Iron isotopes constrain biogeochemical redox cycling of iron and manganese in a Palaeoproterozoic...

Earth and Planetary Science Letters xxx (2010) xxx–xxx

EPSL-10505; No of Pages 10

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Earth and Planetary Science Letters

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Iron isotopes constrain biogeochemical redox cycling of iron and manganese in aPalaeoproterozoic stratified basin

Harilaos Tsikos a,⁎, Alan Matthews b, Yigal Erel b, John M. Moore a

a Geology Department, Rhodes University, Grahamstown 6140, South Africab Institute of Earth Sciences, Hebrew University of Jerusalem, Israel 91904, Israel

⁎ Corresponding author.E-mail address: [email protected] (H. Tsikos).

0012-821X/$ – see front matter © 2010 Elsevier B.V. Adoi:10.1016/j.epsl.2010.07.032

Please cite this article as: Tsikos, H., et al., Irstratified basin, Earth Planet. Sci. Lett. (201

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Article history:Received 20 November 2009Received in revised form 14 July 2010Accepted 16 July 2010Available online xxxx

Editor: M.L. Delaney

Keywords:iron isotopesbanded iron-formationmanganesePalaeoproterozoicSouth Africa

The Hotazel Formation in the uppermost stratigraphic portion of the Neoarchaean–PalaeoproterozoicTransvaal Supergroup of southern Africa is an unusual sedimentary sequence of banded iron-formation (BIF)intercalated with three manganese-rich layers. As such, it is a succession that holds great potential to offer aunique view of one of the most dramatic transitions in early Earth history — the switch to a full oxidativecycle in shallow oceans at ca. 2.3 Ga. We present iron isotope results from BIF and Mn-rich samples collectedacross the entire Hotazel sequence, with a view to constraining processes of biogeochemical redox cyclingfor both metals close to the transition from a reducing to an oxidizing ocean–atmosphere system. Therecorded de-coupling of Fe- and Mn reduction during anaerobic organic carbon cycling in the Hotazel strata,suggests that manganese became an important electron acceptor in stratified marine environments of thePalaeoproterozoic during periods of increased primary manganese precipitation relative to iron. Verylow δ57Fe values registered across the entire Hotazel sequence and especially in manganese-rich samples(−2.4 to −3.5‰) signify deposition of iron and manganese in a terminal, stratified aqueous reservoir thatwas depleted in the heavy iron isotopes. These isotopic signatures, in conjunction with the unusualendowment of the Hotazel sequence in manganese, are interpreted to have evolved by Rayleigh distillationprocesses during protracted deposition of Mn-poor BIFs as preserved in the lower stratigraphic portion of theTransvaal Supergroup (Kuruman and Griquatown BIFs). The unique end-member geochemical and isotopiccharacteristics of the Hotazel rocks may therefore constitute a potential link between the widespreaddeposition of BIF during the Neoarchaean and Palaeoproterozoic, and the postulated rise in atmosphericoxygen levels around 2.3 Ga ago.

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on isotopes constrain biogeochemical redox cy0), doi:10.1016/j.epsl.2010.07.032

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1. Introduction

Iron and manganese are two of the principal electron acceptors inthe Phanerozoic oceanic biogeochemical cycle and are commonlydiscussed together in the literature due to their broad similarities interms of both geochemistry and microbiology (Canfield et al., 2005).The massive deposition of iron as banded iron formations (BIF) in theNeoarchaean and Palaeoproterozoic (2.7–2.2 Ga) also suggests thatredox cycling of iron essentially has been active throughout Earth'shistory. Consequently, major research emphasis has been placed onthe role of iron as a key redox proxy for the evolution of oxygen in theatmosphere–ocean system of the early Earth (Bekker et al., 2004;Rouxel et al., 2005; Kump 2005, 2008; Anbar et al., 2007; Kaufman etal., 2007; Kump and Barley, 2007; Lyons, 2007). Most BIFs, however,contain manganese at levels no higher than 0.5 wt.% on average(Klein, 2005). Accordingly, there has been limited discussion of

manganese as a direct biogeochemical tracer of stratified marineenvironments of the Neoarchaean and Palaeoproterozoic, particularlyin terms of biological versus abiotic processes of manganese oxidationin primary depositional environments and anaerobic reduction ofhigh-valence sedimentary precursors during early diagenesis. Anexplanation for the very low levels of manganese recorded in mostclassic BIF sequences was previously sought in terms of low rates ofMn photo-oxidation relative to Fe (Anbar and Holland, 1992). Morerecently, emphasis has shifted towards biologically-driven ironoxidation pathways (Konhauser et al., 2007).

In this paper, we address the potential significance of manganeseas a proxy of biogeochemical redox cycling and atmosphere–oceanevolution during Precambrian BIF deposition. Our approach is basedon the use of iron isotope data obtained from the Hotazel Formation, aPalaeoproterozoic BIF succession in South Africa that contains themost voluminous sedimentary deposit of manganese on the Earth'scontinents, in the form of three discrete sedimentary layers (Laznicka,1992). Previous research has demonstrated the potential of ironisotopes in tracking metal cycling during deposition and diagenesis ofBIFs on a variety of temporal scales (Rouxel et al., 2005; Johnson et al.,

cling of iron and manganese in a Palaeoproterozoic

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2 H. Tsikos et al. / Earth and Planetary Science Letters xxx (2010) xxx–xxx

2003, 2008a,b; Heimann et al., 2010). The strikingly inter-layerednature of the Hotazel Fe–Mn sequence and presence of iron as acommon component in both the host BIF and the manganese-richlayers, provide a unique archive for examining the biogeochemicalcycling of both metals in a common depositional palaeoenvironmentusing iron isotope geochemistry. The unusually large volume ofmanganese-rich sediment and very low iron isotope values across theentire Hotazel sequence recorded in this study, suggest thatmanganese must have been, at least transiently, an importantcomponent in mechanisms of redox cycling in Neoarchaean–Palaeo-proterozoic stratified basins. The geochemical and isotopic signaturesof the Hotazel deposits are assessed in the context of BIF depositionalmodels, biological evolution, and the history of atmosphere–oceanoxygenation in the early Earth.

