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Precambrian Research 226 (2013) 116– 124

Contents lists available at SciVerse ScienceDirect

Precambrian Research

journa l h o me pa g e: www.elsev ier .com/ locate /precamres

iliciclastic associated banded iron formation from the 3.2 Ga Moodies Group,arberton Greenstone Belt, South Africa

omaso R.R. Bontognali a,∗, Woodward W. Fischerb, Karl B. Föllmic

ETH-Zurich, Geological Institute, Zurich, SwitzerlandCalifornia Institute of Technology, Geological and Planetary Sciences, Pasadena, CA, United StatesUniversity of Lausanne, Institute of Earth Sciences, Lausanne, Switzerland

r t i c l e i n f o

rticle history:eceived 29 June 2012eceived in revised form0 November 2012ccepted 19 December 2012vailable online xxx

eywords:anded iron formationhert

ron cyclearly lifearberton Greenstone Belt

a b s t r a c t

Most models proposed for banded iron formation (BIF) deposition are based on observations of well-preserved Late Archean and Paleoproterozoic BIF. Efforts to push the understanding gained from youngersuccessions deeper in time have been hampered by the high metamorphic grades that characterize EarlyArchean BIF. This study focuses on a unique occurrence of well-preserved and contextualized BIF fromthe Early Archean (∼3.2 Ga) Moodies Group, in the Barberton Greenstone Belt, South Africa. The MoodiesBIF occurs thinly interbedded with fine-grained and cross-stratified sandstones, indicating depositionduring times of decreased clastic sediment supply. In the Moodies BIF, chert is present as concretions,and is never observed in direct contact with the siliciclastic material but is always associated with ironminerals. This observation suggests that the processes leading to the formation of both chert and ironminerals were coupled. The dominant iron-rich minerals within unweathered Moodies BIF are hematiteand magnetite, with less common occurrences of Fe–carbonate phases (mainly ankerite). Petrographictextures reveal that hematite constitutes an early mineral phase, while magnetite and ankerite displaytextures indicative of a late diagenetic or metamorphic origin. Carbonaceous particles are present in closeassociation with the magnetite crystals. These C-bearing phases may be the preserved organic matter ofmicrobes involved in the production of the ferric iron precursor phases, though it is difficult to rule out

an origin from abiotic processes involving thermal decomposition of siderite to magnetite and organiccarbon compounds. Nonetheless, the range of textures, mineralogies, and valence states supports the viewthat diagenetically-stabilized BIF mineralogies reflect the interaction of ferric iron phases with reducingfluids during diagenesis. These patterns are commonly observed in younger Archean and Paleoproterozoiciron formations, and imply a continuity of processes operating in the iron and silica cycles across both arange of paleoenvironments and long intervals of Archean time.

© 2013 Elsevier B.V. All rights reserved.

. Introduction

Banded iron formations (BIF) are chemical sedimentary rocksharacterized by alternating layers of Fe-rich minerals and chertmicrocrystalline quartz) (James, 1954). Despite years of assiduousesearch, several aspects concerning their genesis remain contro-ersial (Bekker et al., 2010; Beukes and Gutzmer, 2008; Clout andimonson, 2005; Klein, 2005; Trendall, 2002). BIF are widespreadn Archean and Paleoproterozoic sedimentary basins, but similaracies is not observed to form in any modern geological setting.

IF clearly result from a suite of non-uniform processes. Secularhanges in their accumulation and sedimentary style continue tootivate efforts to understand their origins, with the goal of being

∗ Corresponding author.E-mail address: tomaso.bontognali@erdw.ethz.ch (T.R.R. Bontognali).

301-9268/$ – see front matter © 2013 Elsevier B.V. All rights reserved.ttp://dx.doi.org/10.1016/j.precamres.2012.12.003

able to link their occurrences to changes in fluid Earth redox chem-istry and geobiology.

It is commonly thought that, during times of BIF formation,ocean basins must have been anoxic and sulfur poor (at least atdepth) in order to allow for the transport and accumulation ofdissolved Fe(II); and that Fe was subsequently concentrated inthe sediments by oxidation, hydration, and precipitation (Canfield,1998; Cloud, 1968; Drever, 1974; Holland, 1973; Klein, 2005). Fe(II)may have been oxidized in the water column forming a hydrousferric oxide phase as a precursor to hematite (Bekker et al., 2010;Lepp and Goldich, 1964). Oxidation may have occurred either inthe presence of O2 produced by photosynthetic organisms or in theabsence of molecular oxygen, through abiotic photochemical reac-

tions (Cairns-Smith, 1978) or through anoxygenic photosynthesiswith iron as a primary electron donor (Widdel et al., 1993). Alterna-tively, direct precipitation from anoxic seawater may have formedsiderite and mixed valence iron–silicate phases.

