Ossa et al., 2016; Eickmannet al., 2018). Based on kinetics, UV photooxidation of Fe in Archeanseawater is not a likely source of oxidation (Konhauser et al., 2007),and based on rates of deposition, oxidation of Fe by photolytically-produced H2O2 is also not a likely source of significant Fe oxidation(Pecoits et al., 2015).

Oxidation in the water column by O2 or by anoxygenic Fe oxidizersis further considered though a comparison with data and modelingresults for the 3.8 Ga BIF of the Isua Supracrustal Belt of southern WestGreenland, the 3.5 Ga jaspilitic chert Marble Bar Chert member of thePilbara region of Western Australia, and the 3.2 Ga Manzimnyama BIFof South Africa (Fig. 7B; Czaja et al., 2013; Li et al., 2013a; Satkoski

Fig. 5. Average differences in hematite and magnetiteδ56Fe values plotted against the average magnetite con-tent of each core sample. Magnetite contents are pre-sented as percent magnetite by area. X-axis error bars are1 s.d. for analyses of multiple core pieces (see Table 1). Y-axis error bars represent 1 s.d. for multiple isotopicanalyses of each Fe mineral phase. There are no errorbars for the magnetite from sample WR4 because therewas insufficient magnetite for more than one isotope ormass analyses.

Fig. 6. Average magnetite compositions for each core section. A–B) Averageconcentrations of Fe and Si (as Fe3O4 and SiO2, respectively) in magnetite asdetermined by electron microprobe (see Supplementary Table 1 for all data).Both electron dense (light) and relatively less electron dense (dark) regions (asseen in high-contrast BSE images; see Fig. 3D, H, and N) were analyzed. Errorbars are 1 s.d. for multiple analyses of each region. C) Average δ56Fe values formagnetites from each core section. See Supplementary Table 1 for data. D)Average total magnetite in each core section as determined by image analysis(see Table 1 for data).

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et al., 2015). The Weld Range BIF is assumed to have been depositedunder conditions similar to those of these older units. These models areone-dimensional dispersion-reaction models that assume open condi-tions, continuous input of Fe(II), and continuous oxidation of Fe withinor near the 100-m-deep photic zone and were run until a steady statewas reached. Two different Fe isotope fractionation factors, 2.5 and4.0‰, were considered as end members to cover all experimentallyderived values. These models assume an atmospheric oxygen con-centration of< 10−5 present atmospheric level (PAL), although theresults are fairly insensitive to variable but low PO2. Full model detailsare reported by Czaja et al. (2012, 2013) and Li et al. (2013a). Therange of δ56Fe values for the Weld Range hematite compared to theresults of previous modeling suggest that if the oxidant source was freeO2, then the water column above the Weld Range BIF depositional basinwould have contained between ~10−2 and 10−3 μM of O2 (Fig. 7B).For comparison, the modern equatorial photic zone value is ~250 μM,although oxygen minimum zone waters are typically< 1 μM (e.g.,Ulloa et al., 2012).