2. Geological background

The Hotazel Formation is an inter-layered succession of BIF andmanganese-rich sedimentary rock in the stratigraphically uppermostsegment of the Palaeoproterozoic Transvaal Supergroup in theNorthern Cape Province of South Africa (Fig. 1A). The lower part ofthe Transvaal Supergroup in this area (Ghaap Group) spans the agebracket 2.64–2.45 Ga (Pickard, 2003; Beukes and Gutzmer, 2008) andis dominated by a thick succession of largely dolomitic carbonates atthe base (Campbellrand Subgroup), overlain by the several hundredmeter thick BIF of the Kuruman–Griquatown Formations (Klein andBeukes, 1989; Beukes and Klein, 1990; Beukes et al., 1990; Beukes andGutzmer, 2008). Stratigraphically above the latter, a glacial diamictiteunit known as theMakganyene Formation (Polteau et al., 2006)marksthe base of the upper portion of the Transvaal Supergroup (the

Fig. 1. (A) Geographical distribution of the Transvaal Supergroup in South Africa (inset) anProvince, including the locality of the Kalahari manganese field. (B) Outline of the Kalahari mfour drill-cores used in this study.

Please cite this article as: Tsikos, H., et al., Iron isotopes constrain biogeocstratified basin, Earth Planet. Sci. Lett. (2010), doi:10.1016/j.epsl.2010.0

Postmasburg Group; Fig. 1A). The Makganyene diamictite is overlainby a sequence of basaltic andesites belonging to the OngelukFormation, with characteristic pillowed structures in its upperportions suggesting sub-aqueous extrusion at least in part.

The Hotazel Formation occurs under a ca. 40 m-thick blanket ofTertiary sand and calcrete in an area of approximately 400 km2 knownas the Kalahari manganese field (Fig. 1A,B). It consists of threelaterally continuous manganese layers inter-bedded with BIF that liestratigraphically above the Ongeluk Formation and below Fe-bearingcarbonate rocks of the Mooidraai Formation (Fig. 2). The BIF portionsof the Hotazel succession are characterised by typically mm- to cm-scale laminated assemblages that are dominated by micro-crystallinequartz and magnetite in roughly equal modal proportions (Fig. 2).Diagenetic carbonates (ankerite, calcite) constitute the next mostimportant species in terms of modal abundance, whereas Fe silicates(mainly greenalite, minnesotaite and stilpnomelane) occur in muchsmaller modal amounts. This mineralogy suggests only very shallowburial metamorphic overprint (essentially below 200 °C) in theHotazel sediments (Gutzmer and Beukes, 1996; Tsikos et al., 2003).The inter-bedded manganese-rich layers are dominated by variablerelative proportions of very fine-grained braunite and hematite, aswell as considerable amounts (up to 40% modal on a hand-specimenscale) of manganiferous carbonate (kutnahorite and lesser man-ganoan calcite) in the form of mm-scale laminae and diageneticovoids (Gutzmer and Beukes 1996; Tsikos et al., 2003).

At the contacts between the BIF andmanganese layers, transitionallutitic lithologies develop, dominated by very fine-grained hematiteand diagenetic Mn-carbonate (kutnahorite, Mn-calcite, minor Mn-dolomite). The stratigraphic thicknesses of these transitional hematitelutites cannot be precisely defined due to their gradational contacts

d simplified regional geological map of the Transvaal Supergroup in the Northern Capeanganese field including mine localities, major structural features and localities of the

hemical redox cycling of iron and manganese in a Palaeoproterozoic7.032


Fig. 2. Representative stratigraphic log of the Hotazel Formation and adjacentstratigraphic units, as these typically develop in the northern Kalahari manganesefield. Broad mineralogical characteristics and relative stratigraphic thickness relationsbetween BIF, hematite lutite transitions and the three manganese-rich layers areshown. Mineral abbreviations: qz: quartz; mt: magnetite; hem: hematite; cc: calcite;ank: ankerite; gr: greenalite; min; minnesotaite; stl: stilpnomelane; kut: kutnahorite;Mn-cc: manganoan calcite; br: braunite.

3H. Tsikos et al. / Earth and Planetary Science Letters xxx (2010) xxx–xxx

with the adjacent BIFs and manganese units proper, but extend fromseveral decimeters up to a fewmeters in thickness, depending on bothstratigraphic level and geographic location (Fig. 2). The transitionallutites typically contain magnetite disseminations and mm-scalelaminae in immediate contact with the BIF units, but becomeprogressively magnetite-free and more braunite-rich relative tohematite (and thus more manganiferous) as the manganese unitsare approached. The stratigraphic extent of the manganese ore-unitsis therefore defined on the basis of vertical variations in bulkmanganese content and Mn/Fe ratios.

Relative primary thickness variations between the BIF and thethree manganiferous layers are substantial and broadly antitheticacross the Kalahari manganese field (Tsikos and Moore, 1997). Thelowermost and commercially most important manganese layer variesin stratigraphic thickness from a maximum of ~45 m in thesouthernmost part of the field, to less than 8 m in the northern partof the field where drill-core selection and sampling for this studywerecarried out. Bulk manganese metal contents for all three manganeselayers range between 15 and 45 wt.%, with the highest average values


generally seen in the stratigraphically lowermost layer. The threemanganese layers attain a collective stratigraphic thickness bet-ween 15 and 70 m across the field, and the entire Hotazel Formationreaches a preserved primary thickness in excess of 200 m in places,but over most of the field it is in the order of 100–150 m (Tsikos et al.,2003). Post-Hotazel erosional events that have variably affected theHotazel succession are manifested by three unconformities developedduring the late Paleoproterozoic (Olifantshoek Supergroup), theCarboniferous (glacial scouring filled by Dwyka tillite), and theTertiary (sand and calcrete of the Kalahari Formation) (Tsikos et al.,2003).

