Beukes and Gutzmer. 2008. Origin and Paleoenvironmental Significance of Major Iron Formations at the...

5IntroductionFOUR REVIEW PAPERS have recently been published that elo-quently summarize the current knowledge of the main com-positional attributes of iron formations and their distributionin space and time (Trendall, 2002; Simonson, 2003; Clout andSimonson, 2005; Klein, 2005). The authors of these reviewsalso highlighted the considerable questions that center on thegenesis and paleoenvironmental significance of this enigmaticrock type that is unique to the Precambrian Era. Rather thanto duplicate their effort, we focus on a critical review of

advances made in the understanding of the origin and pale-oenvironmental significance of the most voluminous and, ar-guably, best preserved iron formations in Earths history.These are the exceptionally well preserved iron formations ofthe Ghaap-Chuniespoort Groups of the Transvaal Super-group on the Kaapvaal craton in southern Africa (i.e., Beukes,1983), and the time-equivalent Hamersley Group of theMount Bruce Supergroup on the Pilbara craton in WesternAustralia (Fig. 1; Trendall, 2002; Pickard, 2003). Understand-ing the paleoenvironmental and paleogeographic setting ofthese giant iron formations and their associated strata is ofparticular significance, as geochemical characteristics of these

Chapter 1

Origin and Paleoenvironmental Significance of Major Iron Formations at the Archean-Paleoproterozoic Boundary

NICOLAS J. BEUKES AND JENS GUTZMERPaleoproterozoic Mineralization Research Group, Department of Geology, University of Johannesburg, P.O. Box 524,

Auckland Park 2006, South Africa

AbstractThis paper provides a critical review of advances made in understanding of sedimentary environments, geo-

chemical processes, and biological systems that contributed to the deposition and diagenetic evolution of theexceptionally well-preserved and large iron formations of the late Neoarchean to very early PaleoproterozoicGhaap-Chuniespoort Group of the Transvaal Supergroup on the Kaapvaal craton (South Africa) and the timeequivalent Hamersley Group on the Pilbara craton (Western Australia). These iron formations are commonlyassumed to have formed coevally but in separate basins, and they are often used as proxies for global oceanchemistry and paleoenvironmental conditions at ~2.5 Ga. However, lithostratigraphic and paleogeographic re-constructions show that the iron formations formed in a single large partly enclosed oceanic basin along themargins of the ancient continent of Vaalbara. Furthermore, although large relative to other preserved iron for-mations, the combined Transvaal-Hamersley basin is miniscule compared to marginal basins of the modernocean system so that the succession probably documents secular changes in depositional environments of thatbasin rather than of the global ocean at the time.

The iron formations comprise a large variety of textural and mineralogical rock types that display complexlateral and vertical facies variations on basinal scale. Based on detailed analyses of these variations it is con-cluded that the iron formations were deposited in environments that ranged from very deep-water basinal set-tings far below storm-wave base and the photic zone to very shallow-platform settings above normal wave base.Precipitation of both iron and silica took place from hydrothermal plumes in a dynamically circulating oceansystem that was not permanently stratified. Ferric oxyhydroxide was the primary iron precipitate in virtually allof the iron formation facies. This primary precipitate is now represented by early diagenetic hematite in someof the iron formations. However, in both deep- and shallow-water iron formations most of the original ferricoxyhydroxides have been transformed by dissimilatory iron reduction to early diagenetic siderite and/or mag-netite in the presence of organic carbon. Precipitation of ferric oxyhydroxides in very deep water below thephotic zone required a downward flux of photosynthetically-derived free oxygen from the shallow photic zone.In these deep-water environments, under microaerobic conditions, chemolithoautotrophic iron-oxidizing bac-teria may have played an important role in precipitation of ferric oxyhydroxides and acted as a source of pri-mary organic matter. With basin fill even shallow-shelf embayments were invaded by circulating hydrothermalplume water from which ferric oxyhydroxides could be precipitated in oxygenated environments with high pri-mary organic carbon productivity and thus iron reduction to form hematite-poor siderite- and magnetite-richclastic-textured iron formations.

