Earth and Planetary Science Letters 303 (2011) 121–132

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Earth and Planetary Science Letters

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Iron and carbon isotope evidence for microbial iron respiration throughoutthe Archean

Paul R. Craddock ⁎, Nicolas DauphasOrigins Laboratory, Department of the Geophysical Sciences and Enrico Fermi Institute, The University of Chicago, 5734 South Ellis Avenue, Chicago, IL 60637, United States

⁎ Corresponding author. Tel.: +1 773 834 3997.E-mail address: craddock@uchicago.edu (P.R. Craddo

0012-821X/$ – see front matter © 2010 Elsevier B.V. Adoi:10.1016/j.epsl.2010.12.045

a b s t r a c t

a r t i c l e i n f o

Article history:Received 17 August 2010Received in revised form 20 December 2010Accepted 22 December 2010Available online 22 January 2011

Editor: R.W. Carlson

Keywords:iron-formationHamersleyIsuairon carbonatesiron respiration

Banded Iron-Formations (BIFs) are voluminous chemical sediments that are rich in iron-oxide, carbonate andsilica and whose occurrence is unique to the Precambrian. Their preservation in the geological record offersinsights to the surface chemical and biological cycling of iron and carbon on early Earth. However, many detailsregarding the role of microbial activity in BIF deposition and diagenesis are unresolved. Laboratory studies haveshown that reactionbetween carbon and iron throughmicrobial iron respiration [2Fe2O3∙nH2O+CH2O+7H+→4Fe2++HCO3

−+(2n+4)H2O+chemical energy] can impart fractionation to the isotopic compositions of theseelements. Here, we report iron (δ56Fe, vs. IRMM-014) and carbon isotopic (δ13C, vs. V-PDB) compositions ofmagnetite and of iron-rich and iron-poor carbonates in BIFs from the late Archean (~2.5 Ga) Hamersley Basin,Australia and the early Archean (~3.8 Ga) Isua Supracrustal Belt (ISB), Greenland. The range of δ56Fe valuesmeasured in the Hamersley Basin, including light values in magnetite and heavy values in iron-rich carbonates(up to+1.2‰), are incompatible with their precipitation in equilibriumwith seawater. Rather, the data togetherwith previously reported light δ13C values in iron-rich carbonates record evidence for diagenetic reduction offerric oxide precursors to magnetite and carbonate through microbial iron respiration (i.e., dissimilatory ironreduction, DIR). Iron and carbon isotope data of iron-richmetacarbonates from the ISB are similar to those of lateArcheanBIFs. The isotopic signatures of thesemetacarbonates are supportive of an early diagenetic origin despitemetasomatic overprint, and preserve evidence of microbial iron respiration within the oldest recognizedsedimentary rocks on Earth.

ck).

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© 2010 Elsevier B.V. All rights reserved.

1. Introduction

Banded Iron-Formations (BIFs) are conspicuously laminated marinechemical sediments that are characterized byhigh concentrations of iron-bearing minerals (20–40 wt.% bulk Fe) commonly interbedded withlayers of silica, and whose occurrence is unique to the Precambrian(James, 1954, 1983). Themineralogyof thebest-preservedBIFs consists ofcombinations of four dominant facies: oxide (magnetite, hematite),carbonate (siderite, ankerite, Fe–dolomite and, less commonly, calcite),chert and silicate (stilpnomelane, riebeckite, greenalite, minnesotaite),and locally sulfide (pyrite) and phosphate (apatite). Most known BIFshave ages in the range ~3.8 to 1.8 Ga, but these formations also occur to alesser extent in theNeoproterozoic at ~700 Ma(Klein, 2005). The studyofthese formations preserved in the rock record offers critical insights tosurface geochemical cycles and chemical evolution of the ocean andatmosphere in the Precambrian, and in particular the Archean (≥2.5 Ga).

Despite significant scientific interest and research, however, there isno consensus on the origin of BIFs. The primarymechanism for oxidationof Fe(II)aq in an Archean ocean that was purportedly anoxic (Canfield

et al., 2000; Farquhar et al., 2000; Kasting, 1987; Ono et al., 2003; Pavlovand Kasting, 2002), is uncertain. Photochemical oxidation of Fe(II)in surface ocean waters owing to interaction with incident UV radiationhas been proposed as an entirely abiological means of accounting forferric iron in BIFs (Braterman et al., 1983; Cairns–Smith, 1978). Oxidationof Fe(II) by O2 produced via photosynthesis has also been suggested(Cloud, 1965, 1973), implying an indirect biological influence on BIFformation and hinting at the presence of free O2 oases in the Archeansurface ocean. Alternatively, direct biological activity has been implicated,via anoxygenic photosynthesis that coupled oxidation of Fe(II) toreduction of inorganic carbon to yield organic compounds (Garrelset al., 1973; Kappler et al., 2005; Konhauser et al., 2007; Widdel et al.,1993). It is also uncertain the extent to which the mineral assemblagespreserved in BIFs reflect either primary precipitates from seawater,possibly in near-chemical equilibriumwith the ocean and atmosphere, orare authigenic minerals formed during early sedimentary diagenesis andburialmetamorphism. For example, themineralogical, chemical (e.g., rareearth element) and isotopic (e.g., δ13C, δ18O) characteristics of iron-richcarbonates such as siderite [FeCO3] and ankerite [Ca0.5(Fe,Mg)0.5CO3] inBIFs have been used to argue either for primary precipitation from ananoxic and stratified water column (Beukes et al., 1990; Kaufman et al.,1990; Klein and Beukes, 1989) or for an authigenic origin (Becker andClayton, 1972; Heimann et al., 2010; Walker, 1984).

Table 1Iron and carbon isotopic compositions of oxides and carbonate from drill core samples of the Brockman Iron Formation, Hamersley Basin.

Samplea Hole a Macroband Mesoband Type Depthb

(m)δ13C(‰)c

δ56Fe(‰) d

δ57Fe(‰)d

Ankerite Siderite Ankerite/Siderite Magnetite Hematite Ankerite/Siderite Magnetite Hematite

Dales Gorge Member, Brockman Iron FormationDrill core from Wittenoom Gorge Area (118° 28′ E; 22° 25′ S)13324 M1 27 BIF 12 Magnetite 85.2 0.424±0.031 0.630±0.04313327 WC2 27 BIF 11 Chert 98.4 −11.41 −11.08 −0.289±0.031 0.205±0.031 −0.442±0.049 0.316±0.04913327 WC1 27 BIF 11 Chert 98.4 −15.05 0.013±0.031 0.465±0.031 0.008±0.046 0.699±0.04313321 HC 27 BIF 10 Chert+hematite 99.5 −6.98 −7.45 −0.705±0.031 0.908±0.038 −1.023±0.083 1.369±0.04413318 M3 27 BIF 7 Magnetite 122.2 −9.24±0.03 0.006±0.036 0.005±0.05713318 Pl 27 BIF 7 Chert 122.2 −9.83±0.04 −0.584±0.029 −0.865±0.05113318 FM1 27 BIF 7 Chert 122.2 −9.72±0.06 0.118±0.031 −0.210±0.034 0.373±0.029 0.198±0.046 −0.298±0.054 0.561±0.05113318 M1 27 BIF 7 Magnetite 122.3 0.159±0.029 0.230±0.05113318 M2 27 BIF 7 Magnetite 122.3 0.112±0.031 0.179±0.04613316 M1 27 BIF 6 Magnetite 128.1 0.650±0.038 0.949±0.05813316 CA1 27 BIF 6 Fine−band combination 128.2 −9.31 0.641±0.031 0.636±0.033 0.933±0.083 0.914±0.06013316 CB1 27 BIF 6 Coarse-band combination 128.2 −9.01 −9.41 0.178±0.033 0.196±0.031 0.283±0.060 0.244±0.04913309 FM1 27 BIF 1 Chert 162.2 −9.87 0.057±0.031 0.588±0.031 0.070±0.046 0.852±0.04313309 PC1 27 BIF 1 Chert 162.3 −10.06 −0.161±0.031 0.490±0.030 −0.236±0.043 0.710±0.04613309 QIO1 27 BIF 1 Chert-matrix 162.4 0.675±0.031 0.790±0.034 0.990±0.046 1.144±0.05413309 M2 27 BIF 1 Magnetite 162.4 0.488±0.030 0.694±0.04413326 M1 28 BIF 12 Magnetite 110.9 0.418±0.031 0.610±0.04613328 WC2 28 BIF 11 Chert 114.9 −11.07 −12.48 −0.555±0.031 −0.757±0.04913328 WC1 28 BIF 11 Chert 115.0 −12.86 0.426±0.038 0.630±0.04413322 WC1 28 BIF 10 Chert 125.2 −9.73 0.453±0.035 0.662±0.04713322 WCIA 28 BIF 10 Chert 125.2 −9.61 0.294±0.027 0.419±0.04013322 M1 28 BIF 10 Magnetite 125.2 0.730±0.033 1.107±0.06013322 HC 28 BIF 10 Chert+hematite 125.2 −6.76 −8.26 −0.792±0.038 0.398±0.030 −1.142±0.044 0.584±0.05713322 WHC 28 BIF 10 Chert+hematite 125.2 −6.50±0.03 −7.80 −1.081±0.033 −1.601±0.06013322 FM1 28 BIF 10 Chert 125.3 −8.96 −9.79 0.787±0.033 1.194±0.033 1.158±0.044 1.778±0.04413322 QIO1 28 BIF 10 Chert-matrix 125.3 0.566±0.038 0.998±0.038 0.847±0.058 1.477±0.05813319 FM3 28 BIF 7 Chert 146.5 −9.74±0.03 −0.318±0.029 −0.287±0.030 −0.470±0.043 −0.422±0.04413319 FM2 28 BIF 7 Chert 146.6 −10.18±0.01 0.086±0.030 0.064±0.030 0.156±0.044 0.077±0.04413319 M1 28 BIF 7 Magnetite 146.6 −0.022±0.030 0.014±0.04413319 FM1 28 BIF 7 Chert 146.6 −9.90 0.075±0.030 −0.149±0.029 0.093±0.044 −0.212±0.04313313 M1 28 BIF 2 Magnetite 174.5 1.085±0.029 1.582±0.04313310 PC1 28 BIF 1 Chert 186.7 −0.181±0.028 0.159±0.035 −0.303±0.047 0.241±0.047M 2 40 BIF 2 Magnetite 196.8 0.551±0.035 0.810±0.047P 2 40 BIF 2 Chert 196.8 −9.70±0.02P 2A 40 BIF 2 Chert 196.8 −9.80±0.06