3. Recent genetic modelling

Current debates on the genesis of BIF revolve around the exactmechanisms of primary iron precipitation, and specifically whetherFe(II) oxidation in the surface waters of stratified marine environ-ments of the Neoarchaean and Palaeoproterozoic was driven byphotosynthetically generated O2 or by anaerobic photo-autotrophicbacteria. Direct evidence for the onset of oxygenic photosynthesis inthe late Archaean is derived from the identification of cyanobacterialbiomarkers in 2.7–2.5 Ga organic carbon-rich shales from Australiaand South Africa (Brocks et al., 1999, 2003a,b; Eigenbrode andFreeman, 2006; Eigenbrode et al., 2008; Waldbauer et al., 2009). Thisevidence is disputed in recent studies that suggest an epigeneticorigin for at least some of these biomarker records (Rasmussen et al.,2008). On the other hand, laboratory-based studies and research onmodern natural analogues for BIF-style depositional environments,provide increasing support for the role in BIF deposition of anaerobicFe(II) oxidation by photo-autotrophic bacteria (Konhauser et al.,2002; Kappler et al., 2005; Crowe et al., 2008).

Recent models of the climatic and biological evolution across thePalaeoproterozoic place the Hotazel Formation in the aftermath of theearliest snowball Earth event, dated at approximately 2.3–2.2 Ga andrepresented by the Makganyene glacial deposits stratigraphicallybelow the Hotazel sediments (Kirschvink et al., 2000; Kopp et al.,2005; Kirschvink and Kopp, 2008). By implication, the deposition ofthe Hotazel Formation would also postdate the Great Oxidation Event(GOE; Holland, 2005), dated at ca. 2.3 Ga and manifested by severallines of independent geological and geochemical evidence, such as theloss of mass independent fractionation in sulphur isotopes (Bekker etal., 2004; Guo et al., 2009). The snowball Earth model proposes thatprimary precipitation of Mn(IV) leading to the formation of theHotazel manganese deposits would have required the availability offree oxygen during post-GOE times, and thus the pre-establishment ofoxygenic photosynthesis. It was argued that the latter commencedshortly before the Makganyene snowball event, and would havecontributed to the collapse of a methane greenhouse over short timescales.

Geochronological constraints with regard to the timing of thePalaeoproterozoic snowball Earth event and, by extension, thedeposition of the Hotazel Formation, are primarlily based on awhole-rock U–Pb model age of ~2.2 Ga for the Ongeluk volcanicsequence (Cornell et al., 1996). This age, however, is at odds with awhole-rock Pb–Pb carbonate age of ~2.4 Ga yielded by the Mooidraaicarbonate rocks overlying the Hotazel Formation (Bau et al., 1999). Onthese grounds, the age bracket for the deposition of the Hotazelsediments can only be broadly constrained between 2.4 and 2.2 Ga,and hence the precise temporal association of the Hotazel Formationwith at least the timing of the GOE remains inconclusive.

4. Sample selection and methods

Samples of Hotazel BIF, transitional hematite lutite andmanganeseore from four drill-cores obtained from the northern part of theKalahari manganese field (Fig. 1B) were selected for iron isotope



Table 1Iron isotope data and selected major element analyses of BIF, transitional hematitelutite and manganese ore samples from essentially unaltered Hotazel Formation.

Sample δ57Fe

(1)δ57Fe

(2)δ57Fe

(3)δ

57Fe Ca Fe Mg Mn Lithology

Mean wt.% wt.% wt.% wt.%

GL26-1 −2.15 −1.83 −1.99 5.31 38.07 0.57 0.45 BIF-2 −2.49 −2.24 −2.37 4.52 38.30 0.85 0.44 BIF-3 −1.15 −1.49 −1.32 7.87 33.07 0.70 0.39 BIF-4 0.05 −0.18 −0.07 20.69 22.99 0.35 0.55 BIF-5 −0.90 −0.83 −0.87 0.80 49.00 0.97 0.44 BIF-6 −0.93 −0.90 −0.92 10.93 30.17 1.27 3.57 BIF-7 −2.90 −2.76 −2.83 9.24 38.37 1.02 6.30 HL-8 −2.88 −2.91 −2.90 27.41 10.21 0.64 10.31 HL-9 −2.72 −2.57 −2.65 23.26 18.18 0.69 8.73 HL-10 −2.35 −2.06 −2.21 17.29 22.31 0.56 1.62 BIF-11 −0.29 −0.24 −0.27 0.73 45.61 0.08 0.30 BIF-12 −2.82 −2.62 −2.72 14.29 6.80 1.17 40.87 Mn ore-13 −3.34 −2.98 −3.16 10.15 11.27 4.90 34.09 Mn ore-14 −3.30 −3.34 −3.32 1.06 38.39 0.70 8.08 HL-15 −0.14 −0.41 −0.29 −0.28 0.11 49.80 0.40 0.55 BIF

BEL7A-1 −1.95 −1.97 −1.96 0.23 42.14 0.01 0.06 BIF-2 −1.58 −1.38 −1.48 0.19 47.46 0.01 0.08 BIF-3 −0.75 −0.77 −0.76 0.45 41.36 bd 0.57 BIF-4 −3.07 −2.84 −2.96 8.95 61.44 2.22 25.77 HL-5 −2.28 −2.16 −2.22 6.46 41.31 2.57 1.65 BIF-6 −2.55 −2.54 −2.55 1.60 39.27 1.95 1.68 BIF-7 −2.31 −2.30 −2.31 3.12 39.62 1.54 0.90 BIF-8 −2.69 −2.72 −2.71 4.70 29.37 2.36 1.40 BIF-9 −2.82 −2.65 −2.74 8.58 26.65 4.04 2.28 BIF-10 −2.35 −2.07 −2.21 1.11 50.32 0.47 0.38 BIF-11 −1.58 −1.70 −1.64 4.54 33.73 2.26 1.58 BIF-12 −1.01 −1.37 −1.19 3.60 42.52 1.84 1.31 BIF-13 −3.37 −3.19 −3.28 11.06 30.94 3.34 5.63 HL-14 −3.29 −3.23 −3.26 2.14 39.92 0.93 11.44 HL-15 −2.62 −2.47 −2.55 3.50 38.04 1.35 1.13 BIF-16 −1.64 −1.95 −1.80 1.97 26.61 0.28 0.24 BIF