The average δ56Fe value of the precursor Fe-oxide precipitate for theFe-Si precipitate that produced the Weld Rage BIF was 0.61‰. Thisvalue stands in contrast to that of the Isua BIF, which was 1.2‰ (Czajaet al., 2013) and that of the Marble Bar Chert, which was 2.1‰ (Liet al., 2013a) (Fig. 7A), but is similar to that of the 3.2 Ga Man-zimnyama BIF, which was 0.55‰ (Satkoski et al., 2015). The O2 con-centrations calculated from the Weld Range Fe isotope compositions areone to two orders of magnitude greater than was estimated for thewaters from which the Isua BIF and the jaspilitic cherts of the MarbleBar Chert Member were deposited (Czaja et al., 2013; Li et al., 2013a),but overlap that estimated for the 3.2 Ga Manzimnyama BIF (Satkoskiet al., 2015) (Fig. 7B). This relation would suggest that if oxygenicphotosynthesis was the ultimate source of oxidizing power, then similaramounts of oxygen might have existed in the waters from which these3.2 and 2.75 Ga BIFs were deposited. It is important to reiterate thatthis model does not require the presence of significant free O2 in theatmosphere and this value was set to 10−5 PAL, the concentrationbelow which sulfur mass independent fractionation (S-MIF) is favored(Pavlov and Kasting, 2002). Notably the S-MIF signal, which is used as aproxy for atmospheric oxygen, contracts in the Mesoarchean suggestingan increase in atmospheric O2, before increasing again in theNeoarchean at around 2.75 Ga (e.g., Johnston, 2011). Thus, S-MIFsignatures as well as Fe isotope data suggest the oxygen concentrationof the atmosphere could have been similar during the deposition of the3.2 Ga Manzimnyama and 2.75 Ga Weld Range BIFs (see Busigny et al.(2017) for an alternate view).

The other option for oxidation considered here, bacterial anoxy-genic photosynthetic Fe oxidation, can also be tested through modeling.Fig. 7B shows these results and indicates that the rate of Fe oxidation byanoxygenic phototrophs in the Weld Range BIF depositional basin couldhave been ~0.006 μM/day. Similar to the oxic model results, this valueis greater than those estimated for the Isua BIF and MBC depositionalbasins, but only by about 20%. This rate is lower than that measured inmodern laboratories but, as discussed by Czaja et al. (2013), the ratedetermined by the modeling is an average rate, taking into accountdiurnal and annual cycles of light and nutrient availability, making itpotentially much lower than the maximum measured rate. Details ofthese oxygenic and anoxygenic models including input parameters andsensitivities are discussed by Czaja et al. (2012, 2013) and Li et al.(2013a).

The fact that the hematite, which most likely resulted from dewa-tering and diagenesis of an initial Fe(III)-Si co-precipitate, retained apositive δ56Fe value indicates that the pore fluids in the system weresufficiently enriched in Fe(II)aq relative to the amount of Fe in the se-diment (L. Wu et al., 2012). Metamorphic re-equilibration with otherminerals such as magnetite would tend to decrease the δ56Fe value ofhematite, given the positive hematite-magnetite 56Fe/54Fe fractionationfactors that have been reported (Polyakov et al., 2007; Blanchard et al.,

A Jaspilite deposition model

Avg. O2 content of photic zone (µM)10-4 10-2 100

56Fe

Fe-S

i CP (

‰)

0

1

2

3

Avg. rate of Fe(II)aq oxidation (µM.day-1)0.004 0.006 0.008

Isua BIF (Czaja et al. 2013) Weld Range

(This study)

B Models of Fe oxidationOXIDATION BY O2

ANOXYGENIC PHOTOSYNTHETIC FE OXIDATION

0.002

Marble Bar Chert(W. Li et al. 2013a)

Manzimnyama BIF(Satkoski et al. 2015)

Fe(II)aq56Fe=0‰

Fe-Si CP56Fe = 0.61 ± 0.05‰

Partialoxidation

10-3 10-1

56Fe

Fe-S

i CP (

‰)

0

1

2

3

=1.0025=1.004

Weld Range (This study)

=1.0025=1.004

Isua BIF (Czaja et al. 2013)

Marble Bar Chert(W. Li et al. 2013a)

Manzimnyama BIF(Satkoski et al. 2015)