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Table 1 (continued)

Samplea Hole a Macroband Mesoband Type Depthb

(m)δ13C(‰)c

δ56Fe(‰) d

δ57Fe(‰)d

Ankerite Siderite Ankerite/Siderite Magnetite Hematite Ankerite/Siderite Magnetite Hematite

Dales Gorge Member, Brockman Iron FormationDrill core from Wittenoom Gorge Area (118° 28′ E; 22° 25′ S)

FM 2 40 BIF 2 Chert 196.8 −10.28±0.02M 3 51 BIF 12 Magnetite 104.5 0.314±0.028 0.478±0.047QIO 51 51 BIF 5 Chert-matrix 155.4 −0.189±0.028 −0.282±0.047M 1 51 BIF 5 Magnetite 155.6 −0.092±0.035 −0.111±0.047P l 51 BIF 5 Chert 155.6 −8.65

Drill core from Yampire Gorge Area (118° 37′ E; 22° 27′ S)13325 WC2 Y3 BIF 12 Chert 66.2 −10.95±0.02 −0.305±0.041 0.202±0.029 −0.448±0.072 0.298±0.04113325 M2 Y3 BIF 12 Magnetite 66.2 0.351±0.029 0.515±0.04113325 M1 Y3 BIF 12 Magnetite 66.3 0.344±0.051 0.519±0.13313325 FM1 Y3 BIF 12 Chert 66.3 −10.96 0.301±0.033 0.297±0.029 0.410±0.083 0.433±0.04113325 WC1 Y3 BIF 12 Chert 66.4 −10.84±0.01 −0.404±0.029 0.194±0.029 −0.588±0.043 0.282±0.13313329 M1 Y3 BIF 11 Magnetite 70.7 −12.59±0.50 −0.005±0.031 −0.017±0.08313329 FM1 Y3 BIF 11 Chert 70.8 −11.11 0.244±0.033 0.408±0.033 0.674±0.051 0.395±0.083 0.632±0.083 1.057±0.13313323 M1 Y3 BIF 10 Magnetite 78.2 0.727±0.034 1.029±0.05413320 FM1 Y3 BIF 7 Chert 98.4 −9.53 0.013±0.033 −0.156±0.029 −0.029±0.083 −0.253±0.04113317 CA1 Y3 BIF 6 Fine-band combination 104.4 −9.36 −9.34 0.605±0.038 0.870±0.05813317 CBI Y3 BIF 6 Coarse-band combination 104.4 −9.06 −9.31 0.456±0.038 0.710±0.04413317 M1 Y3 BIF 6 Magnetite 104.5 0.665±0.085 0.887±0.05813314 M3 Y3 BIF 3 Magnetite 125.3 0.735±0.029 1.073±0.04113314 M2 Y3 BIF 3 Magnetite 125.3 0.877±0.028 1.283±0.04713314 H1 Y3 BIF 3 Hematite 125.3 1.055±0.033 1.568±0.08313314 M1 Y3 BIF 3 Magnetite 125.3 0.886±0.041 1.301±0.07213314 FM1 Y3 BIF 3 Chert 125.3 0.963±0.028 1.419±0.04713314 Carb 1 Y3 BIF 3 Chert+carbonate 125.5 −9.48 1.212±0.041 1.041±0.041 1.850±0.072 1.549±0.07213311 PC1 Y3 BIF 1 Chert 137.6 −9.75 −0.033±0.029 0.600±0.027 −0.053±0.041 0.868±0.04013311 PC1 Y3 BIF 1 Chert 137.6 0.096±0.035 0.135±0.047

a Sample numbers follow the nomenclature adopted by Becker (1971). Core samples were recovered by drilling operations in two areas: Wittenoom Gorge and Yampire Gorge. Hole designates core number.b Depth in meters below modern surface.c Carbon isotopic (δ13C) ratios are per mil difference relative to PDB standard and are those published by Becker and Clayton (1972). Quoted uncertainties are the standard deviation of replicate analyses as reported by Becker and Clayton

(1972).d Iron isotopic (δ56Fe, δ57Fe) ratios are reported in per mil difference relative to iron metal standard IRMM-014 (δ57Fe~1.5xδ56Fe). Uncertainties are 95% confidence intervals.

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In order for BIFs to offer a faithful and critical perspective of theEarth's early ocean and atmosphere biogeochemical cycles, an under-standing of processes leading to their formation [i.e., primary Fe(II)oxidation and precipitation versus diagenetic transformations] isessential. Here, we report coupled iron, δ56Fe=[(56Fe/54Fe)sample/(56Fe/54Fe)IRMM-014−1]×103, and carbon, δ13C=[(13C/12C)sample/(13C/12C)V-PDB−1]×103, isotope compositions of iron oxides and carbonatesfromoneof themostwell-preserved and leastmetamorphosedBIFs, the~2.5 Ga Brockman Iron Formation of the Hamersley Basin, westernAustralia, and in one of the oldest known Fe-rich chemical sedimentaryformations, that of the~3.8 Ga Isua Supracrustal Belt (ISB), southernWestGreenland. Our data are used to constrain the probable chemical andbiological pathways for the precipitation of iron-bearing minerals inArcheanBIFs. Aprimarymotivation for this research is the study of Beckerand Clayton (1972) that reported δ13C of iron-rich carbonates (siderite,ankerite) in the Brockman Iron Formation, Hamersley Basin, mostlybetween−8 and−11‰, significantlymore depleted than that of typicaliron-poor marine carbonates (calcite, dolomite) with δ13C between −2and+2‰. The source of light carbon in the iron-formationwas suggestedto derive from organic carbon, possibly resulting from oxidation–reduction reactions involving microbially metabolized Fe(III). Weexamine specifically whether the range of iron and carbon isotope ratiosrecorded in iron-rich carbonates and magnetite from the Brockman IronFormation is consistentwith such a process and thus records evidence formicrobial iron respiration (i.e., dissimilatory iron reduction; DIR) in thelate Archean, a process that was originally suggested by Walker (1984)and Baur et al. (1985). Johnson et al. (2008) have recently reported theiron isotope compositions of magnetite and siderite from the BrockmanIron Formation in order to distinguish between primary precipitation inseawater and diagenetic iron transformations. These authors suggestedthat the rangeofmeasured iron isotopic compositions couldbe consistentwith an authigenic origin, but they did not have access to carbon isotopicdata for the same carbonates to validate a DIR pathway. Complementaryiron and carbon isotopic studies ofmagnetite and carbonates from the 2.5to 2.4 Ga Kuruman Iron Formation, Transvaal Craton, South Africa, havealso explored evidence for microbial iron respiration during this periodof Earth's history (Heimann et al., 2010; Johnson et al., 2003). Heimannet al. (2010) concluded that the range of δ13C and δ56Fe valuesdocumented in iron-rich carbonates was consistent with authigenicpathways for carbonate formation through microbial DIR.