BEL6B-1 −0.99 −1.20 −1.10 2.83 64.07 0.70 1.32 BIF-2 −1.41 −1.15 −1.28 0.05 39.58 0.18 0.33 BIF-3 −1.52† 2.27 42.93 0.43 0.37 BIF-4 −2.22 −2.03 −2.13 2.41 39.76 0.75 0.36 BIF-5 −2.10 −2.02 −2.06 2.35 30.01 0.89 0.44 BIF-6 −1.57 −1.34 −1.46 4.21 37.19 0.62 0.27 BIF-7 −1.61 −1.63 −1.62 4.94 38.00 0.43 0.15 BIF-8 −2.36† 7.85 45.39 0.08 0.88 BIF-9 −3.14 −3.03 −3.09 7.81 39.97 0.60 7.93 HL-10 −2.33 −2.50 −2.42 2.39 34.77 0.53 10.90 HL-11 −2.89 −2.40 −2.77 −2.69 2.29 52.57 0.71 0.14 BIF-12 −2.74 −2.45 −2.60 0.37 50.01 0.13 0.32 BIF-13 −2.07 −2.06 −2.17 −2.10 5.32 28.45 2.97 0.57 BIF-14 −3.48 −3.40 −3.68 −3.52 11.26 24.56 2.99 6.42 HL-15 −2.84 −3.00 −2.92 6.89 38.93 2.23 2.05 HL-16 −2.30 −2.59 −2.45 2.07 48.94 0.18 0.18 BIF

All samples tabulated in stratigraphic order for each drill-core.δ

57Fe values normalised against the IRMM-14 standard.

†: single δ57Fe analysis.

bd: below detection; BIF: banded iron-formation; HL: hematite lutite.

Table 2Bulk iron isotope data and selected major element analyses of altered BIF (hematite-quartz), Hotazel Formation.

Sample δ57Fe

(1)δ57

Fe(2)

δ57Fe Ca Fe Mg Mn

Mean wt.% wt.% wt.% wt.%

N12D-1 −2.27 −2.45 −2.36 1.88 72.91 0.05 2.47-2 −2.04 −2.17 −2.11 0.13 45.22 bd 0.03-3 −2.46 −2.29 −2.38 0.15 41.63 bd 0.02-4 −2.35 −2.21 −2.28 0.13 42.05 bd 0.02-5 −1.91 −2.51 −2.21 0.25 42.47 bd 0.02-6 −1.41 −1.69 −1.55 0.29 36.16 0.06 0.10-7 −2.49 −2.73 −2.61 0.20 46.37 0.18 0.08

δ57Fe values normalised against the IRMM-14 standard.

bd: below detection.


analyses. In two of these drill-cores, i.e. GL26 and BEL6B, the Hotazelrocks are devoid of any visible secondary effects such as supergeneoxidation or hydrothermal alteration controlled by faults andsupradjacent unconformity contacts with younger sedimentarysequences (Gutzmer and Beukes, 1995; Tsikos et al., 2003). Samplesfrom the upper half of drill-core BEL7A display some incipientalteration effects on the BIF which translate into limited martitisationof magnetite and calcite replacement by dolomite. The fourth drill-core intersection (N12D) captures a hydrothermally de-carbonatedand oxidised intersection of the Hotazel strata where iron in the BIF ispreserved exclusively in the form of secondary micro-platy hematiteacross the entire stratigraphic section sampled. Samples of the lattertwo drill-cores were analysed for iron isotopes with the addedobjective of monitoring any potential isotopic shifts related to latefluid processes.

Mineralogically representative slices of core samples were made bysectioning with a diamond saw along their length. Our samplingfocused particularly on the hematite–lutite transitions between the BIFand manganese layers, since these macroscopically display sharp redoxgradients in terms of the iron species present. The rock slices werecrushed and ground to fine powder in a corundum mortar. Dissolutionof the crushed rock powders was performed by digesting 0.05 g ofpowder with 5 ml 9.5 M HCl solution at 85 °C for 12hours. A portion ofthe solutions was also taken for chemical analysis of Fe, Mn, Ca, Mg,using ICP-OES. Since dissolution of Fe-silicate minerals in our rocks isnot possible through this method, the Fe isotope data reported in thisstudy are not strictly representative of whole-rock compositions butrather of the total oxide and carbonate component of our samples.Comparisons of the Fe content of Hotazel BIF samples as reported herewith previously published bulk-rock data (Tsikos et al., 2003), suggestthat the total fraction of Fe-bearing silicateminerals would contribute anegligible amount of Fe, and thus, to a close approximation, the HCldissolved fraction represents the bulk value of the rock.

The protocol for the preparation of pure Fe-bearing solutions foriron isotope analysis is modified after that used by Zhu et al. (2002)and is described in detail by Matthews et al. (2004). Appropriateamounts of solution to give a 10 μg/g Fe solution in 0.1 M HCl wereevaporated to dryness, re-dissolved in 7 M HCl and passed twicethrough an anion exchange resin for chromatographic separation of Fefrom other elements. Fe isotopic measurements were made on a NuInstruments MC-ICP-MS at the Geological Survey of Israel. All theanalyses reported here (Tables 1 and 2) are given as 57Fe/54Fe ratiosand reported in the δ57Fe notation relative to the Institute forReference Materials and Measurements (IRMM)-14 standard. Aprocessing bias of δ57Fe=0.05±0.09% (1 standard deviation) wasdetermined by running the IRMM standard solution through the samedouble chromatographic separation protocol used for samples.