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sandstones, mudstones, minor conglomerates, and discrete inter-calations of BIF (Anhaeusser, 1973; Eriksson, 1977, 1979, 1983;Heubeck and Lowe, 1994, 1999), on which this study focused. Themetamorphic grade in the Moodies Group in the study locality

Barberton

South Africa

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Moodies Group

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Fig Tree Group

Study Area

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T.R.R. Bontognali et al. / Precam

Although now common, the hypotheses that microbes werenvolved in the primary oxidation of Fe(II) to Fe(III) (via oxygenicr anoxygenic photosynthesis, or via chemoautotrophy at low oxy-en concentrations (Brown et al., 1995; Cloud, 1973; Harder, 1919;artman, 1984; Kappler et al., 2005; Konhauser et al., 2002; Perryt al., 1973; Posth et al., 2008)) contrast with the general lack oficrobial biomass (e.g. accumulation of organic carbon, microfos-

ils or biomarkers) within BIF (Beukes and Klein, 1992; Klein andeukes, 1989). A reasonable explanation for this discordance is pro-ided by diagenetic processes that respired much of the organicarbon back to dissolved inorganic carbon (DIC) during interactionsith ferric oxide or mixed valence phases (Baur et al., 1985; Fischer

nd Knoll, 2009; Konhauser et al., 2005; Perry et al., 1973; Walker,984). This scenario is consistent with the well-documented pres-nce, in many BIF, of diagenetic iron-bearing carbonates (sideritend ankerite) characterized by a 13C-depleted isotopic compositionBaur et al., 1985; Becker and Clayton, 1972; Beukes et al., 1990;ischer and Knoll, 2009; Goodwin et al., 1976; Kaufman et al., 1990;erry et al., 1973).

The origin of chert – the most abundant phase in BIF – in theseocks is no less enigmatic than that of iron. In the absence of silicify-ng organisms, Precambrian oceans were likely close to saturation

ith respect to amorphous silica and evaporation may have pro-ided an important driver for the precipitation of chert (Siever,992; Trendall and Blockley, 1970). However, this interpretationoes not explain why chert is common in BIF, which are com-only manifest as a deep-water facies. One hypothesis to explain

he transport and precipitation of silica in deep waters, as wells its close association with iron minerals, has been proposed byischer and Knoll (2009). This mechanism is based on the tendencyf ferric hydroxides to bind and shuttle silica to basinal waters andediments. Fe(III) respiration taking place within sediments wouldhen return the majority of iron to the water column, while silica,hich does not undergo reductive dissolution, remains reactive,

s concentrated in pore waters, and is ultimately precipitated asiagenetic mineral phases.

Finally, not only is the origin of the BIF mineralogy contro-ersial, but also the processes resulting in the interlaminationetween the iron-rich and chert-rich beds producing three differ-nt scales of banding – microbands (≤1 mm), mesobands (∼1 mmo 10 cm) and macrobands (≥1 m) (Trendall and Blockley, 1970).roposed explanations include temperature variations (Posth et al.,008), microbial blooms (Trendall and Blockley, 1970), episodiccean mixing (Hamade et al., 2003), deposition by density currentsKrapez et al., 2003), and internal dynamics of the geochemicalystem (Wang et al., 2009).