Jaspilitic chert

Fe-poor chert

Weld Range BIF formation model: Stage 1

Fig. 7. A conceptual model of the first stage of formation of the Weld Range BIF(jaspilite deposition) and models of Fe oxidation to explain the measured Feisotope values. A) Partial oxidation of volcanic hydrothermal Fe(II)aq and de-position of the resulting Fe(III)-Si co-precipitate (Fe-Si CP) along with Fe-poorsilica producing alternating Fe-poor (white) and relatively Fe-rich (pink) layers.After deposition, the Fe(III)-Si CP was converted to hematite and micro-crystalline silica by dewatering and diagenesis, which produced alternatinglayers of jaspilitic chert and Fe-poor chert. B) Fe(III) was likely produced eitherby anoxygenic photosynthetic Fe oxidizing bacteria, or by reaction with freeoxygen produced by oxygenic photosynthesis. Models of each process and theresulting O2 content or the rate of anoxygenic Fe(II)aq oxidation needed toproduce the measured δ56FeHem values are indicated. These models are adaptedfrom those of Czaja et al. (2013) (see Section 4.1 for details). For comparison,the shaded boxes indicate the ranges of δ56Fe values for the Fe oxide precursorof the hematite in this study (dark gray) as well as those of the 3.77 Ga Isua BIF(green; Czaja et al., 2013), the 3.5 Ga Marble Bar Chert (orange; Li et al.,2013a), and the 3.2 Ga Manzimnyama BIF (blue; Satkoski et al., 2015). The thindownward pointing arrows in the upper panel of part B indicate the maximummodeled O2 content of the photic zone based on the average δ56Fe values.

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2009; Frierdich et al., 2014). However, there is no textural evidence forequilibration between these two phases at the mm-scale of sampling, sothe preservation of primary bedding textures in the hematite suggestpreservation of primary Fe isotope compositions as well.

4.2. Origin of magnetite

Magnetite formed during Stage 2 of the Weld Range BIF formationmodel (Fig. 8A) and has slightly lower average δ56Fe values than thoseof the “primary” hematite. Such compositions, as well as petrographicevidence that the magnetite layers are secondary to the jaspilite, reflectnet addition of reducing fluids that remobilized hematite Fe alongfractures that were developed largely parallel to the bedding planes

(Fig. 8A), but that locally crosscut the bedding (e.g., Fig. 2D). Althoughthe magnetite δ56Fe values fall over a narrow range, they are sig-nificantly different between the three portions of the drill core studied,with differences directly proportional to the magnetite content of eachsample (Fig. 5). This relation suggests that the samples with low δ56Fevalues for magnetite and low magnetite contents (e.g., WR4) formed bya small amount of reduction from a solution that contained a sufficientmass of Fe to allow isotope equilibration between the hematite andmagnetite in the layers where reduction occurred, as indicated by theirhigh temperature equilibrium fractionation of Δ56FeMag–Hem=0.29‰(Fig. 8A). Again, it is important to stress that the hematite sampled forthis study had primary bedding textures suggesting they also preservedprimary Fe-isotope compositions. Conversely, the samples with high

Fig. 8. Secondary and tertiary stages of Weld Range BIF formation (cf. Fig. 7). A) Stage 2: after conversion of Fe(III)-Si CP to jaspilitic chert, the jaspilite wasinfiltrated along bedding planes by reducing fluids, converting hematite to magnetite. In some portions of the unit (e.g., WR4), only a small amount of reductionoccurred and there was sufficient aqueous Fe to allow exchange and Fe isotope equilibration between hematite and magnetite (Stage 2a). Equilibrium modeling ofΔ56FeMag–Hem (right) based on published beta values (model 1: based on the magnetite-Fe(II)aq fractionation factor of Frierdich et al. (2014) and the hematite-Fe(II)aqfractionation factor of Blanchard et al. (2009); model 2: based on the magnetite-Fe(II)aq fractionation factor of Frierdich et al. (2014) and the hematite-Fe(II)aqfractionation factor of Polyakov et al. (2007)). In other portions of the unit (e.g., WR6), a large amount of reduction occurred with insufficient aqueous Fe to allowexchange so the precipitated magnetite retained the same Fe isotope composition as the jaspilitic hematite (Stage 2b). B) Stage 3: sulfide-rich fluids infiltrated certainsections of the unit (e.g., WR6) and reacted to form pyrite. Assuming equilibrium conditions, an average pyrite δ56Fe value of 1.03‰ indicates a δ56FePyr–Hem value of0.44‰ and formation temperatures of> 200 °C depending on the beta factors used (model 1 and 2: based on the pyrite-Fe(II)aq fractionation factor of Blanchardet al. (2009) and the hematite-Fe(II)aq fractionation factor of either Polyakov et al. (2007) (model 1) or Blanchard et al. (2009) (model 2); models 3 and 4: based onthe pyrite-Fe(II)aq fractionation factor of Polyakov and Soultanov (2011) and the hematite-Fe(II)aq fractionation factor of either Polyakov et al. (2007) (model 3) orBlanchard et al. (2009) (model 4).