Our machine is configured for δ57/54Fe measurements and not forδ56/54Fe measurements. Mass fractionation dependence of the isotopicmeasurementswas independently verified by running a set of syntheticand natural samples through two analytical cycles of δ57/54Fe andδ56/54Fe. Most sample solutions measured in this work were processedwith duplicate column chemistries, and the reproducibility of repli-cate MC-ICP-MS analyses are compatible with a 1 standard deviationvalue of ±0.1%, similar to that reported by Matthews et al. (2004). Ironisotope fractionations are given as Δ57Fe values, where:

Δ57FeðA –BÞ ¼ δ57FeðAÞ – δ57FeðBÞ:

Isotopic compositions reported in the literature using the δ56Fenotation (i.e., 56Fe/54Fe ratios) are corrected to δ57Fe values using themass-dependent fractionation relationship:

δ56Fe ¼ 0:68 ⁎ δ57Fe:

Please cite this article as: Tsikos, H., et al., Iron isotopes constrain biogeochemical redox cycling of iron and manganese in a Palaeoproterozoicstratified basin, Earth Planet. Sci. Lett. (2010), doi:10.1016/j.epsl.2010.07.032


Fig. 3. Distribution of δ57Fe values versus Mn/(Mn+Fe) ratios for unaltered BIF (drill-cores GL26, BEL6B and BEL7A), altered BIF (drill-core N12D), hematite lutite andmanganese ore samples from the Hotazel Formation.


5. Results

The range of δ57Fe values for all Hotazel BIF samples that showvery limited or no secondary alteration (i.e. those from drill-coresGL26, BEL6B and BEL7A) is between−2.74 and−0.07‰, with a meanat −1.70‰ (Fig. 3, Table 1). These values are significantly lower thanthose of the underlying BIF succession of the Transvaal Supergroup(Kuruman and Griquatown Formations) and the broadly coevalBrockman BIF of the Hamersley Basin, Australia (Pickard, 2003), theδ57Fe values of which rarely drop below−1.50‰ on amineral-specificlevel (Fig. 4; Johnson et al., 2003, 2008a). δ57Fe data of highly alteredHotazel BIF from drill-core N12D (Table 2) that contains iron entirelyas secondary hematite are similarly low (−1.98±0.31‰), indicatingthat iron isotopesmust have behaved conservatively during later fluidalteration.

The eleven samples collected from the transitional hematite lutitesand the two from the lowermost manganese layer of drill-core GL26,are characterised by δ57Fe values ranging from −2.42‰ to as low as−3.52‰ (Figs. 3 and 4). These are, to our knowledge, the lowest δ57Fevalues yet reported from Precambrian hematite-rich sedimentaryrocks. From a stratigraphic viewpoint, sharp declines in δ57Fe of 1–2‰are recorded across the hematite–lutite transitions relative to theadjacent BIF, in concert with corresponding increases in bulk Mn/Feratios (Fig. 5); very low δ57Fe values also occur in the two samples

Fig. 4. Plot of δ57Fe data for all samples from the Hotazel Formation against its maximum possvalues for other classic BIF sequences deposited during the 3.8–1.8 Ga period. Note that the dthose for Fe-oxide (magnetite, hematite) and siderite separates from these sequences. Data fBIF, Australia, from Johnson et al. (2008a); and Belingwe BIF, Zimbabwe, from Rouxel et alranges displayed in Fig. 5 of Johnson et al. (2008b) (see also references therein for specific


from the lowermostmanganese layer. In terms of ironmineralogy, theforegoing trends essentially correspond to the increase in theabundance of hematite at the expense of magnetite with distancefrom the BIF-lutite contacts, with samples from the manganese layerscontaining iron entirely as hematite.

6. Discussion

6.1. Redox cycling during early diagenesis

Interpretation of the iron isotope dataset hinges to a considerabledegree on previously reported carbon isotope data from across theentire Hotazel BIF-Mn succession (Tsikos et al., 2003). Ankerite andcalcite from unaltered Hotazel BIF samples exhibit low δ13C (PDB)values relative to typical marine bicarbonate, ranging from −4 to aslow as −18‰. Kutnahorite and manganoan calcite from the inter-bedded manganese layers have comparably low δ13C values, between−8 and−12‰. Similar signatures from other BIF sequences as well asfrom Phanerozoic manganese carbonate deposits, are cited asevidence for microbially mediated, sub-oxic diagenetic reduction ofhigher-valence iron and manganese oxy-hydroxide precursors re-spectively, coupled with oxidation of co-precipitated organic matter(e.g. Baur et al., 1985; Okita et al., 1988; Kaufman et al., 1990;Heimann et al., 2010). The magnitude of heavy carbon isotopedepletion in each instance can be attributed to the relative supply oforganic carbon versusmarine bicarbonate to the initial sediment, theirrespective end-member isotopic compositions, and the abundanceand speciation of high-valence Fe and Mn solid precursors as electronacceptors for anaerobic carbon cycling during diagenesis.

Mineralogical evidence suggests that reduction of Fe andMn in theHotazel sediments was both incomplete and strongly decoupled: BIFunits are dominated by mixtures of partly- to wholly reduced Feminerals such as magnetite, ankerite, and to a much lesser extent Fe-silicate minerals, and contain very low bulk Mn contents (Table 1),chiefly as Mn(II) in ankerite (Tsikos et al., 2003). By contrast, themanganese layers and adjacent transitional hematite lutites containmanganese entirely in the divalent and trivalent states in the formof kutnahorite Ca(Mn,Mg)(CO3)2, manganoan calcite and brauniteMn2+Mn3+

6SiO12, whilst Fe is present almost entirely in the trivalentstate (i.e. as hematite). The virtual absence of sulphide minerals fromboth the BIF and manganese-rich portions of the Hotazel Formation

ible age, along withmineral-specific δ57Fe data ranges (as grey bars) and selected meanata range for the ~2.5 Ga Kuruman–Griquatown–Brockman BIFs represents collectivelyor the Kuruman–Griquatown BIFs of South Africa from Johnson et al. (2003); Brockman. (2005). δ57Fe ranges for the Isua and Biwabik BIFs adapted from corresponding δ56Fedata sources).