Further, we assess whether a record of microbial iron respirationextends to the early Archean. Dauphas et al. (2007) have reported ironisotopic compositions of iron-rich and iron-poor metacarbonates fromthe ~3.8 Ga ISB. Metacarbonates from the ISB typically show ametasomatic relationship with their host rocks (Rose et al., 1996; Rosinget al., 1996). Rose et al. (1996) and Rosing et al. (1996) have argued thatmetacarbonates were formed by metasomatic alteration of igneousprotoliths. However, the iron isotopic and trace element (e.g., REE, Fe/Ti)characteristics of iron-richmetacarbonates aremore similar to that of co-existingBIFs in the ISBandaremore supportiveof a chemical sedimentaryorigin (Bolhar et al., 2004; Dauphas et al., 2007). Here, we integrate newcarbon and existing iron isotopic data of iron-rich and iron-poormetacarbonates from the ISB to re-examine the origin of these rocks, inparticular evaluating whether these isotopic signatures are supportive ofa sedimentary origin prior tometasomatic overprint and record evidenceof microbial iron respiration within the oldest recognized sedimentaryrocks on Earth.

2. Materials and methods

2.1. Geology and sample descriptions

The geological setting, description and analytical procedures forisotopic measurement of iron-formations in this study are describedin detail in the Supporting Online Material, and summarized below.The Hamersley Basin, western Australia is a roughly elliptical basin in

which late Archean and Paleoproterozoic (~2.6 to 2.45 Ga) volcanicsand sediments of the Hamersley Group were deposited (Trendall andBlockley, 1970). The Hamersley Group is of particular significanceowing to the presence of four major and several minor conspicuouslylaminated, iron-rich stratigraphic units (Banded Iron Formation, BIF),including the Brockman Iron Formation (MacLeod, 1966; Trendall,2002; Trendall and Blockley, 1970). The Hamersley Group alsocontains the Wittenoom Dolomite, which comprises between 200and 300 m of interbedded carbonate, chert and shale with thecarbonate occurring typically as mostly massive calcite and dolomite.Samples from the Hamersley Basin were selected both from an iron-formation (Dales GorgeMember of the Brockman Iron Formation) andfrom a carbonate formation (Wittenoom Dolomite). Samples of iron-formation included iron oxide (magnetite) and iron-rich carbonates(siderite, ankerite) and were from fresh drill core material that wasrecovered by the Australian Blue Asbestos Co. as part of a prospectingprogram under subsidy from the Western Australian Government(Trendall and Blockley, 1970). Samples of carbonate formation werehand specimens of massive, iron-poor carbonate (calcite, dolomite)obtained from chipping out fresh surfaces of theWittenoom Dolomitefrom weathered exposures. All samples were originally collected andprepared for carbon and oxygen isotope studies by Becker and Clayton(1972,1976).

The Isua Supracrustal Belt (ISB) is part of the ~3.8 Ga high-gradeItsaq Gneiss Complex and is one of the oldest metasedimentary-bearing formations on Earth. The ISB assemblage comprises a varietyof metavolcanic, clastic and metasedimentary rock types, thestratigraphy of which has been discussed in detail by previouscontributions (Boak and Dymek, 1982; Dymek and Klein, 1988;Myers, 2001; Nutman and Friend, 2009; Nutman et al., 1984, 1996;Rosing et al., 1996; van Zuilen et al., 2003). Protolith identification ofrocks in the ISB is complicated owing to the diversity of lithologicalunits observed in the ISB and metamorphic and/or metasomaticoverprint. For this study, metacarbonate sequences were targeted inorder to better identify their probable origin (chemical sedimentaryversus metasomatic) and possible association with genuine iron-formations. Iron-poor metacarbonates occur typically at the contactswith ultramafic intrusions in the southern ISB and have beeninterpreted as originating by leaching of ultramafic protoliths bymetasomatic fluids (Rose et al., 1996; Rosing et al., 1996). Iron-richmetacarbonates occur within a range of host rocks in the northern ISB,including in association with banded magnetite–quartz rocks of achemical sedimentary origin. We report and contrast here the carbonisotopic ratios of iron-rich and iron-poor metacarbonates andcombine these new data with existing iron isotopic data for thesame samples (Dauphas et al., 2007) to better constrain their primaryformation, in particular identifying whether the isotopic signaturesrecord a chemical sedimentary origin despite metasomatic overprint.

2.2. Sample preparation and analytical methods

Samples of magnetite and iron-rich carbonate from the BrockmanIron Formation were originally taken from drill core material andseparated and crushed by Richard Becker to obtain milligram-sizedmonomineralic powders for the analysis of carbon and oxygenisotopes (Becker and Clayton, 1972, 1976). In addition, samples ofiron-poor carbonate from the Wittenoom Dolomite were cut fromunweathered faces of hand specimens and approximately one gram ofthis material was crushed to a powder using an agate mortar andpestle. Magnetite and iron-rich and iron-poor carbonates from the ISBwere taken from powders of hand specimens originally prepared foriron isotopic analysis (see Dauphas et al., 2007).

Iron isotopic analyses of samples from the Brockman Iron Formationand Wittenoom Dolomite were carried out following the proceduresdeveloped in our laboratory (Craddock and Dauphas, 2010; Dauphas etal., 2004a, 2009b). Iron isotopic measurements were performed on a

-1.5 -1.0 -0.5 0.0 0.5 1.00

5

10

15

-14 -12 -10 -8 -6 -4 -2 0 2

freq

uenc

y

Fe-rich carbonate

Fe-poorcarbonate

Kuruman~2.5 Ga

magnetiteFe-rich carbonate

Fe-poorcarbonate

-2.0

0

5

10

15

freq

uenc

yseawater

Fe2+aq

seawater HCO3

-, aq

magnetite Fe-rich carbonate

Fe-poorcarbonate

Hamersley~2.5 Ga

Fe-rich carbonate

Fe-poorcarbonate

δ56Fe, IRMM-014 (‰) δ13C, V-PDB (‰)

Fig. 1. Histograms of iron (δ56Fe) and carbon (δ13C) isotope ratios in iron-rich carbonates(siderite, ankerite), iron-poor carbonates (calcite, dolomite) andmagnetite from the ~2.5 GaBrockman Iron Formation, Hamersley Basin (top panel) and Kuruman Iron Formation,Transvaal Craton (bottompanel). Iron isotope data for theBrockman Iron Formation are fromTable 1. Carbon isotope data for the Brockman Iron Formation are from Becker and Clayton(1972). Carbon and iron isotope data for the Kuruman Iron Formation are from Johnson et al.(2003) and Heimann et al. (2010). Bin widths for iron and carbon isotope values are 0.025‰and 1.0‰, respectively. The vertical arrows at δ56Fe=−0.2‰ and δ13C=0‰ are theestimated compositions of Fe(II)aq anddissolved inorganic carbon (DIC) inArchean seawater,respectively. The range of iron and carbon isotopic ratios in these samples are incompatiblewith formation of magnetite and iron-rich carbonates as direct precipitates from seawater,but can be reconciled by a model invoking diagenetic mobilization of primary ferric oxidesand organic carbon through dissimilatory iron reduction (DIR).

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Thermo Scientific Neptune MC-ICPMS at the University of Chicago. Alliron isotope data are calibrated relative to the IRMM scale (Craddockand Dauphas, 2010; Taylor et al., 1992).