Depositional models derived from the study of the iron formations along the Neoarchean-Proterozoicboundary can be applied to iron formations of all ages in both the Archean and later Paleoproterozoic. The fa-cies architecture of the iron formations determines to a large degree the textural attributes, composition, andstratigraphic setting of high-grade iron ores hosted by them. Detailed facies information thus would assist inimproving genetic models for high-grade iron ore deposits. Future research should be guided in this direction,especially in some of the very large iron ore districts of Brazil and India where very little is known about thecomposition and facies variations of the primary iron formation hosts and possible controls on localization ofhigh-grade ores.

Corresponding author: e-mail, [email protected]

2008 Society of Economic GeologistsSEG Reviews vol. 15, p. 547

Beukes_Gutzmer 6/11/08 7:40 AM Page 5

6 BEUKES AND GUTZMER

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Greenstone and Granitoids ofthe Pilbara Craton

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FIG. 1. Simplified geologic maps indicating the distribution of the Transvaal Supergroup on the Kaapvaal craton in south-ern Africa (A) (modified after Beukes, 1983) and the Hamersley Group on the Pilbara craton in Western Australia (B) (mod-ified after Trendall, 1983). Note that the two maps are presented on approximately the same scale, illustrating the muchlarger area of preservation of the Transvaal Supergroup relative to that of the Hamersley Group.


successions are often used as proxies for global ocean chem-istry at the Archean-Paleoproterozoic boundary (e.g., Anbaret al., 2007; Kaufman et al., 2007). The reliability of theseproxies is dependent entirely on the assumption that the twosuccessions were deposited in two contemporaneous but spa-tially separate open marine sedimentary basins. However, thismay not have been the case, as some authors (e.g., Cheney,1996) have provided stratigraphic evidence for deposition ofboth iron formation-bearing successions in one single sedi-mentary basin, thus requiring a much more careful differen-tiation of global and local basinal depositional and paleoenvi-ronmental signatures.

The aim of this paper is to constrain depositional environ-ments, geochemical processes, and biological systems thatcontributed to the deposition and diagenetic evolution of ironformations at ca. 2.5 Ga (i.e., the Archean-Proterozoic bound-ary). A modern sequence stratigraphic approach is taken, fo-cusing on depositional system facies tracts that were in placeduring deposition of the iron formations. We shall use termi-nology as defined by Van Wagoner et al. (1988) for the se-quence analyses. Important implications of sequence strati-graphic analyses for the understanding of the nature of thePaleoproterozoic-Archean ocean (Isley, 1995; Simonson andHassler, 1996; Isley and Abbott, 1999), and the role of micro-bial organisms in the precipitation of iron (Klein and Beukes,1989; Konhauser et al., 2002) are illustrated.

Within the context of this volume, it is hoped that an im-proved understanding of the stratigraphic and lithologic ar-chitecture of these giant iron formations will provide new im-pulses for exploration for high-grade iron ore deposits. Boththe Transvaal and the Hamersley Iron Formations are host toa number of economically important high-grade iron ore de-posits (see reviews by Gutzmer et al., 2005; Alchin et al.,2008; Thorne et al., 2008) and have a long history of explo-ration. Compositional differences, in particular the abun-dance of carbonates and iron oxides versus that of iron sili-cates and chert, in different iron formation lithofacies may beof critical importance to determining the suitability of partic-ular iron formation units to host epigenetic high-grade ironorebodies (e.g., Gutzmer et al., 2005). Indeed, it is throughthe detailed study of abundant exploration diamond drill corethat the lithostratigraphy and lateral and vertical facies varia-tions in these two giant iron formation successions are ar-guably the best established of all iron formations in the world(Trendall and Blockley, 1970; Beukes, 1983, 1984; Trendall,1983, 2002; Klein and Beukes, 1992a; Klein, 2005).