The abovementioned concepts, models, and hypotheses wererimarily developed from observations of Late Archean and Paleo-roterozoic BIF, including the spectacular craton-wide occurrences

n the Hamersley Group of Western Australia, the Transvaal Super-roup in South Africa, and in the Lake Superior Region in the USABekker et al., 2010). Many of these successions were only mildlyffected by metamorphism and deformation, and these depositsffer good sedimentological context into their paleodepositionalettings and appreciable textural preservation such that informa-ion regarding their petrogenesis can be obtained. However, thextent to which models proposed for these younger successions cane tested and applied for interpreting BIF deposited during Earlyrchean time remains unclear. Many of these deposits are thin com-onents of severely deformed and metamorphosed (to greenschistnd granulite grade) Greenstone belt successions (Bekker et al.,010). Indeed, it is challenging to ascertain whether some of the

ldest iron-rich metamorphic rocks even were once BIF, or even ifhey had a sedimentary origin (Dauphas et al., 2007; Eiler, 2007).

oreover, younger and well-studied BIF show local differences inedimentological textures and mineralogies (Bekker et al., 2010),

Research 226 (2013) 116– 124 117

suggesting that a global theory explaining all BIF occurrences maynot exist. To answer the question of whether current models canbe extrapolated back in time to explain these Early Archean BIF it isimportant to identify well-preserved examples of earlier ArcheanBIF, which can be compared in terms of their sedimentary geology,geochemistry, and petrography, with their younger equivalents.

This study focuses on the BIF from the ∼3.2 Ga Moodies Groupfrom the Barberton Greenstone Belt, South Africa. These sedimen-tary rocks have been noted (Eriksson, 1977, 1983; Heubeck andLowe, 1999), but not studied in detail because the few outcropswhere they are exposed at the surface are strongly affected by surfi-cial weathering obscuring the original mineralogies. For this study,we were able to collect a suite of samples directly from the under-ground tunnels of an active gold mine. Coupled to observationsfrom an outcrop located at the surface, these materials provide aunique window into the processes responsible for the depositionof BIF in Early Archean time.

2. Geological setting

The Barberton Greenstone Belt (BGB) is situated in thecentral-east part of South Africa, along the border between theMpumalanga Province and Swaziland (Fig. 1). The BGB contains adiverse suite of sedimentary strata deposited in one of the oldestrecognized foreland basins; despite their early Archean age, regionsof the BGB have remarkably good preservation and provide a uniqueand rich source of insight about sedimentary processes and envi-ronments on the early Earth (Byerly et al., 1986; Eriksson, 1977;Eriksson and Simpson, 2000; Javaux et al., 2010; Noffke et al., 2006;Simpson et al., 2012). The successions of rocks that comprise theBGB were subdivided into three different groups (Hall, 1918; Loweet al., 1999) (Fig. 2). The Onverwacht Group (3.5–3.3 Ga) is pre-dominantly composed of mafic and ultramafic volcanic rocks but italso includes some thin cherty units thought to be sedimentary inorigin (Lowe et al., 1999). The overlying Fig Tree Group (3.3–3.2 Ga)consists mainly of fine-grained sedimentary rocks including BIF,carbonaceous shales, siltstones, sandstones, and cherts. And finallythe Moodies Group (3.2–3.1 Ga) includes alluvial to shallow-marine

Fig. 1. Study area of the 3.2 Ga Moodies Group, Barberton Greenstone Belt, SouthAfrica.

Modified from Heubeck and Lowe (1994).

118 T.R.R. Bontognali et al. / Precambrian

Fig. 2. General stratigraphy of the Barbertone Greenstone Belt within the MoodiesH

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ills Block.

odified from Lowe et al. (1999).

s constrained to lower greenschist facies (Heubeck and Lowe,999), with estimated maximum burial temperatures of ∼220 ◦CDe Ronde et al., 1997; Toulkeridis et al., 1998). The minimum ageonstraints on the Moodies strata are provided by several igneousntrusions, and the maximum age by dates of detrital grains withinhe strata (Kamo and Davis, 1994; Kröner et al., 1991). In spite of theight structural folding that affected the region, sedimentary struc-ures are preserved unmodified and are widespread within the

oodies lithologies. These sedimentary structures and associatedacies reveal a wide range of terrestrial and marine sedimentaryaleoenvironments including aeolian, alluvial fan, braided stream,ide-dominated delta, open marine shelf that were distributedcross the foreland basin (Anhaeusser, 1973; Eriksson, 1977, 1979,983; Eriksson and Simpson, 2000; Eriksson et al., 2006; Heubeck,009; Heubeck and Lowe, 1994, 1999; Simpson et al., 2012).