Table 2Iron and carbon isotopic compositions of iron-poor carbonate hand specimens from the W

Samplea Lithology δ13C(‰)b

Calcite Dolomite

Wittenoom Dolomite Formation12523 Dolomite 0.05±0.0112523 vein Dolomite vein 0.38±0.0213354 Calcite+quartz −1.09±0.0213355 top Calcite±dolomite −1.80±0.02 −1.5613355 bottom −1.83 −1.8413356 Calcite −0.13±0.0513357 chert Calcite+quartz −1.2213358 #1 Calcite −0.30±0.0513358 #213359 chert Black chert+calcite −1.06±0.0513360 Calcite −0.82±0.06RB-22a* Calcite+silicate −4.91±0.10RB-22b* −4.68±0.02RB-23a* Calcite+dolomite+ −6.36±0.04RB-23b* silicate −4.75±0.03 −5.67±0.03

a Sample numbers follow the nomenclature adopted by Becker (1971). Detailed descriptb Carbon isotopic (δ13C) ratios are per mil difference relative to PDB standard and are t

deviation of replicate analyses as reported by Becker and Clayton (1972).c Iron isotopic (δ56Fe, δ57Fe) ratios are reported inpermil difference relative to ironmetal stan

iron-poor carbonates sampled from the base of the Wittenoom Dolomite that occur as fine-layappearance to more massive calcite and dolomite sampled elsewhere in the Wittenoom Dolom

Carbon isotopic analyses were carried out on iron-rich and iron-poor metacarbonates from the ISB at RSMAS, University of Miami. Theprocedures followed those developed by Swart et al. (1991). Carbonisotope compositions are reported relative to the V-PDB scale.

3. Results

Iron isotopic (δ56Fe) compositions of magnetite samples from theBrockman Iron Formation, Hamersley Basin, analyzed in this study rangefrom −0.3 to +1.2‰ (Table 1; Fig. 1). These data are similar to thosemeasured in magnetite samples from the same formation by Johnsonet al. (2008). In the two studies combined, magnetite from the BrockmanIron Formation extends down to δ56Fe=−1.0‰ and has a mean ironisotopic composition of +0.17‰ (n=94). Two types of carbonates fromthe Hamersley Basin were studied: iron-rich carbonates (siderite,ankerite) from the Brockman Iron Formation and iron-poor carbonates(calcite, dolomite) from the underlying Wittenoom Dolomite [usingnotations from van Zuilen et al. (2003) and Dauphas et al. (2007), iron-rich carbonates have an Fe atomic ratio, Fe/(Fe+Mn+Ca+Mg), greaterthan 0.40,whereas iron-poor carbonates have an Fe atomic ratio less than0.15]. Theδ56Fevaluesof iron-rich carbonates analyzed in this study rangefrom−1.1 to +1.2‰ (Table 1; Fig. 1). Considering the iron isotopic datafor siderite published by Johnson et al. (2008), the range extends down to−2.0‰. The iron isotopic compositions of magnetite and iron-richcarbonates sampled fromadjacent bands in theBrockman Iron Formationhave a broadly positive correlation. The δ56Fe values of iron-poorcarbonates from the Wittenoom Dolomite analyzed in this study are allverynegative, between−0.5 and−1.0‰ (Table 2; Fig. 1). Carbon isotopecompositions (δ13C) of the carbonate samples analyzed in this study fortheir iron isotopic compositions were previously reported by Becker andClayton (1972). All iron-rich carbonates have light δ13C compositions(Fig. 1). Siderite δ13C range from−12.6 to−7.5‰ and cluster around theaverage of −9.6‰ (n=11). Ankerite δ13C are slightly more variable,ranging from −15.0 to −6.5‰, but cluster around a similar average of−9.9‰ (n=32). Iron-poor carbonates from the Wittenoom Dolomitehave a limited range of heavier δ13C with an average −2‰ (n=15;Fig. 1). For comparison, the iron and carbon isotope compositions of iron-rich and iron-poor carbonates and of magnetite from the penecontem-poraneous Kuruman Iron Formation, Transvaal Supergroup, South Africa(Heimann et al., 2010; Johnson et al., 2003) are illustrated (Fig. 1). The

ittenoom Dolomite, Hamersley Basin.

δ56Fe(‰)c

δ57Fe(‰)c

Calcite Dolomite Calcite Dolomite

−0.995±0.068 −1.427±0.092−0.774±0.068 −1.135±0.092

−0.728±0.055 −1.059±0.038−0.472±0.055 −0.674±0.038

−0.921±0.055 −1.313±0.038−0.860±0.055 −1.344±0.038−0.882±0.055 −1.338±0.038−0.450±0.055 −0.692±0.038−0.887±0.057 −1.366±0.085−0.685±0.057 −1.024±0.085

ions given by Becker (1971).hose published by Becker and Clayton (1972). Quoted uncertainties are the standard

dard IRMM-014 (δ57Fe~1.5xδ56Fe). Uncertainties are 95% confidence intervals. * Indicatesers in apparently microbanded rocks containing silica and silicate and that are distinct inite (see Becker, 1971). These samples were not available for the analysis of iron isotopes.

Table 3Iron and carbon isotopic compositions of bulk metacarbonates from Isua Supracrsutal Belt, southern West Greeland.

Sample Description Fe (wt.%) δ13C (‰) a δ56Fe (‰) b δ57Fe (‰) b

Fe-rich metacarbonatesIS-04-01 Siderite, northwest of belt 23.6 −5.50±0.06 0.236±0.059 0.296±0.093IS-04-02 Siderite, northwest of belt 23.7 −4.52±0.05 0.318±0.125 0.500±0.169IS-04-03 Siderite, northwest of belt 24.7 −4.73±0.02 0.448±0.074 0.763±0.151IS-04-07 Metacarbonate mix, east of belt 59.4 −4.10 0.759±0.038 1.129±0.074IS-04-07* Replicate, this study 58.1 0.743±0.039 1.098±0.045IS-04-08 Siderite, east of belt 31.1 −4.52±0.06 0.400±0.034 0.587±0.048IS-04-09 Graphite-rich metacarbonate 47.3 −4.75±0.05 0.294±0.069 0.496±0.165IS-04-10 Siderite, east of belt 49.6 −5.94±0.05 0.281±0.084 0.400±0.110IS-04-11 Graphite-rich metacarbonate 45.7 −4.58±0.04 0.204±0.056 0.383±0.114

Fe-poor metacarbonatesIS-04-05 Calcite/dolomite, west of belt 2.6 −1.98±0.07 −0.742±0.046 −1.096±0.093IS-04-05* Replicate, this study 2.9 −0.754±0.037 −1.138±0.043AL-04-G18 Carbonated Amitsoq gneiss 3.2 −1.52±0.04 −0.901±0.177 −1.456±0.219Al-04-G23 Carbonated Amitsoq gneiss 2.5 −0.34±0.06 −0.725±0.177 −1.051±0.219

a Carbon isotopic (δ13C) ratios are reported in per mil difference relative to the V-PDB standard. Uncertainties are two standard deviation for replicate analyses of the same sample.b Iron isotopic (δ56Fe, δ57Fe) ratios are those previously published by Dauphas et al. (2007b), except for two samples IS-04-07 and IS-04-05 indicated by *, which were

independently measured for their iron isotopic composition as part of the present study for confirmation of accuracy. Iron isotopic ratios are reported in per mil difference relative toiron metal standard IRMM-014. Quoted uncertainties are 95% confidence intervals.

126 P.R. Craddock, N. Dauphas / Earth and Planetary Science Letters 303 (2011) 121–132

data showadistributionof ironandcarbon isotopic ratios that is similar tothat documented in the Hamersley Basin.

To compare the isotopic signatures of metacarbonates from the~3.8 Ga Isua Supracrustal Belt (ISB) with those from genuine lateArchean BIFs, we combine new carbon isotope data of iron-rich andiron-poor metacarbonates measured in this study with previouslyreported iron isotope data for the samemetacarbonates fromDauphaset al. (2007). Iron-poor metacarbonates have light δ56Fe (−0.90 to−0.73‰), whereas iron-rich metacarbonates have heavy δ56Fe(+0.24 to +0.76‰; Table 3, Fig. 2). Iron-poor metacarbonates haveδ13C values between−2.0 and 0‰, whereas iron-richmetacarbonateshave distinctly lighter δ13C values between −6.0 to −4.1‰ (Table 3;Fig. 2). Several studies have previously reported δ13C values ofmetacarbonates from the ISB (Oehler and Smith, 1977; Perry andAhmad, 1977; Schidlowski et al., 1979; Ueno et al., 2002; van Zuilenet al., 2003); only van Zuilen et al. (2003) have distinguishedgeochemically between the carbon isotopic signatures of iron-poorand iron-rich variants and none have iron isotopic data for the samesamples that enable a direct comparison with our study.