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δ56Fe values for magnetite and high magnetite contents (e.g., WR6) areinterpreted to have formed by a large amount of reduction, where thetotal mass of Fe in the solution was low and hence did not provide alarge reservoir that could promote isotopic exchange (Fig. 8A). To thedegree that this approximated a closed system with respect to Fe, onewould expect minimal re-equilibration between magnetite and hema-tite, where the δ56Fe value of magnetite would largely be inheritedfrom hematite, given the fact that 2/3 of Fe in magnetite is Fe(III) andhence could be incorporated directly into magnetite.

The difference in Fe content between the jaspilite layers (≤7.2 wt%) and magnetite layers (≥54.5 wt%) bears on magnetite paragenesis.If reduction of hematite occurred in situ, a large volume decrease ofportions of the unit must have occurred during Stage 2 of the system(Fig. 8A). Brecciated portions of the unit (e.g., Fig. 2G) appear tosupport this interpretation. As illustrated in Fig. 8, the jaspilitic chert isenvisioned to have had a substantially larger initial volume than themagnetite-bearing units. It is also possible that a reducing fluid addedFe(II), as illustrated in Fig. 8, although Fe isotope compositions suggestthat the δ56Fe value of such a fluid would have needed to have hadpositive δ56Fe values like that of the jaspilite oxides. However, thishypothesis is not favored because it is difficult to produce a Fe(II)-richfluid that has a positive δ56Fe value given the known Fe isotope com-positions of hydrothermal fluids (δ56Fe≤ 0‰) and Fe isotope fractio-nation factors that show ferrous Fe to be isotopically light (Heimannet al., 2008). The possibility that the magnetite layers represent primarysedimentary layers that were later reworked is in conflict with thepetrographic evidence that they formed by the inflation of the jaspilitelayers.

Modeled equilibrium Δ56FeMag–Hem fractionation factors over arange of temperatures (models 1 and 2, Fig. 8A) can be used to estimatethe metamorphic temperature under which the magnetite formed. Thesample with the maximum Δ56FeMag–Hem value (−0.29‰) and model 1indicates a formation temperature of ~350 °C, whereas model 2 in-dicates a temperature of ~550 °C. We prefer equilibrium model 1,based on the close agreement with the previous maximum metamorphictemperature estimate of 320 °C (Gole, 1980).

The chemical compositions of magnetite in the Weld Range BIFprovide additional constraints on its paragenesis, most notably Si con-centration. Weld Range magnetite is similar to that of magnetite inportions of the c. 2.45 Ga Dales Gorge Member BIF of the BrockmanIron Formation, Hamersley Basin, Western Australia (Li et al., 2013b) interms of BSE contrast (Fig. 3) and silica content. Magnetite in theseyounger, Hamersley BIFs was divided into “silician magnetite” and “Si-poor magnetite”, based on the Si concentration. Silician magnetite wasdefined by Li et al. (2013b) as magnetite having an SiO2 content be-tween 1 and 3wt%, and Si-poor magnetite as< 1wt%. In the DalesGorge Member, such magnetite was demonstrably hydrothermal inorigin based on textures and O isotope compositions (Huberty et al.,2012; Li et al., 2013b). The average SiO2 content of the silician mag-netite of the Weld Range magnetite is 1.17 ± 0.23wt%, but two ana-lyzed areas had contents below 1wt% (Supplementary Table 1). Thesilica-poor zones have average SiO2 contents of 0.35 ± 0.18 wt%.However, the silica-rich and silica-poor zones do not differ significantlyin Fe or Si content between the core sections (Supplementary Table 1and Fig. 6). These data provide powerful constraints that the magnetitein the Weld Range is metamorphic or hydrothermal in origin and notrelated to early microbial diagenesis prior to lithification.