Fig. 5. Stratigraphic profiles of δ57Fe and Mn/Fe ratios from three of the four selected drill-cores intersecting essentially unaltered Hotazel Formation in the northern Kalaharimanganese field (vertical scale only approximate). Mn-rich stratigraphic intervals sampled for δ57Fe analyses (i.e. transitional hematite lutites andmanganese ore where applicable)are highlighted by laterally extended shaded bands. Note that the stratigraphic extent of sampling was limited in each drill-core by the relative position of unconformity contactswith younger sedimentary sequences. mbs: meters below surface.


(Tsikos and Moore, 1997; Tsikos et al., 2003) also suggests thatsulphate reduction must have been limited in extent, either dueto very low availability of dissolved sulphate in the ambient waters orto the predominance of precursor Fe and Mn oxy-hydroxides overdissolved sulphate as the thermodynamically favoured electronacceptors (Froelich et al., 1979; Canfield et al., 2005). We thereforeassume that early diagenetic redox cycling in the Hotazel sedimentswas primarily controlled by the abundance of iron and manganeseoxy-hydroxide precipitates. We further assume that the preservedmineral assemblages do not record any major redox changesassociated with the post-diagenetic burial history of the Hotazeldeposits.

Previous studies on the iron isotope composition of BIFs focused onsmall-scale isotopic inhomogeneities on a band-to-band scale,recorded in most instances in individual mineral phases such asmagnetite and siderite (e.g. Johnson et al., 2003, 2008a; Heimannet al., 2010). Low iron isotope signatures in such minerals have beeninterpreted to result from dissimilatory iron reduction (DIR) of initial,high δ57Fe ferric precipitates (Bullen et al., 2001; Croal et al., 2004;Balci et al., 2006) via a variety of kinetic and/or equilibrium diageneticpathways (Johnson et al., 2005; Crosby et al., 2007; Johnson et al.,2008b). The consistently low δ57Fe values in almost all BIF samples ofthe Hotazel Formation represent essentially average values for totaloxide and carbonate across several individual Fe-rich and Fe-poorbands for each sample, and thus cannot alone provide sufficientconstraints as to the exact role of DIR in terms of Fe cycling duringdiagenesis of the Hotazel strata. The very low δ57Fe values of ourthirteen manganese-rich samples (Figs. 3 and 4), however, provideimportant insights. Here, the prevalence of hematite in co-existencewith aMn(II)-rich assemblage dominated by low-δ13CMn-carbonatesand braunite, suggests that the initial ferric component must haveremained largely inert in favour of thermodynamically preferred


bacterial reduction of co-precipitated Mn(III) and/or Mn(IV) solidprecursors. Such processes have been demonstrated in manganese-rich sediments from modern marine settings such as the Skagerrakdeep-site in the Baltic-North Sea transition (Jensen et al., 2003).Paucity of Fe cycling by DIR in the manganese-rich portions ofthe Hotazel would also imply that the lowest δ57Fe value of themanganese-rich Hotazel samples (i.e. −3.52‰) represents a maxi-mum value of the dissolved Fe(II) pool at the time of deposition. Thelatter argument is based on the added assumption that initial Fe(III)precipitation was accompanied by some fractionation of the heavy Feisotopes into the latter (Bullen et al., 2001; Croal et al., 2004; Balci etal., 2006; Johnson et al., 2008b). The relatively higher, yet still lowδ57Fe values of the Hotazel BIF samples can then be interpreted torecord essentially the same isotopic depletion in the ambient dis-solved Fe(II) pool, with variable positive departures from sample tosample likely to reflect variable degrees of DIR-driven loss of iso-topically light Fe(II) from the initial sediment.

6.2. Primary depositional environment

The Palaeoproterozoic snowball Earthmodel outlined in the earliersections of this paper (Kirschvink et al., 2000; Kopp et al., 2005;Kirschvink and Kopp, 2008), provides an elegant construct in whichthe Hotazel deposits are interpreted in terms of an extreme climaticevent, the timing of the GOE and the evolution of oxygenicphotosynthesis. With respect to its specific applicability to the Hotazelsedimentary processes, however, there are a number of points ofcontention:

• The snowball model assumes an age of approximately 2.2–2.3 Ga forthe Makganyene glaciation on the basis of stratigraphic relations inthe broader Transvaal basin and the age of the Ongeluk volcanic




rocks stratigraphically below the Hotazel strata and above theMakganyene diamictite (see Beukes and Gutzmer, 2008, andreferences therein). This invalidates without discussion the~2.4 Ga age of the Mooidraai Formation.

• The ~2.2 Ga age assumed for the Makganyene event requires a ca.200 Ma erosional unconformity between the Ghaap and Postmas-burg Groups in the Transvaal Supergroup. The existence of such ahiatus has recently been challenged on the basis of field evidence(Moore et al., 2001; Polteau et al., 2006).

• The snowball model does not adequately account for the highlyproductive, Mn-rich depositional environment of the HotazelFormation, with ca. 50% of the world's total land-based reserves ofMn being apparently sourced from a single submarine hydrother-mal system causally associated with the Ongeluk volcanic event.Indeed an explanation for the presence of huge amounts of Mn in aBIF-dominated sequence is generally called for in light of the verylow levels of manganese in Archaean and Palaeoproterozoic BIFsequences world-wide (Klein, 2005).

• The snowball model requires that deposition of manganese tookplace in an oxygenated environment in the form of primary Mn (IV)oxy-hydroxides, despite the fact that the mineralogy of themanganese layers provides no direct evidence for such precursors.

• Finally, the snowball model does not account for the cyclicintercalation between BIF andMn-rich sediment, whereby transientevents of manganese-rich deposition took place in a stratifiedenvironment that otherwise also favoured deposition of Mn-poorBIF.