Isua~3.8 Ga

0

5

10

freq

uenc

y

Fe-poormetacarbonate

Fe-rich metacarbonate

magnetite

Fe-poormetacarbonate

Fe-rich metacarbonate

-1.5 -1.0 -0.5 0.0 0.5 1.0 -14 -12 -10 -8 -6 -4 -2 0 2-2.0

seawater Fe2+

aq

seawater HCO3

-, aq

δ56Fe, IRMM-014 (‰) δ13C, V-PDB (‰)

Fig. 2. Histograms of iron (δ56Fe) and carbon (δ13C) isotope ratios in iron-richmetacarbonates (siderite, ankerite), iron-poor metacarbonates (calcite, dolomite) andmagnetite from the ~3.7 to 3.8 Ga Isua Supracrustal Belt, SW Greenland. Iron isotopic dataare fromDauphas et al. (2007). Binwidths for iron and carbon isotope values are 0.025‰ and1.0‰, respectively. The vertical arrows are the same as those depicted in Fig. 1. Despite acomplexmetamorphic history andmetasomatic overprint, the iron and carbon isotopic dataare similar to those of genuine late-Archean BIFs and are compatible with derivation of iron-rich metacarbonates in the ISB from reduction of ferric oxides (ferrihydrite), probablycoupled to oxidation of organic carbon.

4. Discussion

4.1. Evidence for an authigenic origin of iron-rich carbonates andsedimentary microbial iron respiration at 2.5 Ga

Iron-rich carbonates in BIFs from the Hamersley Basin exhibit awide range of iron and carbon isotopic compositions. Can this isotopicheterogeneity be explained entirely by precipitation of dissolved Fe(II) and inorganic carbon in the water column in near-isotopicequilibrium? Or, are the isotopic signatures consistent with anauthigenic origin in marine sediments, possibly involving microbialiron and carbon respiration?

The δ13C of iron-poor carbonates of different ages ranging from~3.8 to 2.5 Ga are remarkably uniform and similar to that of platformcarbonates deposited throughout the Phanerozoic (~0‰) (Becker andClayton, 1972; Schidlowski et al., 1975; Shields and Veizer, 2002;Veizer et al., 1989, 1992). This argues for no significant changes (i.e.,b10‰) in the bulk carbon isotopic reservoir of dissolved inorganiccarbon (DIC) in seawater (c.f. Beukes et al., 1990; Kaufman et al.,1990). The uniform δ13C of iron-poor carbonates deposited synchro-nously in the late Archean across a range of water column depths fromcontinental shelves (platform) to abyssal plains (basinal) argues alsoagainst vertical stratification with respect to carbon isotopes of DIC atany given time (Fischer et al., 2009). Based on the experimentallydetermined carbon isotopic fractionation between siderite and DIC,Δ13CHCO3-Sid=−0.5±0.2‰ (Jimenez-Lopez and Romanek, 2004),siderite precipitated from the water column in isotopic equilibriumwith DIC should have near-uniform δ13C ~0±2‰ (Fig. 3). The rangeof δ13C between −6 and −15‰ documented in iron-rich carbonatesfrom the Brockman Iron Formation in the Hamersley Basin is clearlydifferent from that expected for, and argues against, formation ofthese carbonates in isotopic equilibriumwith Archean seawater (Bauret al., 1985; Becker and Clayton, 1972).

Evidence from iron isotope data of the same iron-rich carbonatesfurther argues against primary precipitation in Archean seawater. Theprevailing consensus is that iron was delivered to the Archean ocean asferrous Fe(II)aq, primarily via hydrothermal activity (e.g., Bau andMöller,1993; Dymek and Klein, 1988; Jacobsen and Pimental-Klose, 1988). Theδ56FeFe(II) composition of modern hydrothermal fluids ranges from−0.1to −0.6‰, averaging −0.20‰ (n=19) (Beard et al., 2003; Severmannet al., 2004; Sharma et al., 2001). Given that the range of iron isotopiccomposition of Archean igneous rocks is limited and similar to modern(e.g., Dauphas et al., 2009a), hydrothermal alteration of oceanic crustin the Archean should contribute Fe(II) with an isotopic composition also

Brockman

Fe-rich

0

seawater

Isua, Fe-poor metacarbonates

Isua, Fe-rich metacarbonates

Kuruman

Fe-rich

Fe-poor carbonates

~2.5 Ga

~3.8 Ga

1.5

1.0

0.5

0.0

-0.5

-1.0

1.5

1.0

0.5

0.0

-0.5

-1.0

-1.5-16 -14 -12 -10 -8 -6 -4 -2

carbonate δ13C (‰)

carb

onat

e δ56

Fe

(‰)

carb

onat

e δ56

Fe

(‰)

Fig. 3. Coupled δ56Fe and δ13C ratios iron-rich and iron-poor carbonates from theBrockman Iron Formation (Becker and Clayton, 1972, this study) and Kuruman IronFormation (Heimann et al., 2010; Johnson et al., 2003) (top panel) and from the IsuaSupracrustal Belt (Dauphas et al., 2007, this study) (bottom panel). The white boxdelineates the estimated isotopic composition of Archean seawater. The gray box is thepredicted range of isotopic compositions for iron-rich carbonates precipitated directlyfrom and in isotopic equilibriumwith Archean seawater based on experimental (Jimenez-Lopez and Romanek, 2004; Wiesli et al., 2004) and theoretical estimates (Anbar et al.,2005; Blanchard et al., 2009; Schauble et al., 2001) of equilibrium isotope fractionationbetween Fe(II)aq-siderite and HCO3-siderite. Iron-poor carbonates from both the ~2.5 GaBrockman and Kuruman Iron Formations and the ~3.8 Ga Isua Supracrustal Belt have ironand carbon isotopic compositions very similar to that expected for carbonate precipitationin Archean seawater. In contrast, iron-rich carbonates from these iron-formations have awide range of iron and carbon isotopic ratios that are inconsistent with their precipitationin equilibrium with common seawater.

127P.R. Craddock, N. Dauphas / Earth and Planetary Science Letters 303 (2011) 121–132

~−0.20‰. It is likely that the Archean ocean, at least below the surfacemixed layer (~100 m),waswell-mixedandhomogeneouswith respect toiron isotopes because the reservoir of Fe(II)aq required to form BIFs mustbe considerable (20 to 100 mg Fe L−1) and continually replenished(Ewers and Morris, 1981; Morris, 1993). In effect, the residence time ofiron in the Archean ocean was significantly longer than the oceanicmixing time (Johnson et al., 2003). Iron-rich carbonates precipitated inequilibrium with Archean seawater are predicted to have light ironisotopic ratios between~−0.5 and−2‰ (Fig. 3), based on thenet isotopefractionation between carbonate (siderite) and Fe(II)aq determined byexperiments, Δ56FeSid–Fe(II)=−0.5±0.2‰ (Wiesli et al., 2004), andtheoretical calculations, Δ56FeSid–Fe(II)=−1.6 to −2.1‰ (Anbar et al.,2005; Blanchard et al., 2009; Schauble et al., 2001). Such low δ56Fe valuesare found in iron-poor carbonates from theWittenoomDolomite (Fig. 3),a platform carbonate formation in the Hamersley Group underlying theBrockman Iron Formation. Similar light δ56Fe values are observed inplatform carbonates associated with other late Archean and Paleoproter-ozoic BIFs (Heimann et al., 2010). In contrast, the iron isotopiccompositionsmeasured inmost iron-rich carbonates from the Brockman

Iron Formation are heavier than−0.5‰ and extendup to+1.2‰ (Fig. 3).These isotopic signatures cannot be explained by direct precipitation innear isotopic equilibrium with bulk Archean seawater.