4.3. Origin of pyrite

The coarsely crystalline pyrite of metamorphic/hydrothermal originin the Weld Range BIF (e.g., Fig. 3F and I–K) formed during Stage 3(Fig. 8B). The pyrite δ56Fe values measured here are some of the highestyet reported from sedimentary rocks and probably reflect interactionbetween dissolved sulfide and excess Fe(II)aq flowing through thesystem after the magnetite formed (Fig. 8B). Equilibrium modeling

suggests the pyrites formed at a temperature of> 200 °C, but variousreported beta factors suggest temperatures over a range of> 300 °C(Fig. 8B) making an accurate temperature estimate difficult. The δ56Fevalues of the pyrite are all positive and greater than those of both thehematite and magnetite (Supplementary Table 1 and Fig. 4). Thisstands in dramatic contrast to pyrite in Archean black shales, whichextend to some of the most negative δ56Fe values measured in nature(e.g., Rouxel et al., 2005).

Pyrites sampled by hand mill have significantly lower δ56Fe valuesthan those measured by laser ablation (an average of 0.16‰ lower;compare Supplementary Tables 1 and 2). This discrepancy is likely aresult of the coarseness of the hand scribe sampling technique, whereinsome Fe-oxides were sampled in addition to pyrite. In some cases, theadditional material would have been jaspilite, but the concentration ofFe in the jasper layers (maximum of ~7wt%; Supplementary Table 1) islow enough to have contributed an insignificant amount of Fe to thesample. In other cases, the grains contained, or were surrounded by,significant magnetite (e.g., Fig. 3I and J) that resulted in apparentpyrite δ56Fe values that were lower than expected. Thus, the laser ab-lation δ56Fe values are more reliable and represent those of pure pyrite.

4.4. Comparison to other Archean Algoma-type IFs

The overall formation model of the Weld Range BIF is distinct fromother Algoma-type IFs from the Archean, most notably the 3.8 Ga IsuaBIF from southern West Greenland (Czaja et al., 2013) and the ~2.7 GaOld Wanderer IF from Zimbabwe (Steinhoefel et al., 2009b). The Isua IFis composed largely of magnetite and chert, with minor pyrite andcarbonate. Based on Fe isotope compositions, petrology, and modeling,the magnetite in this unit was interpreted to have formed via additionof Fe(II)-rich fluids to the original ferric oxyhydroxide precipitateduring early diagenesis to produce the mixed valence state Fe3O4 (Czajaet al., 2013). The Old Wanderer IF is largely composed of magnetite,chert, siderite, and ankerite. Based on Fe and Si isotopes and petrology,these Fe minerals are interpreted to have formed via varying degrees ofdiagenetic reduction of a ferric oxyhydroxide protolith during earlymetamorphism by co-occurring organic matter (Steinhoefel et al.,2009b).

The fact that all Fe-bearing phases in the Weld Range BIF havepositive δ56Fe values relative to the likely near-zero or slightly negativeδ56Fe values of hydrothermal Fe(II)aq sources (Yamaguchi et al., 2005;Johnson et al., 2008b) indicates that the Fe inventory was likely smallrelative to the hydrothermal flux, based on simple isotopic mass bal-ance. The isotope compositions of the jaspilite is interpreted to havebeen inherited from the Fe(III)-Si precursor, which would be predictedto have had a δ56Fe value of 2.3–4.0‰ (Wu et al., 2011; Wu et al.,2012; Wu et al., 2017), assuming equilibrium fractionation duringprecipitation and, importantly, a small extent of Fe(II) oxidation. Thelower δ56Fe value of Weld Range hematite suggests partial oxidation ofthe hydrothermal plume resulting in lower than maximum fractiona-tion.