Our iron isotope data allow us to gain new insight into the aboveissues. Primarily, our interpretation that the depletion in heavy ironisotopes directly reflects the ambient dissolved Fe(II) pool precludes adirect hydrothermal source for Fe and Mn associated with theOngeluk volcanic event. We propose instead that the low δ57Fe valuesrepresent an evolved isotopic signature, which would have beeninherited from long-term Rayleigh fractionation effects due tosequestration of isotopically heavy Fe as BIF in the lower stratigraphicportions of the Transvaal Supergroup. This suggestion shares parallels,albeit on a spatially and temporally more restricted scale, with themodel of Rouxel et al. (2005) for the Precambrian oceans during theworld-wide peak in BIF deposition between 2.7 and 2.3 Ga. Wespecifically envisage that the palaeo-depositional environment of theHotazel Formation would have evolved from a much larger stratifiedbasin characterized by the precipitation of typical, Mn-poor BIF, asmanifested in the Kuruman and Griquatown BIFs and possibly theBrockman BIF of the Hamersley Basin in Australia. The inclusion of theBrockman BIF is based on the widely-held view that the Kaapvaal andPilbara cratons constituted a common geotectonic entity during BIFdeposition (Nelson et al., 1999; Pickard, 2003; Beukes and Gutzmer,2008; de Kock et al., 2009).

In order to quantify our argument for a Rayleigh isotopic process,we assume a starting δ57Fe value of 0‰ for dissolved Fe(II) in thebottom anoxic waters of the Transvaal proto-basin, as conventionallyused in the literature for the Archaean ocean inventory (Johnsonet al., 2008a,b). An additional parameter in our considerations isthe initial and/or long-term hydrothermal flux of Fe and Mn intothe original basin. In the absence of independent constraints forthe Fe/Mn ratio in submarine hydrothermal fluids of the Precam-brian, we assume that a maximum value reflecting that of averageoceanic crust (i.e., ~50/1; Qin and Humayun, 2008) would representthe most conservative estimate for this ratio. Low sulphate content insuch fluids in a pre-GOE ocean would also have ensured minimaltitration of Fe as sulphide (Kump and Seyfried, 2005), and thereforeminimal fractionation of Fe relative to Mn prior to hydrothermaldischarge.

Published bulk geochemical analyses from the Kuruman, Griqua-town and Brockman BIFs (Klein and Beukes, 1989; Beukes et al., 1990;


Kaufman et al., 1990) suggest that Fe and Mn were sequestered at aFe/Mn mass ratio at least twice that of the assumed averagesubmarine hydrothermal flux (i.e. ≥100/1). This high ratio can beattributed to a combination of several factors depending on theprimary oxidative mechanism. Oxidation rates in aerobic environ-ments at pHb8 are known to be faster for Fe(II) relative to Mn(II),whereas more rapid photo-autotrophic bacterial oxidation of Fe(II)relative to Mn(II) may have been an alternative mechanism in theabsence of oxygenic photosynthesis (Canfield et al., 2005). In eithercase, any oxidation of Mn(II) would have been expected to take placeas a two-step electron transfer with Mn(III) as the intermediate step(Tebo et al., 2005; Webb et al., 2005). Mn(III) however, would also bea thermodynamically favourable electron acceptor for the oxidation ofthe abundant co-existing Fe(II) in primary BIF basins (Canfield et al.,2005):

MnOOH þ Fe2þ→FeOOHþMn

2þ

resulting in effective re-cycling of Mn back into the water column asMn(II). This latter pathway indeed may have been a key factor for thevery low levels of manganese observed inmost Palaeoproterozoic BIFsof South Africa and world-wide (Klein, 2005).

Based on the above parameters, net removal of isotopically heavyFe as BIF relative to the initial and/or long-term hydrothermal flux inthe Transvaal proto-basinwould have resulted both in the progressiveformation of isotopically light Fe(II) in the residual dissolved pool anda relative enrichment in dissolved Mn(II). In the Rayleigh fraction-ation model of Fig. 6, we assume that the long-term kinetic Fe isotopefractionation between BIF precursor sediment and the total dissolvedFe(II) pool:

Δ57Fe ¼ δ57FeðBIFÞ−δ57FeðaqÞ;

was a function of the combined fractionation effects of oxidation ofFe(II)aq to particulate Fe(III) species and partial reduction of thelatter into magnetite and lesser Fe carbonate via DIR (Bullen et al.,2001; Croal et al., 2004; Balci et al., 2006; Crosby et al., 2007; Johnsonet al., 2008b). Considering that oxides, and particularly magnetite,are the dominant Feminerals in BIFs on regional scales, we hypothesizethat both above processes would have cumulatively resulted in apositive net Δ value. Fig. 6 demonstrates that the low δ

57Fe values

recorded in the elevenMn-rich hematite lutites and the twomanganeseore samples of the Hotazel Formation could have ensued in a terminaldepositional environment that evolved from the precipitation of morethan 80% of the total dissolved Fe(II) inventory as BIF, at an average Δvalueof amagnitudebetween0.8 and1.2‰. This range falls between theexperimentally measured fractionation factors (as δ57Fe) of 2.0‰ formagnetite-Fe(II) and 0‰ for siderite-Fe(II), respectively (Johnson et al.,2005). The resulting Fe/Mn ratio in the evolved solution is broadlyconstrained to a value of 15 or less (Fig. 6), which would have beenconducive to the formation of a Hotazel-like BIF-Mn sequence uponquantitative cyclic precipitation.

7. Conclusions — implications for palaeoenvironmental models

The key elements of our interpretation are summarised in the two-stage ocean model of Fig. 7 for the Transvaal/Hamersley proto-basin.Initial BIF deposition (Fig. 7A) would have taken place in a stratifiedbasin with bottom waters enriched in dissolved Fe(II); an Fe/Mn ratiobroadly recording that of average oceanic crust (i.e. ~50/1); and anear-zero δ

57Fe composition reflecting the long-term input of

dissolved iron from submarine hydrothermal activity and/or riverineinput. In such a setting, typical Mn-poor BIF deposition (Kuruman/Griquatown/Brockman BIFs) would have been driven during periodicupwelling/mixing events, by either aerobic or anaerobic Fe(II)oxidation mechanisms, whilst long-term Mn re-cycling in the water



Fig. 6. Model calculations showing the variation of δ57Fe in BIF as a function of : (1) F(the fraction of remaining aqueous Fe reservoir); and, (2) the Fe/Mn mass ratio in theremaining solution, assuming that both metals are removed at an average rate of 100/1from an initial reservoir of average oceanic crust signature (50/1). The calculated curvesare made for variable values of the mean, long-term kinetic fractionation factor:Δ57Fe=δ57Fe(BIF)–δ57Fe(II)aq, using Rayleigh equations that indicate the Fe isotopiccomposition of the ferrous solution reservoir at any value of F. Data points for thethirteen hematite lutite and Mn ore samples of the Hotazel Formation represent δ57Fevalues plotted against values of F, the latter calculated from the bulkMn and Fe contentsof the samples (see also Table 1). Note the clustering of data at values of Fb0.2,corresponding Fe/Mnmass ratios b15, and Δ values broadly between 0.8 and 1.2‰ (seetext for detailed discussion).