If iron-rich carbonates were not directly precipitated from seawater,then what were the chemical and/or biological transformations of ironand carbon that prevailed to their formation? The range of light δ13Cratios measured in iron-rich carbonates from the Brockman IronFormation requires a source of carbon in the iron-formation with alight isotopic composition. Becker and Clayton (1972) have previouslyargued that themost reasonable source of this light carbon was organicmatter. The carbon isotopic compositions of fossil organic matter in theHamersley Basin and other Precambrian iron-formations fall in arestricted range of very light δ13C between−30 and−33‰ (Barghoornet al., 1977; Brocks et al., 1999; Perry et al., 1973). Fossil biomarkerspreserved in the Marra Mamba Iron and Jeerinah Formations of theHamersley Basin (Brocks et al., 1999) and in iron- and carbonate-formations of the penecontemporaneous Transvaal Supergroup (Wald-bauer et al., 2009) provide support for oxygenic/anoxygenic photosyn-thesis and organic carbon production in surface seawater in the lateArchean. Diagenetic oxidation of organic carbon would deliver a poolof isotopically-light carbon (as CO2 or dissolved carbonate) thatcould dilute the existing pore water reservoir of DIC and accumulatein marine sediments until saturation with respect to authigenic car-bonates was obtained. Whereas carbonates with δ13C ratios ~−30‰would have formed entirely from an organic-derived pool of carbonatecarbon, iron-rich carbonates in the Brockman Iron Formation (averageδ13C=−9.8±0.2‰) precipitated from a pore water reservoir ofcarbonate carbon contributed by both organic and DIC sources, thelatter possibly supported by partial exchange of DIC across thesediment–seawater interface.

Oxidation of organic carbon must be coupled to reduction of anappropriate electron acceptor. Inmodernmarine sediments the order ofoxidant consumption is O2NMn-oxidesNnitrateNFe-oxidesNsulfate(Froelich et al., 1979). In anoxic marine sediments of the Archean,oxygenwould have been severely limited (if not absent) as the primaryelectron acceptor, as would Mn-oxide and nitrate (e.g., Anbar andHolland, 1992; Anbar et al., 2007; Chapman and Schopf, 1983; Walker,1984). Banded iron-formations are anomalously rich in iron andreduction of ferric oxides may have permitted oxidation of organiccarbon (Dimroth and Chauvel, 1973; Walker, 1984). Studies havedemonstrated that oxidation–reduction between ferric oxide andorganic carbon is effectively mediated by microbes via the process ofDIR (Lovley, 1993), which can be illustrated by the stoichiometricreaction for ferrihydrite,

2Fe2O3d nH2O + CH2O + 7Hþ→4Fe2 + + HCO–3 + 2n + 4ð ÞH2O:

ð1Þ

Wewrite HCO3− as the carbonate product, reflecting the dominance of

this species in seawater at equilibrium (e.g., Walker, 1983). Previousstudies have applied simple mass balance models to estimate theoverall potential for Fe(III) reduction in the formation of iron oxidesand carbonates in BIFs and implicate a significant role for microbialiron respiration in the late Archean (e.g., Konhauser et al., 2005;Walker, 1984).

Our isotopic data for magnetite and iron-rich carbonates from theBrockman Iron Formation provide an independent assessment ofthese results and demonstrate that the isotopic signatures imprintedin BIFs are consistent with extensive microbial iron respiration at~2.5 Ga. Ferric iron in magnetite requires the oxidation andprecipitation of dissolved Fe(II)aq. The process by which largeamounts of Fe(II)aq were oxidized in the Archean ocean that wasanoxic is debated, but the primary Fe(III) precipitates in all cases werelikely amorphous ferric oxides (Ewers and Morris, 1981; Kappler andNewman, 2004; Morris, 1993; Trendall and Blockley, 1970), such as

-1.8

‰

-2.3

‰

1:1

Sid-Mgt

equilibrium

Δ

2.5

2.0

1.5

1.0

0.5

0.0

-0.5

-1.0-1.0 -0.5 0.0 0.5 1.0 1.5 2.0 2.5

magnetite δ56Fe (‰)

carb

onat

e δ56

Fe

(‰)

Fig. 4. Iron isotopic compositions of paired magnetite and iron-rich carbonates (siderite,ankerite) from the same laminations (chert-matrix mesobands) from the Brockman IronFormation. The isotopic fractionation between siderite and magnetite (light gray lines) isestimated to be between−1.8‰, based on laboratory experiments (Johnson et al., 2005;Wiesli et al., 2004), and −2.3‰, based on spectroscopy and theoretical calculations ofvibrational states of isotopic bonding (Blanchard et al., 2009; Polyakov et al., 2007). Thegray parallelogram delineates the range of iron isotopic ratios expected for iron-richcarbonates andmagnetite precipitated in isotopic equilibriumwith Archean seawater. Themeasured iron isotopic compositions of most magnetite and iron-rich carbonates areincompatible with this model. Note the overall positive correlation between the ironisotopic compositions ofmagnetite and siderite,whichmay reflect formation ofmost iron-rich carbonates and magnetite in association during diagenetic iron and carbon cycling.The blue band defines the 95% confidence interval of the regression (slope=0.69±0.33;the twodatawith light carbonate δ56Fe areomitted fromthis regressionbecause thesedatamay reflect near-isotopic equilibrium with seawater).

128 P.R. Craddock, N. Dauphas / Earth and Planetary Science Letters 303 (2011) 121–132

ferrihydrite. Experimental studies have demonstrated that amor-phous ferric oxides from partial Fe(II)aq oxidation are enriched in theheavy isotopes of iron relative to starting Fe(II)aq, with an equilibriumfractionation factor ranging from Δ56Fe ~ +0.9‰ for O2-mediatedoxidation (Bullen et al., 2001) to+2.0‰ for direct microbial oxidation(Balci et al., 2006; Beard et al., 2010; Croal et al., 2004). Thus, partialoxidation of Fe(II)aq in Archean seawater should produce ferrihydritewith positive δ56Fe values up to +1.5‰. Near-complete Fe(II)oxidation in surface seawater would produce ferrihydrite with aδ56Fe value of ~0‰, similar to that of the starting reservoir of Fe(II)aq,which is a lower limit on the range of δ56Fe in the primary ferricoxides. Magnetite in the Brockman Iron Formation was likely formedby the reaction of primary ferrihydrite with ferrous iron (Ewers andMorris, 1981; Morris, 1993):

Fe2O3d nH2O + Fe2 + + 2OH–→Fe3O4 + n + 1ð ÞH2O: ð2Þ

Magnetite formed by reaction between ferrihydrite and Fe(II)aq in thewater column can have positive δ56Fe values between ~0 and +1.0‰,reflecting the contributions of iron from ferrihydrite (δ56Fe ~0 to+1.5‰)and Fe(II)aq (δ56Fe~ −0.2‰) in the ratio 2:1. Measured δ56Fe ratios ofmagnetite from the Brockman Iron Formation are as light as −1.0‰,which is difficult to explain unless the pool of Fe(II)aq from whichmagnetite formedwasalsovery light, approaching−1 to−2‰. Themostlikely source of Fe(II)aq with very light δ56Fe values is from partialreduction of ferrihydrite in marine sediments via microbial DIR (Eq. (1);also see Johnson et al. 2008; Heimann et al. 2010). Indeed, experimentalstudies have documented a significant isotopic effect associated withpartial Fe(III) reduction via DIR that yields reduced Fe(II)aq depleted by−1.0 to −2.5‰ in δ56Fe relative to precursor ferric oxides (Beard et al.,1999, 2010; Crosby et al., 2007; Icopini et al., 2004; Johnson et al., 2005;Tangalos et al., 2010).

Iron-rich carbonates in the Brockman Iron Formation exhibit awide range of δ56Fe values (Fig. 1). Iron-rich carbonates with heavyδ56Fe compositions up to +1‰ could only have precipitated from apool of Fe(II)aq with positive δ56Fe. This reservoir of iron was notcommon seawater, but was an authigenic reservoir contributed fromnear-complete reduction in marine sediments of primary ferrihydritethat hadheavy δ56Fe values following partial Fe(II) oxidation in surfaceseawater. Microbial DIR (Eq. (1)) produces both dissolved Fe(II) andcarbonate that would accumulate in sediment pore waters untilsaturation with respect to iron-rich carbonates was attained:

Fe2 + + HCO–3→FeCO3 + Hþ

: ð3Þ

Light δ13C values of the same iron-rich carbonates (Fig. 3) havebeen shown to be consistent with derivation of carbonate ion fromoxidation of organic matter during microbial DIR (Baur et al., 1985;Becker and Clayton, 1972; Walker, 1984).