. Locations, sampling, and methods

A detailed stratigraphic section was measured at bed-by-bedesolution (∼0.5 cm) and samples were collected from an outcropocated along Ameide Road (S 25◦49′56′′–E 031◦00′50′′, in a struc-ural sub-basin referred to as the Moodies Hill Block; Heubecknd Lowe, 1994) and directly from the underground tunnels of thegness Gold Mine (Ramp West, Access South, −600 m). Access toubsurface materials is particularly important to obtain pristineamples and mitigate the ever-present effects of oxidative sur-ace weathering in South Africa. Samples collected from the mineere not affected by weathering and show good preservation of

edimentary fabrics and textures. These materials were used foretrographic and geochemical analyses. The total amount of sam-

les collected was approximately 175 kg. Transmitted and reflected

ight microscopy was performed on 30 �m-thick polished thin sec-ions. Bulk mineralogy was determined by X-ray diffraction using

Scintag XRD 2000 diffractometer. Scanning electron microscopy

Research 226 (2013) 116– 124

(SEM) analyses were performed with a Philips XL-30 FEG equippedwith an EDAX energy dispersive X-ray spectrometer (EDX). Imagesand EDX analysis were obtained with a backscatter detector, anaccelerating voltage of 25 kV, and a working distance of 10 mm.The thin sections were coated with 7 nm of Au prior to analysis toensure a proper conducting sample surface.

4. Results

4.1. Stratigraphy and sedimentology

The stratigraphic section through the Moodies Group exposedin the Moodies Hills Block (Heubeck and Lowe, 1994) in out-crop along Ameide Road contains a BIF-rich interval within asuccession of shallow marine sandstones (Fig. 3). The section is62 m thick and is composed of five different lithofacies, includ-ing: (1) siliciclastic beds, (2) cm-scale packages of closely packedsub-mm scale laminations of iron minerals, (3) chert beds andconcretions, (4) non-laminated sandstone, and (5) volcanic tuff.Fine-grained sandstones, showing common small-scale parallel orwavy laminations with occasional trough cross-stratification dom-inate the base of the section. Continuing upward, the first BIFiron-rich packages begin to appear, interbedded with cm-scalewavy-laminated sandstone beds. Initially, these iron-rich pack-ages are not associated with conspicuous chert. The frequency ofiron packages (relative to sandstone interbeds) gradually increasesrelative to the amount of accumulated siliciclastic sediment andthe first chert concretions appear. Interestingly, chert beds andconcretions are always preceded and followed by layers of ironminerals and do not occur in direct contact with the siliciclasticbeds. The chert concretions are often characterized by lenticularshapes and their thickness is laterally heterogeneous – texturesseen in cherts of all ages and commonly in younger iron for-mation deposits (Beukes, 1984; Fischer and Knoll, 2009; Krapezet al., 2003). The alternation of sandstone and BIF lithologies isabruptly interrupted at 10.5 m by a 0.6 m-thick bed of volcanictuff. Gradually the thickness of the chert and the iron packagesincreases upsection while the frequency of interbedded sandstonebeds diminishes into an interval several meters thick of predom-inantly BIF with occasional sandstone beds. Where the bandingtypical of BIF is best expressed, the siliciclastic interbeds are absent.Continuing upward, the thickness and the frequency of iron andchert beds gradually decrease and siliciclastic sediment again dom-inates the section. These upper sandstones also contain commonsedimentary structures, including wavy laminations, trough cross-stratification, and soft-sediment deformation. In the uppermostpart of the studied section, two coarsening-upward sequenceswere apparent. These intervals consist of massively beddedsandstones.

4.2. Iron-bearing minerals in Moodies BIF

In unweathered samples, the iron-rich component of MoodiesBIF appears dark gray, shows a typical metallic luster, and is char-acterized by several scales of banding (Fig. 4). The thickness of theiron mineral laminations are commonly submillimetric but groupinto rough bundles at centimetric scales (an approximate averageis 0.5–1.5 cm, the thickest observed interval is about 3.5 cm). Inthe nomenclature of Trendall and Blockley (1970), these would bedescribed as microbands (millimeter to sub-millimeter units) andmesobands (centimeter thick units), respectively. The iron-rich

constituent of the Moodies BIF is always finely laminated, and con-tains no macroscopic granular, globular or oolitic morphologies likethose commonly observed in Paleoproterozoic-age granular ironformations (Fig. 5). The iron-rich packages are mainly composed

T.R.R. Bontognali et al. / Precambrian Research 226 (2013) 116– 124 119

Fig. 3. Stratigraphic section of Moodies Group strata exposed at the Ameide Road outcrop. Iron-rich packages and chert beds are thinly interbedded with siltstones ands horizt ing th