Fig. 7. Diagrams illustrating key chemical and isotopic parameters of (A) an initial stratiTransvaal–Hamersley Supergroups; and, (B) an evolved stratified basin conducive to the deppool in terms of δ57Fe (≤−3.5‰) and relative abundance of the two metals (Fe/Mn≤15/1),Formation, and the Rayleigh fractionation model of Fig. 6, respectively (see text for details)



column would have been achieved through effective reduction ofdissolvedMn(III) species or transient Mn(III)/Mn(IV) solid precursorsby co-existing aqueous Fe(II). Rayleigh fractionation processes relatedto protracted deposition of oxide-dominated BIF in such an environ-ment, would have led to the terminal depositional basin for theHotazel BIF-Mn sequence (Fig. 7B), characterized by an enrichmentin dissolved manganese in the bottom anoxic waters reflected in aFe/Mn solution ratio of at least 15/1 (Fig. 6), and an evolved, depletedδ

57Fe signature of dissolved Fe(II) with a maximum value of −3.5‰.

The latter signature is quantitatively reflected in the very low ironisotope signatures of the Mn-rich, hematitic portions of the HotazelFormation, where iron cycling via DIR is interpreted to have beencurtailed in favour of bacterial manganese reduction during earlydiagenesis.

The model presented in Fig. 7 provides a simplistic re-constructionof what must have been a chemically and biologically very complexpalaeo-depositional environment. Previous attempts to model thedeposition of the Hotazel sequence have invoked the superposition oftransgression-regression cycles in order to explain the cyclicintercalation of manganese-rich sediment with typical BIF (Beukes,1983; Tsikos and Moore, 1997). Such models remain to be tested inthe absence of any direct constraints available to date on themechanism of primary Fe and Mn precipitation in the Hotazelpalaeodepository, and on the speciation of the Fe and Mn precursors.

The suggestion that the Hotazel Fe–Mn Formation represents anend-member sedimentary component in a common, evolving BIFbasin removes the requirement for a ca. 200 Ma depositional breakbetween the Ghaap and Postmasburg Groups in the TransvaalSupergroup, and thus implies a pre-2.4 Ga age for the Hotazelsuccession as constrained by the overlying Mooidraai carbonates(Bau et al., 1999). This contention calls for a re-appraisal ofstratigraphic correlations in the Transvaal Supergroup, particularlywith regard to the absolute timing of the Makganyene glaciation and

fied basin for the deposition of the Kuruman/Griquatown and Brockman BIFs in theosition of a Hotazel-like BIF–Mn sequence. The evolved character of the dissolved Fe–Mnis based on the lowest δ57Fe value in the hematite lutite/Mn ore samples of the Hotazel.




Ongeluk volcanism at the centre of an essentially continuous mega-cycle of biochemical deposition (Moore et al., 2001; Polteau et al.,2006). Furthermore, the Hotazel sedimentary environment is nolonger considered in isolation from the underlying BIF sequence of theTransvaal Supergroup, nor does it require a geochemically andisotopically exceptional hydrothermal metal source in the context ofBIF genesis in general.

The interpreted end-member character of the Hotazel deposits inan evolving stratified basin also suggests that manganese may havebeen overlooked as an important proxy of biogeochemical redoxprocesses in the Neoarchaean–Palaeoproterozoic atmosphere–oceansystem. We consider that the very low manganese levels recorded inmost classic BIF sequences of the world could themselves constitutean important constraint in understanding primary Fe precipitationmechanisms in stratified basins of that time. The tentative placementof the Hotazel depositional event at ~2.4 Ga, may also be of directsignificance with regard to the timing of the GOE: until now, it hasbeen argued that the Hotazel deposits represent evidence for aerobicoxidation of Mn and Fe in a post-GOE world. Until the specificmechanisms and absolute timing of primary Mn and Fe precipitationin the Hotazel strata are adequately resolved, we present analternative interpretation whereby the Hotazel Fe–Mn formationmay in fact represent a precursor to, rather than a consequence of, theGOE: specifically, if it is assumed that oxygenic photosynthesis hadindeed commenced by ca. 2.7 Ga ago, the large-scale deposition of Feas BIF in sequences such as those of the Transvaal–Hamersley basinmay have acted as an important sink for photosynthetic oxygen in theatmosphere–ocean system. In such a context, extensive manganeseprecipitation during the latter stages of basin evolution may similarlybe regarded as one of the terminal marine oxygen sinks of that period,and thus a strong signal for the onset of the GOE thereafter. Furtherdetailed research on the remarkable Hotazel deposits with emphasison the rigorous re-construction of their depositional environment,will hopefully assist in the resolution of this enigma.

Acknowledgements

The iron isotope measurements were funded by grants no2002-338 from the US–Israel Binational Science Foundation (BSF)and 1102/07 from the Israel Science Foundation (ISF). The technicalhelp of Mr Uri Ryb is greatly appreciated. HT acknowledges ASSMANGLtd and specifically Mr B. Ruziwe, for permitting unconstrained accessto drill-core material. AM acknowledges receipt of the Raymond F.Kravis Chair in Geology. Reviews by S. Severmann, N. Planavsky andtwo anonymous reviewers greatly improved the earlier versions ofour manuscript.

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Iron isotopes constrain biogeochemical redox cycling of iron and manganese in a Palaeoproterozoic...

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