4.5. Contrasts with Superior-type BIFs and secular trends in iron formationdeposition

The model proposed here stands in contrast to theNeoarchean–Paleoproterozoic Superior-type IFs, which are interpretedto have been formed via a significant degree of biological Fe cyclingwithin the sediment; where biological cycling has been inferred basedon Fe isotope compositions of IFs, one of the key lines of evidence hasbeen Fe isotope disequilibrium among the phases (Johnson et al.,2008b; Heimann et al., 2010; Li et al., 2015). For alternate views, seerecent work by McCoy et al. (2017) and Rasmussen and Muhling(2018). Most notably, dissimilatory iron reduction has been invoked toexplain the large amount of reduced Fe phases (e.g., magnetite) thathave negative δ56Fe values (e.g., Johnson et al., 2008b). The average

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δ56Fe value of −0.2‰ for Fe oxides from the 2.45 Ga Transvaal andHamersley Superior-type BIFs differs significantly from the positiveδ56Fe values of the earlier Algoma-type BIFs (Fig. 9). The measuredδ56Fe values of the Weld Range BIFs are similar to the average mag-netite δ56Fe value of ~0.7‰, but lower than the estimated averageδ56Fe values of 1.2‰ for the initial Fe(III)-Si precipitate that formed themagnetite in the 3.8 Ga Isua Supracrustal Belt BIF (Czaja et al., 2013)and the average measured δ56Fe values of ~2.0‰ for the hematite ofthe 3.46 Ga Marble Bar Chert (Li et al., 2013b). Thus, when consideringaverage δ56Fe values of sedimentary Fe oxides for various time periodsin the Archean, there is an apparent overall Fe isotope trend from highpositive δ56Fe values in the Paleoarchean to lower positive δ56Fe valuesof the early Neoarchean to average δ56Fe values of ~0‰ for IFs in theNeoarchean (Fig. 9). Note that averages are used here to remove localspatial and short-term temporal heterogeneity. There is perhaps a de-viation from this trend in the late Mesoproterozoic, but the number ofsamples analyzed at this time is low (n=13; Fig. 9) so it is difficult toknow if this average is truly representative of the average Fe isotopecomposition at that time. The overall trend, however, suggests a generalincrease in oxidative power on Earth's surface with decreasing age overthe course of the Archean, which is consistent with an overall change inapparent abundance of dominantly abiologically-cycled Algoma-typeIFs in the Paleoarchean to early Neoarchean to biologically-cycled Su-perior-type IFs in the Neoarchean–Paleoproterozoic, and independentevidence for increasing oxidation of the surface throughout the Archean(e.g., Anbar et al., 2007; Wille et al., 2007; Kendall et al., 2010;Voegelin et al., 2010; Czaja et al., 2012; Czaja et al., 2013; Satkoskiet al., 2015). This increased oxidative power resulted in increased Fedeposition and an expansion of dissimilatory iron reduction in parti-cular, an increase in significance of a number of other microbial me-tabolisms (e.g., Johnson et al., 2008a), an increase in productivityoverall, and ultimately the irreversible oxygenation of Earth's atmo-sphere during the Great Oxidation Event.

5. Conclusions

A hematite–magnetite–(± )pyrite Algoma-type iron formation ofthe 2.75 Ga Wilgie Mia Formation in the Weld Range of the YilgarnCraton, Australia, was studied by petrographic and iron isotopic ana-lyses. Petrographic analyses reveal that fine-grained hematite in jaspi-lite was the initial Fe phase and formed from deposition in a quiescentbasin, followed later by the intrusion of Fe-rich fluids to split jaspilitealong sedimentary surfaces and precipitate magnetite, and later pyrite.