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andstones, very fine to fine in grain size. Chert is always associated with iron-richhe processes leading to chert accumulation were tightly coupled with those involv

f hematite and magnetite crystals with a microcrystalline quartzement (Fig. 6). The presence of these two iron-bearing phases isasily recognized by optical microscopy, but was further confirmedy XRD-analysis of bulk rock powders. The laminations are oftenmphasized by magnetite and fine-grained hematite intercalatedith thin laminae and discontinuous lenses of detrital siliciclasticarticles (Figs. 5 and 6B). Where iron minerals are predominantith little intervening clastic material, the lamination is often lessell expressed. Texturally, magnetite occurs as a diagenetic phase

hat primarily replaced or overgrew prior hematite – a featureommonly observed in younger iron formations (Beukes et al.,990; Fischer and Knoll, 2009; Han, 1978). Euhedral crystals ofagnetite overgrowing hematite (with relic hematite still visible

ons and does not occur in direct contact with siliciclastic sediment, indicating thate precipitation of the iron minerals.

in the middle) were observed. The abundance and petrographicalrelationships of magnetite are closely linked to the accumulationof fine-grained hematite. Some chert lenses include undulated, dis-continuous laminations almost exclusively comprised of hematite(i.e. without magnetite overgrowths) (Fig. 6C); the occlusion ofporosity in these cherty beds may have better preserved hematiteby allowing limited permeability of later reducing fluids. Underlow magnification, hematite appears as a gray homogeneouspowder. SEM microscopy reveals that the size of single hematite

spheroids is often less than 1 �m. Sulfide-bearing minerals,including bravoite, pyrite, and chalcopyrite, are also presentwithin the iron mineral-rich lamina. They are significantly lessabundant relative to hematite and magnetite and they often

120 T.R.R. Bontognali et al. / Precambrian Research 226 (2013) 116– 124

Fig. 4. Photograph of a polished slab of Moodies BIF. (A) Chert concretion. (B) Sub-mm scale laminations of iron-bearing minerals (mainly magnetite and hematite).(C) Siliciclastic bed. Chert concretions variably show mm to sub mm-scale internalllt

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Fig. 5. Large-scale photograph of a representative Moodies BIF polished sample. (A)Contact between two chert beds. The color of the chert varies from dark red to almosttransparent depending on the concentration of microscopic iron minerals (primar-ily hematite). (B) Iron-rich microbands finely interlaminated with siliciclastic grains(mainly fine quartz sand). (C) Siliciclastic bed showing laminations comprised of

often associated with magnetite crystals. Their size is rather regular(∼10 �m) and their distribution within the iron-rich layers appearshomogeneous: no clear laminations or morphologies resemblingmicrofossils were observed.

Table 1Elemental composition of the carbonaceous particles imaged in Fig. 7 from energydispersive X-ray spectroscopy.

Element Fig. 7Awt%

Fig. 7Cwt%

C 89.51 67.33O – 12.94Mg 0.16 0.05Al 0.01 0.04Si 0.55 0.86

aminations due to the variable amount of incorporated iron minerals. Iron-richaminations are due to both the fine interlamination with siliciclastic material andhe variable amount of surrounding chert.

ccur in association with small veins and cracks that cut primaryaminations – mineralization related to later metasomatic events.

.3. Chert in Moodies BIF

In unweathered samples, chert concretions appear red (Fig. 4).icroscopy reveals that the intensity of red color is related to

he concentration of fine-grained hematite laminations within thehert (Figs. 4 and 5A and C). In the literature, hematitic chert isometimes referred to as jasper. The thickness of the chert laminaearies from submillimetric to centimetric, the thickest observednterval measures about 7 cm. The chert beds often show a nodu-ar shape absent in iron-rich and clastic-rich layers. In some cases,mall (<1 mm) chert concretions are embedded in packages ofron minerals. In addition to hematite, domains of iron–carbonatemainly ankerite) and magnetite crystals are present within thehert. The growth of the carbonate crystals cut and overprint thehert laminations (Fig. 6C). A later weak overgrowth of chert overhe edges of the same carbonate crystals was also observed.