Mineralogical and Textural Classification of Iron Formations

Two broad textural types of iron formation are widely rec-ognized (Trendall, 2002; Simonson, 2003; Clout and Si-monson, 2005), namely banded iron formation (BIF) thatwas deposited as chemical muds and granular iron forma-tion deposited as endoclastic sands. However, this classifi-cation is too broad for the purpose of a detailed analysis ofdepositional facies and systems in Archean-Paleoprotero-zoic iron formations. Furthermore, the use of the term BIFfor iron formations that were deposited as chemical muds isconfusing as it is at odds with the commonly accepted de-finition of the term banded iron formation (Jackson, 1997,

p. 51). The latter definition restricts the term banded ironformation to iron formations with distinct chert banding, irre-spective of textural and/or mineralogical composition. This in-consistency is further illustrated in Klein (2005), who ex-tended the definition of the term banded iron formation to besynonymous to the term iron formation, irrespective of thepresence and/or absence of chert bands and texture, i.e.,muddy versus granular. There is thus a clear need to stan-dardize the nomenclature. In this paper, we propose andapply a revised classification of iron formation that acknowl-edges the classical definitions of iron formation and bandediron formation and subdivides these two rock types furtheraccording to textural and mineralogical attributes.

Following the Glossary of Geology (Jackson, 1997, p. 335),iron formation (IF) is defined as a finely-laminated to thinbedded chert-bearing chemical sedimentary rock containingat least 15% iron of sedimentary origin. Banded iron forma-tion (BIF), on the other hand, is a variety of iron formationthat contains distinct chert bands (Jackson, 1997, p. 51), i.e.,mesobands of Trendall and Blockley (1970). Subdivision ofiron formations further takes place according to textural fea-tures that are in many respects similar to those of carbonaterocks (Dimroth and Chauvel, 1973; Dimroth, 1976). Basedon mineralogical and textural attributes Beukes (1980a, 1983,1984) developed a comprehensive classification scheme thatcovers the full range of iron formation rock types and lithofa-cies. However, that classification scheme specifically appliesto detailed sedimentological analyses of iron formations; forthis paper a simplified classification will suffice. This simpli-fied classification is based on the three basic textural com-ponenets of iron formations, namely, allochem particles(granules), matrix (chemical iron-rich muds, i.e., femicrite),and microcrystalline quartz (chert; Fig. 2A). The latter oftenrepresents the cement component in iron formations (Fig.2A; Simonson, 1987). Iron formations composed mainly ofgranules are referred to as granular iron formations (GIF).Granular iron formation encompasses both conglomeraticferudite and sandy fearenite (Fig. 2B). Fearenite can in turnbe subdivided into grainstone, packstone, or wackestone, de-pending on sorting and the amount of femicrite matrix pre-sent (Fig. 3), following definitions of Dunham (1962) for cal-carenite. It is recognized that virtually all of the granules areendoclastic in origin (Simonson, 1985) and derived from re-working of earlier lithified iron formation components, andthat chemically precipitated grains, like oolites, are in factscarce (Simonson, 2003; Clout and Simonson, 2005). Granu-lar iron formations that are distinctly chert mesobanded couldbe described as banded granular iron formations.

In contrast to granular iron formations, iron formationsdominantly composed of femicrite (chemical mud) are clas-sified as micritic iron formation (MIF) and when distinctlychert-mesobanded they are referred to as banded micriticiron formation (BMIF) (Fig. 2B). Femicrite in micritic ironformation displays three distinct types of bedding, namely(1) microlamination or microbanding (on a submillimeterscale), (2) lamination on a millimeter to centimeter scale,and (3) thin to thick, poorly defined bedding with rathermassive appearance. Beukes (1980a, 1983, 1984) applied theterm ferhythmite to encompass microbanded and well-lami-nated femicrite; in contrast, the term felutite is applied to

MAJOR IRON FORMATIONS AT THE ARCHEAN-PALEOPROTEROZOIC BOUNDARY 7

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poorly bedded femicrite with a muddy massive appearance.Based on these parameters micritic iron formations can bebroadly classified into microbanded, laminated, and lutiticvarieties (Fig. 2A, B). In laminated banded micritic iron for-mation and micritic iron formation, individual laminae mayhave a massive or graded appearance and in some cases beseparated from each other by micritic microbands. This tex-tural variability may reflect the difference between current-reworked femicrite (massive or graded laminae) versus set-tling of fine suspended mud in microbands. Felutite beds areoften associated with granular iron formations and may havebeen reworked by currents as indicated by the presence of

fine femicrite intraclasts (Beukes, 1980a, 1983, 1984). Thereis thus a whole range of textural types of micritic iron forma-tion, ranging from microbanded ones that represent chemi-cal precipitates (orthochemical iron formations) throughmixtures between chemical precipitates and reworked femi-crite in laminated iron formation (best described as ferhyth-mite) to reworked femicrite in felutite, i.e., allochemical mi-critic iron formation (Fig. 2B). In sedimentological contextgranular iron formation and lutitic micritic iron formationcan be grouped as clastic-textured iron formations, in con-trast to the orthochemical microbanded micritic iron forma-tion (Fig. 2B).