These data combined with Fe isotope analyses reveal that the jaspilitewas deposited either through direct biological oxidation of Fe(II)aq inthe ancient Weld Range basin or through indirect biological processesvia biologically-produced O2. Also, extensive hydrothermal and meta-morphic recrystallization and mineral formation under greenschist-fa-cies conditions followed to produce the magnetite and pyrite phases.The results and interpretations reported here suggest that despitebroadly similar Fe isotope compositions (positive δ56Fe values spreadover narrow ranges), and a similar initial mode of formation by pre-cipitation of an Fe(III)-Si co-precipitate from a hydrothermal source ofFe, the Weld Range BIF is distinct from other Archean volcanogenic IFsin the timing and mode of magnetite formation.

The Weld Range BIF data fit an overall secular trend in Fe-oxideδ56Fe values over the Archean that reflect changes in the oxidation stateof the atmosphere-hydrosphere system and thus changes in IF forma-tion pathways. The IFs of the Paleoarchean through the earlyNeoarchean (including the time when the Weld Range BIF was de-posited) are almost entirely of the Algoma type, and the δ56Fe values ofFe-oxides grade from highly positive to moderately positive with a re-latively small spread over this time, suggesting limited but increasingoxidants. However, the large Superior-type IFs of the Neoarchean andPaleoproterozoic included significant oxidation of Fe(II)aq resulting inan average fractionation of approximately δ56Fe=0‰, but also amajor component of biological Fe cycling that spread the range overalmost the entire extent of natural Fe isotope fractionation.

This study also demonstrates the complexity of Fe cycling in theArchean, and natural settings in general, making fine-scale sampling ofindividual mineral phases and laser-ablation analyses of individualmineral grains critical to interpret formation pathways.

Acknowledgements

This research was funded by the NASA Astrobiology Institute grantNNA13AA94A (CMJ, BLB, ADC). The Weld Range iron formationsamples were provided by Atlas Iron. The Geological Survey of WesternAustralia provided support to MVK during his regional mapping of theWeld Range area. The authors also thank two anonymous reviewers forcomments and suggestions that improved the manuscript.

Appendix A. Supplementary data

Supplementary data to this article can be found online at https://doi.org/10.1016/j.chemgeo.2018.04.019.

Fig. 9. Box plots of δ56Fe data for iron oxides (magnetite,hematite, or mixed) from Archean iron formations andother iron rich deposits from the Archean and earlyPaleoproterozoic. Data include δ56Fe values for ironoxides from this study (gray box plot centered at2750Ma) and from the literature (white box plots),measured by either conventional analyses (Johnson et al.,2003; Dauphas et al., 2004; Rouxel et al., 2005; Dauphaset al., 2007a, 2007b; Yamaguchi et al., 2007; Johnsonet al., 2008b; Johnson et al., 2008b; Steinhoefel et al.,2009a; Czaja et al., 2010; Tsikos et al., 2010; Planavskyet al., 2012; Czaja et al., 2013; Li et al., 2013a; Li et al.,2015; Satkoski et al., 2015; Busigny et al., 2017; Smithet al., 2017; Teixeira et al., 2017; Lantink et al., 2018) orlaser ablation (Steinhoefel et al., 2009b; Steinhoefel et al.,2010; Czaja et al., 2013; Li et al., 2013; Li et al., 2015).Each box represents the 25th to the 75th percentile of thedata, the horizontal line within each box is the median,the whiskers indicate the 10th and 90th percentiles, andthe symbols above and below the whiskers (• – WeldRange BIF.; ✕ – all other data) indicate the outliers foreach population. The number of measurements includedin each 250Ma bin is listed on each box.

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