.4. Carbonaceous particles in Moodies BIF

SEM observations revealed the presence of carbonaceous parti-les within the Moodies BIF (Fig. 7). Carbon content of the particlesas measured by energy dispersive X-ray spectroscopy and it rep-

esents more than the 90% of their total weight (Table 1). These

sand draped by iron-bearing minerals. (D) Chert bed including discontinuous lam-inae comprised of iron minerals. Boxes marked A–D correspond to the microscopyimages presented in Fig. 6.

particles preferentially occur within the iron mineral-rich bands,

S 0.35 0.31Ca 0.06 0.08Mn 0.08 0.05Fe 4.49 3.02

T.R.R. Bontognali et al. / Precambrian Research 226 (2013) 116– 124 121

Fig. 6. Transmitted light photomicrographs of centimeter-scale Moodies BIF mesobands. (A) Chert concretions include fine-grained hematite spheroids (gray dots). Thelarger concentration of hematite in the upper band is responsible for the darker color of the chert and its laminated habit. Although less abundant, hematite is also present inthe lower band, which includes lenticular concretions rich in hematite (e.g. black arrow). Magnetite occurs as large opaque crystals with euhedral shapes developed along,but clearly crosscutting primary laminations (e.g. white arrow). (B) Iron-rich laminations associated with quartz sand. Coarse euhedral magnetite crystals (e.g. black arrow)overgrowing hematite. At this magnification, micron-sized hematite crystals appear like a gray powder, barely visible behind the euhedral, larger, and darker magnetitecrystals. Siliciclastic particles are clear (e.g. white arrow) and have diameters in thin section up to 150 �m (i.e. fine sand), and were likely transported as a part of bed loads zed bya ); thea w).

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ediment flux. (C) Interbedded sandstone beds show internal laminations, emphasi chert concretion. Crystals of ankerite are visible within the chert (e.g. white arrowlso includes discontinuous laminations comprised of iron minerals (e.g. black arro

. Interpretations

.1. Depositional setting of the Moodies BIF

BIF deposition constitutes a non-uniformitarian suite of pro-esses, but the sedimentary rocks in contact with these lithologiesrovide substantial insight into the environments of their deposi-ion (Fischer and Knoll, 2009; Ojakangas, 1983; Simonson, 1985).bservations of BIF deposits hosted in siliciclastic sequences high-

ight that BIF formed in a large variety of geological settingsnd paleodepths, ranging from abyssal fans to shallow marinenvironments (Bleeker et al., 1999; Eriksson, 1983; Fralick andufahl, 2006; Hofmann and Kusky, 2004; Srinivasan and Ojakangas,986). The siliciclastics beds interbedded with BIF lithologies inhe Moodies Group carry sedimentary structures (i.e. trough cross-tratification, wavy laminations, and graded siliciclastic beds tooarse sandstone) that were interpreted by previous studies ofhe Moodies Group to characterize subtidal deposition in a shal-ow marine shelf (Eriksson, 1977, 1979; Heubeck and Lowe, 1994).

oodies BIF likely represents the most distal facies of this progra-ational shelf sequence (Eriksson, 1983). Though generally shallow,

he precise water depth of deposition is difficult to constrain. Theetrital layers that are interbedded with the iron layers and chert doot show structures that unambiguously indicate deposition withinhe wave action zone. Furthermore, the parallel lamination found

fine layers of hematite. (D) Contact between an iron-rich bed (top dark band) andy preferentially occur at the boundary between iron-rich beds and chert. The chert

throughout the iron-rich layers suggests deposition of fine-grainedBIF precursor phases from suspension and not saltation. This dif-fers from the sedimentary structures that are common in granularPaleoproterozoic BIF and that are interpreted as the result of waveand current reworking of chemical clasts in shallow waters (Klein,2005). Parallel or wavy lamination and “soft sediment deforma-tion” are the only sedimentary structures that are observed in theclastic material that is intercalated at the cm scale between theiron and chert layers. Such structures may have formed in a tidalenvironment under the influence of offshore–onshore currents (e.g.Watchorn, 1980), but may also have formed deeper in to the basin,where the clastic material was transported by turbidity or com-bined flow currents (Myrow et al., 2002). Independently from thesetwo possible paleoenvironmental interpretations, the gradationalnature of increasing and decreasing frequency of BIF mesobandsinterbedded with siliciclastics suggests that the accumulation ofBIF was greatest during intervals of reduced sand supply, as asediment-starved or condensed facies.