Additional support for microbial iron respiration is provided by thecorrelated iron isotopic ratios of magnetite and iron-rich carbonatessampled from the same mesobands and microbands within the Brock-man Iron Formation (Fig. 4). Following reduction of ferrihydrite,authigenic magnetite and siderite can precipitate from the samereservoir of Fe(II)aq. Thus, the δ56Fe signature of Fe(II)aq produced bymicrobial iron reduction can be inherited by both magnetite and iron-rich carbonate. Our interpretations are consistent with recent micro-analytical studiesof theBrockman IronFormation thathavedocumentediron isotopic variations in magnetite and iron-rich carbonates on spatialscales less than one millimeter, which are suggested to reflect post-depositional redistribution of iron within the sediment (Steinhoefelet al., 2010). Geochemical studies have shown that for the penecontem-paraneous Kuruman Iron Formation, carbonate facies contain higherresidual organic carbon contents up to an order of magnitude higherthan in oxide (magnetite) facies in the iron-formation (Beukes et al.,1990; Fischer et al., 2009). These observations are consistent with the

idea that the relative proportions of organic carbon and ferric ironcontrolled the diagenetic fate of the sediment; preserving magnetiteonly when ferric iron was in excess relative to organic matter.

Most, if not all, Archean and Paleoproterozoic BIFs have experi-enced varying degrees of metamorphism, which can affect the stablemineral assemblage observed in iron-formations (e.g., French, 1973;Klein, 1983, 2005). A possible burial metamorphic origin for carbonatein the Brockman Iron Formation via inorganic reaction betweenprimary ferric iron, such as ferrihydrite, and organic carbon atelevated temperature and pressure can, however, be discounted onthe basis of mass balance for iron and carbon (DIC) in BIFs. Duringmetamorphism, the sediment is isolated from exchange with theocean. The maximum concentration of inorganic carbonate in porewater at the onset of metamorphism can be estimated by assumingequilibrium with atmospheric CO2. The partial pressure, pCO2, of theArchean atmosphere is debated (e.g., Hessler et al., 2004; Kasting,1993; Lowe and Tice, 2004; Rosing et al., 2010; Rye et al., 1995;Walker, 1985), but pCO2 between 0.1 and 1 bar (≫100 times present)is a reasonable estimate. A kilogram of BIF in the Dales Gorge Memberof the Brockman Iron Formation contains on average 6 moles of iron(46.37 wt.% Fe2O3; Ewers and Morris, 1981). This quantity ofsediment corresponds to a pore volume of ~3.4×103 cm3 assuminga density of the iron-formation of ~3.4 g/cm3 (Ewers and Morris,1981) and accounting for post-depositional compaction by up to90% (Trendall and Blockley, 1970). This pore volume could containup to ~0.12 moles of total dissolved carbonate. According to reaction 4(net reaction between ferrihydrite and organic matter to yield Fe-rich carbonate; note that 3 moles of DIC is needed for each mole oforganic C):

2Fe2O3:nH2O þ CH2O þ 3HCO−3 þ 3H

þ→4FeCO3 þ ð2nþ 4ÞH2O; ð4Þ

4500 650 750700600550

5

6

7

8

9

Pre

ssur

e (k

bar)

Temperature (˚C)

Isua P-T

Siderite + Qtz = Ferrosilite + CO2

b

c

d

a

e

f

XCO2 = 1

Fig. 5. Pressure–temperature phase relationship for the reaction siderite+quartz→ferrosilite+CO2,g (shown by black line, a). Shown in the solid gray lines are equivalentphase boundaries for magnesite+quartz→enstatite+CO2,g (b), dolomite+quartz→diopside+CO2,g (c), ankerite+quartz→hedenbergite+CO2,g (d), calcite+enstatite+quartz→diopside+CO2,g (e) and calcite+ferrosilite+quartz→hedenbergite+CO2,g (f).Phase boundaries are reproduced from Lamb (2005). The peakmetamorphic P–T conditions(~5 kbar, 550 °C) to which rocks in the ISBwere subjected (e.g., Boak and Dymek, 1982) areshown by the gray box. The phase relations indicate that siderite (and other carbonates)would have been stable during peak metamorphism.

129P.R. Craddock, N. Dauphas / Earth and Planetary Science Letters 303 (2011) 121–132

this reservoir of dissolved carbonate would be sufficient to convert upto 0.15 moles of iron in ferrihydrite to siderite, which is only ~2% ofthe total inventory of Fe in the iron-formation. Magnetite and ironcarbonate constitute on average 30 wt.% and 15 wt.%, respectively, ofthe Dales Gorge Member in the Brockman Iron Formation (Trendalland Blockley, 1970). Assuming carbonate is siderite, the average molefraction of iron present as carbonate is 25% (assuming ankerite gives aminimum of 10%), which is considerably greater than the 2% of Fe thatcould be converted to carbonate through metamorphism. Thecalculations imply that an external source of carbonate is requiredto yield the large quantities of iron-rich carbonates observed in lateArchean iron-formations. We suggest that this carbonate was madeavailable both from oxidation of organic carbon and through partialexchange of DIC between sediment pore water and the overlyingwater column, and could only proceed during early diagenetictransformation. This idea is consistent with several petrographicstudies that suggested an early diagenetic origin for iron-richcarbonates in BIFs (e.g., Ewers andMorris, 1981; Morris, 1993; Pecoitset al., 2009; Trendall and Blockley, 1970).

Finally, we note that an interpretation to explain the light δ56Fevalues of minerals from late Archean BIF sequences (e.g., pyrite inshale units) has been proposed by Rouxel et al. (2005) wherebypartial oxidation and precipitation of ferric oxides in BIFs in the watercolumn leaves the residual Fe(II)aq reservoir enriched in the lightisotopes of iron. This reservoir is subsequently precipitated as ferrous-bearing minerals. We have shown that the light δ56Fe values of iron-bearing minerals in BIFs can be produced by iron cycling duringdiagenesis and concur with Johnson et al. (2008) that the highly-fractionated and stratigraphically variable iron isotopic compositionsof iron oxides and carbonates in oxide facies BIFs primarily reflectdiagenetic pathways for their formation.

4.2. Iron and carbon isotope evidence to support microbial ironrespiration in the early Archean

We now turn our attention to interpreting the iron and carbonisotope signaturesofmetacarbonates fromthe~3.8 Ga Isua SupracrustalBelt (ISB). The origin of ISB metacarbonates is contentious, withprotolith identification ranging from sedimentary (Bolhar et al., 2004;Dymek and Klein, 1988; Mojzsis et al., 1996; Nutman et al., 1984) toentirely metasomatic (Myers, 2001; Rose et al., 1996; Rosing et al.,1996). Iron-richmetacarbonates were initially interpreted as carbonatefacies iron-formations (Dymek and Klein, 1988; Nutman et al., 1984). Inthis interpretation framework, metacherts and calc-silicates associatedwith iron-richmetacarbonateswere formed by reaction of sedimentarycarbonates and quartz during burial and high-grade metamorphism,and the banding documented in metacherts was indicated to preservethat of the sedimentary protolith (Dymek and Klein, 1988; Nutmanet al., 1984). Trace element characteristics (e.g., REE+Y) of these rockswere interpreted as being consistent with their deposition as chemicalsediments in seawater (Bolhar et al., 2004; Frei and Polat, 2007).Alternatively, iron-rich and iron-poormetacarbonates and calc-silicateshave been interpreted as metasomatic in origin, formed by carbonationand desilication of igneous protoliths. A metasomatic origin for somemetacarbonates, in particular iron-poor variants, was indicated by thediscordant nature of calc-silicate and metacarbonate units, veining,replacive textures and of the igneous lithologies within which theseunits occur (Rose et al., 1996; Rosing et al., 1996).

In a recent publication, Dauphas et al. (2007) have reported the ironisotopic compositions and trace element geochemistry of iron-rich andiron-poor metacarbonates from the ISB in order to better distinguish ametasomatic versus sedimentary origin. Iron-poor and iron-richmetacarbonates have light and heavy δ56Fe ratios, respectively(Fig. 2). The light δ56Fe of iron-poor metacarbonates was interpretedas reflecting mobilization of isotopically light iron from pre-existingmafic or ultramafic protoliths by fluid, consistent with a metasomatic

origin. In support, alteration of oceanic crust at modern seafloorhydrothermal environments yields secondary mineral assemblageswith a heavy iron isotopic composition and an implied fluid with acomplementary light iron isotopic ratio (Rouxel et al., 2003). Heavyδ56Fe ratios up to +0.80‰ in iron-rich metacarbonates, however, areinconsistent with mobilization by metasomatic fluids of isotopicallylight iron from an igneous protolith. Dauphas et al. (2007) showed thatiron-rich metacarbonates had similar heavy iron isotopic compositionstomagnetite fromBIFs but the studywas inconclusive as to the nature ofthis relationship. New carbon isotopic data coupled to existing ironisotopic data for the iron-rich and iron-poor metacarbonates from theISB provide further constraints on a possible chemical sedimentaryversusmetasomatic origin for themetacarbonates. An authigenic origin,similar to that indicated for iron-rich carbonates from the youngerHamersley Basin, would implicate the evolution of microbial metabolicpathways (oxygenic or anoxygenic photosynthesis, DIR) by ~3.8 Ga.