5.2. Origin of banding

The presence of interbedded siliciclastics with the iron-richpackages and chert beds in the Moodies BIF provides new insighthelpful for understanding the origin of the typical BIF banding. Thechert occurs as beds and nodules with a laterally heterogeneous

122 T.R.R. Bontognali et al. / Precambrian Research 226 (2013) 116– 124

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ig. 7. Backscatter scanning electron photomicrographs of polished Moodies BIF thilosely associated with magnetite crystals. Note the amorphous shape of the particol–Ank, dolomite–ankerite. (B) Enlargement of region boxed in (A).

hickness; these textures are indicative of diagenetic precipitationrom pore fluids and do not require that the chert originally pre-ipitated from the water column (Beukes, 1984; Fischer and Knoll,009; Krapez et al., 2003). Furthermore, in stratigraphic context,hert is always preceded and followed by iron-rich packages (i.e.oes not occur in direct contact with siliciclastic beds) (Fig. 3).hese observations reveal a tight coupling between the accumu-ation of iron-bearing minerals and chert, supporting hypotheseshat naturally link the cycles of iron and silica. Grenne and Slack2005) proposed that gels protolith of chert were deposited byallout from hydrothermal fluids in silica-rich seawater, in whichlume-derived Fe oxyhydroxide particles promoted flocculation ofolloidal particles of silica–iron oxyhydroxide. Wang et al. (2009)efined this model, explaining the banding of BIF in terms of self-rganized chemical oscillations – ferric hydroxide precipitationrom Si- and Fe-rich hydrothermal fluids would decrease ambientH that subsequently causes silica precipitation. Thus, precipitationf chert is the result of a positive feedback linked to the precip-tation of iron. A different linkage was suggested by Fischer andnoll (2009), wherein the silica constituting the chert mesobandsas shuttled to basinal waters and sediments by adsorption on

he surfaces of ferric hydroxides. It is important to note that thisypothesis does not imply that BIF deposition was strictly associ-ted with hydrothermal settings.

.3. The mechanisms of Fe(II) oxidation

In the Moodies BIF, magnetite and hematite are the dominante-bearing mineral phases. Ferrous iron minerals are present, but

ions. (A, C and D) Carbonaceous particles with amorphous shape (white arrows) areplaced by the ingrowing magnetite (e.g. Beukes and Klein, 1992). Mag, magnetite;

somewhat rare compared to younger Archean deposits (Fischerand Knoll, 2009; Klein, 2005). Though ankerite was likely derivedfrom recrystallization of preexisting siderite, the sulfide-bearingminerals occur within small veins and cracks, and have a clearmetasomatic origin. Euhedral magnetite grains commonly showcrosscutting relationships with regard to hematite. Though thisphase represents a substantial amount of ferric iron in the MoodiesBIF, it has a late diagenetic or metamorphic origin. Similar observa-tions and interpretations have been made for the younger EarlyArchean and Paleoproterozoic BIF (Bekker et al., 2010; Beukes,1984; Ewers and Morris, 1981; Krapez et al., 2003).

The observation that hematite is an early phase while mag-netite a diagenetic or metamorphic mineral is consistent with thehypothesis that iron was concentrated in the sediments by oxida-tion of dissolved Fe(II) in seawater to form an insoluble hydrousoxide precipitate, a precursor that can spontaneously transforminto hematite on early diagenetic timescales (Ayres, 1972; Bekkeret al., 2010; Trendall and Blockley, 1970). Diagenetic and meta-morphic transformation of these ferric minerals to mixed valenceand ferrous minerals may have been driven by several processes,including microbial respiration of organic matter and abiotic ther-mal reactions, and are discussed further in the next section.

5.4. Role of biology and organic carbon in BIF deposition

In many models for BIF formation, primary oxidation of Fe(II)to Fe(III) is considered either the direct or indirect (via metabolicintermediates like O2) result of microbial activity (Brown et al.,

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995; Cloud, 1973; Harder, 1919; Hartman, 1984; Konhauser et al.,002; Posth et al., 2008; Walker, 1987). However, bona fide micro-ossils from BIF lithologies are very rare and none of the fewocumented occurrences (Barghoorn and Schopf, 1966; Klein et al.,987; Knoll and Simonson, 1981; Planavsky et al., 2009; Waltert al., 1976) provides evidence for a genetic relationship betweenicrofossils and Fe-rich minerals. For these reasons, the effective

nvolvement of microbes has been questioned and alternative abio-enic precipitation mechanisms have been suggested (Bratermant al., 1983; Draganic, 1991; Franc ois, 1986; Wang et al., 2009).

In the Moodies BIF, we observed carbonaceous particles withinhe iron-rich packages. The close spatial association with magnetitehases is consistent with the hypothesis that magnetite may haveormed during diagenesis by reduction of hematite (or its hydrousrecursor) with organic matter. This hypothesis, first proposed byerry et al. (1973) and subsequently refined by Walker (1984), Baurt al. (1985), Konhauser et al. (2005), and Fischer and Knoll (2009)mplies a two-fold role of microbes in BIF formation. First, microbesuild biomass in the water column either by directly oxidizingissolved Fe(II) through anoxygenic photosynthesis or by othersxygenic metabolism that cause passive precipitation of dissolvedron. Then, the precipitated Fe(OH)3 sinks from the water columnlong with part of the produced biomass. Within the sedimentshe biomass is consumed during a second microbial process as, forxample, dissimilatory iron reduction that couples organic matterxidation to Fe(III) reduction. This would explain both the min-ralogy observed in many BIF (i.e. the co-occurrence of mineralsith different Fe oxidation state that cannot precipitate simulta-eously from the same solution) and the lack of organic material

n BIF, although microbes might have played a key role in their for-ation. The organic carbon distribution in the Moodies BIF may

ecord these diagenetic reduction processes, though it is difficulto rule out abiotic alternatives. Importantly, simultaneous forma-ion of magnetite and organic carbon compounds through thermalecomposition of siderite was observed in laboratory experimentshat were carried out to evaluate the origin of putative biosigna-ures present in Martian meteorite ALH84001 (McCollom, 2003).he experiments were conducted at relatively mild temperaturesf 300 ◦C, a temperature not much higher than that experienced byhe Moodies BIF facies during lower-greenschist metamorphism,hough for a much shorter time. The thermal decomposition ofiderite to magnetite and organic carbon, thus, provides a possiblelternative explanation for the presence of carbonaceous particlesithin the Moodies BIF.

. Conclusions

The Moodies BIF share many features with younger Late Archeannd Paleoproterozoic iron formations, including a mineralogy thateflects a ferric iron phase as the earliest sedimentary precursor andhe secondary interaction of these ferric iron phases with reducinguids during diagenesis and metamorphism, and a sedimentaryode that is characterized by alternating iron-rich packages and

ilica-rich beds and concretions. The interbedded siliciclastic mate-ial indicates deposition during times of muted clastic input,erhaps at the base of a progradational shelf sequence. The precisealeodepth is difficult to constrain with any certainty – the Mood-

es BIF may have formed in a shallow, current-swept environmentr deeper into the basin, during the intervals between successiveurbidite events. The presence of siliciclastic beds provides an addi-ional insight on the origin of chert. The lack of chert accumulating

irectly on the siliciclastic material (i.e. the lack of chert not associ-ted with iron-rich bands) indicates a tight coupling between ironnd silica, and supports models positing that silica precipitationas a direct consequence of iron oxidation (Fischer and Knoll, 2009;

Research 226 (2013) 116– 124 123

Grenne and Slack, 2005; Wang et al., 2009). The presence of organicmatter within the Moodies BIF is consistent with the hypothesisthat BIF formed through the activities of iron-oxidizing and reduc-ing microbes as a part of the iron cycle. However, a metamorphicorigin of the observed organic matter through thermal dispro-portionation processes cannot be excluded. Despite some clearidiosyncratic differences, the similarities between the Moodies BIFand those deposited almost one billion years later support broadextrapolation of BIF forming processes back into Earth Archeantime.

Acknowledgments

We would like to thank Jan Kramers for introducing us to theBarberton Greenstone Belt, Chris Rippon, Charles Robus and JohnRobertson for supporting the fieldwork in the Barberton region andin the Agness gold mine, Drummond and Jacky Holman for theirsupport at Grace Farm, Mariarita Bontognali-Valli for organizingthe shipment of the BIF samples, Alexey Ulianov for his assistancewith optical microscopy and SEM investigations, Thierry Adatte forhis help with the XRD analyses, André Villard for the preparationof the thin sections.

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