The ISB metacarbonates fall into two isotopically distinct groups(Fig. 3). Iron-poor metacarbonates that have light δ56Fe have near-zeroδ13C ratios (−3 to 0‰), whereas Fe-rich metacarbonates that haveheavy δ56Fe have distinctly lighter δ13C ratios (mean of −4.8±0.6‰).An important consideration in interpreting these data is whether theseisotopic signatures are primary and record the conditions of precipita-tion, or if these have been disturbed by metamorphism. All rocks in theISB have been subject to amphibolite facies metamorphism, with peaktemperatures ~500 to 550 °C and pressures ~5 kbar (e.g., Boak andDymek, 1982). P–T phase relations of the reaction (Lamb, 2005),

FeCO3 + SiO2→ FeSiO3 + CO2;g ð5Þ

suggest that siderite formed during diagenesis would have survivedpeak metamorphic conditions (Fig. 5).

We interpret the different iron and carbon isotopic compositions ofiron-poor and iron-rich metacarbonates in the ISB as reflecting theiroriginal formation through distinct pathways. The field relationshipbetween iron-poor metacarbonates and ultramafic host rocks aresupportive ofmetasomatic overprint by leaching of iron by a CO2-bearingfluid from an ultramafic protolith (Rose et al., 1996; Rosing et al., 1996).The light iron isotopic compositions of iron-poor metacarbonates arepossibly consistent with derivation from iron mobilized from ultramaficrocks (Dauphas et al., 2007). The carbon and iron isotopic data combined,however, reveal that the isotopic characteristics of iron-poor metacarbo-nates from the ISB are very similar to those of iron-poor carbonates

130 P.R. Craddock, N. Dauphas / Earth and Planetary Science Letters 303 (2011) 121–132

preserved in association with late Archean and Paleoproterozoic BIFs(Fig. 3). By analogy, these isotopic signatures can insteadbe interpretedasreflecting formation of iron-poormetacarbonates as primary precipitatesin Archean seawater. The REE patterns of these carbonates also appear tobe supportive of their formation as chemical precipitates from seawater(Dauphas et al., 2007). Thus, while iron-poor metacarbonates havepreviously been suggested to be entirely metasomatic in origin, wecaution that metasomatic mobilization of a primary (i.e., chemicalsedimentary) carbonate cannot be excluded. Continued study of theisotopic and chemical signatures of iron-poormetacarbonates is requiredto unambiguously determine the primary derivation of these rocks.

Iron-richmetacarbonates from the ISB have iron and carbon isotopiccompositions similar to thoseof other iron-rich carbonates of knownBIFassociation (Fig. 3). To the extent that iron-rich metacarbonates of theISB share the same isotopic characteristics with those of late Archeaniron-formations, theymay have formed by the same process. The heavyδ56Fe values of iron-rich metacarbonates cannot be explained bymetasomatic mobilization of isotopically-light iron from igneousprotoliths. Instead, it is now recognized from studies of genuine BIFsthat iron-rich carbonates of a known authigenic origin can carry heavyδ56Fe resulting from microbial iron respiration of primary ferric oxideprecipitates also with heavy δ56Fe (Heimann et al., 2010, this study).Magnetite in BIFs from the ISB has heavy δ56Fe (Dauphas et al., 2004b,2007; see also Whitehouse and Fedo, 2007). Microbial Fe(III) reductionof this magnetite or other precursor ferric oxide would transfer theheavy isotopic composition to the ferrous iron product and subse-quently to iron-rich carbonates, which is exactly that observed. Furtherevidence that microbial iron respiration (DIR) was involved in theauthigenic formation of iron-rich metacarbonates comes from the lightδ13C values of these samples. Direct evidence for reduced carbon ofbiogenic origin in the ISB is disputed (e.g., Perry and Ahmad, 1977;Mojzsis et al., 1996; Rosing, 1999; Schidlowski et al., 1979; van Zuilenet al., 2002; 2003). Still, the light carbon isotopic compositions of iron-richmetacarbonates from the ISBmirror thosemeasuredof lateArcheanBIFs (Baur et al., 1985; Becker and Clayton, 1972; Heimann et al., 2010),which are consistent with derivation of carbonate from oxidation oforganic carbon during iron respiration. We conclude that iron-richmetacarbonates associated with iron-formation in the ISB formed bysimilar microbially-mediated iron and carbon transformations asdocumented in late Archean BIFs. The antiquity of microbial ironrespiration, as suggestedbyour ironand carbon isotopedata fromoneofthe oldest known sedimentary sequences, is consistent with results ofphylogenetic studies. Ferric iron reduction has been documented as ametabolic pathway within a large diversity of extent Bacteria andArchaea including those most closely related to the last commonancestor of modern life, which points toward iron respiration as one ofthe earliest forms of microbial metabolism (Liu et al., 1997; Lovley,1991; Vargas et al., 1998; Weber et al., 2006).

5. Summary and conclusions

Iron and carbon isotopic analyses of iron oxides and carbonates inBIFs can be used to constrain the pathways of their formation. Here, wereport the iron (δ56Fe, vs. IRMM-014) and carbon (δ13C, vs. V-PDB)isotopic compositions of magnetite and of iron-rich and iron-poorcarbonates from the 2.5 Ga Brockman Iron Formation in the HamersleyBasin, Australia, and the ~3.8 Ga Isua Supracrustal Belt (ISB), WestGreenland. The key results and implications from this study are:

1. Magnetite and iron-rich carbonates (siderite, ankerite) from theHamersley Basin preserve a wide range of δ56Fe values that areincompatible with direct precipitation in isotopic equilibrium withArchean seawater.

2. Magnetite with light δ56Fe (≪0‰) must have precipitated from apool of Fe(II)aq with light δ56Fe between−1 and−2‰, which waslikely produced through microbial partial reduction of ferrihydrite.

3. Iron-rich carbonates with heavy δ56Fe (up to +1.2‰) must haveprecipitated from a reservoir of Fe(II)aq with very positive δ56Fe.The source of heavy δ56Fe–Fe(II)aq was likely from near-completemicrobial reduction of ferric oxides (ferrihydrite) in marinesediments. The light δ13C of the same iron-rich carbonates wasderived from carbonate produced by the oxidation of organiccarbon that was probably coupled to reduction of iron.

4. The combined iron and carbon isotopic data support an authigenicorigin for iron-rich carbonates in the Hamersley Basin via coupledorganic carbon oxidation and ferrihydrite reduction. This process iseffectively mediated by microbes in marine sediments thoughdissimilatory iron reduction (DIR), which implicates extensivemicrobial iron respiration in the formation of late Archean BIFs.

5. Iron-rich metacarbonates from the ~3.8 Ga ISB have δ56Fe and δ13Csignatures similar to those of carbonates in late Archean BIFs. Theisotopic data are interpreted as reflecting formation of these iron-rich metacarbonates as marine authigenic precipitates throughmicrobial iron respiration. Despite metasomatic overprint, iron-rich metacarbonates in the ISB preserve primary isotopic char-acteristics supporting evolution of microbial iron catabolism by~3.8 Ga during the formation of some of the oldest recognizedsedimentary-bearing rocks on Earth.

Acknowledgments

Drill core material of the Brockman Iron Formation from Hamersleywasmade available by the Geological Survey ofWestern Australia. R. N.Clayton provided hand specimens of the Wittenoom Dolomite fromHamersley.Wegratefully acknowledge contributions to this researchbyR. H. Becker who previously carried out the mineral separation of theHamersley samples for carbon and oxygen isotope analysis. Metacarbo-nate samples from Isua were provided byM. van Zuilen and A. Lepland.Themanuscript benefited from discussionswith R. N. Clayton, and fromconstructive reviews by K. Konhauser and an anonymous reviewer. Thisresearch was supported by National Science Foundation through grantEAR-0820807 (Geobiology), National Aeronautics and Space Adminis-tration through grant NNX09AG59G (Cosmochemistry) and a PackardFellowship to N.D.

Appendix A. Supplementary data

Supplementary data to this article can be found online atdoi:10.1016/j.epsl.2010.12.045.

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