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Chert CementChert Bands

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Notes:1) Description according to form of chert bands: Podded BIF, Wavy BIF, Even BIF2) Laminated and microbanded MIF or BMIF can be grouped as rhythmitic MIF or BMIF3) Mineralogical composition described by prefix according to mineralogical classification above

CLASSIFICATION SYSTEMEdgewise conglomerate

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FIG. 2. Classification and nomenclature of iron formation based on the nature of its major components (allochem parti-cles, femicrite, and chert) (modified after Beukes, 1980a, 1983). See text for explanation of acronyms.


Chert mesobands are the third important component ofiron formations. The chert mesobands can have either mas-sive or microbanded internal appearance (Fig. 2A). Mi-crobanded chert mesobands are typically, but not exclusively,associated with microbanded femicrite and massive ones withfelutite. The bedding character of chert mesobands can varyfrom even, wavy to podded forms (Fig. 2A). This geometrycan be incorporated in the description of BIF lithofaciesthrough terms such as podded, wavy, or even BIF (Fig. 2B).Chert pods have a flat lenticular form in cross section and aredisks in three dimensions. They vary from millimeter todecimeter in diameter and thickness.

Jaspilite is a term applied only to BIF composed of alter-nating mesobands of jasper (i.e., chert with minor, finely dis-seminated microcrystalline dusty hematite) and hematite magnetite (Beukes, 1980a). It should be noted that, accordingto our own observations, apart from being a primary precipi-tate, jaspilite in iron formation successions could also be aproduct of supergene or hydrothermal alteration.

Iron formations and BIFs can be further classified accordingto their mineralogical composition. Well-preserved, unmeta-morphosed iron formations, like those of the Transvaal andHamersley basins, are characterized by three groups of veryfine crystalline iron-rich minerals, namely the oxides hematiteand magnetite, the carbonates siderite and ankerite, and theiron silicates greenalite and stilpnomelane with rare chamosite(Klein, 2005). These diagenetic to low-grade metamorphicmineral assemblages define three end-member mineralogicalfacies of iron formations, i.e., oxide, carbonate, and silicatewith several mixed facies (Fig. 2A). James (1954) recognizeda sulfide facies characterized by the presence of either pyriteor pyrrhotite. Because of its high siliciclastic content, it is ap-propriate to describe the latter as pyritic black shale, with orwithout interbeds of black carbonaceous chert, and not as aniron formation lithofacies. Indeed, it is our experience thatearly diagenetic pyrite is conspicuously absent from iron for-mations that have not been contaminated by an influx of exo-genic detrital or pyroclastic material.

The effects of regional metamorphism on the mineralogy ofiron formations have recently been reviewed by Klein (2005).Incipient regional metamorphism of iron formation leads torecrystallization of components, increased grain size, and de-velopment of such minerals as biotite from stilpnomelane,minnesotaite from greenalite and grunerite, and cumming-torite-grunerite from reactions between carbonates or ironoxides with silica. High-grade metamorphic iron formationscontain minerals such as coarse magnetite, clino- and or-thopyroxene, and fayalite. Magnetite, a very common compo-nent of iron formation, is most abundant as a product ofmetamorphic or hydrothermal alteration and less commonlypreserved as a product of early diagenesis. The origin of mag-netite and its relationship to hematite, which is one of the ear-liest diagenetic components of iron formation, can in mostcases only be established through very careful petrographicstudies (Han, 1978). Neither calcite nor dolomite is consid-ered primary or early diagenetic constituents of iron forma-tion. Rather, they are of metamorphic or hydrothermal originwith specifically calcite a by-product in metamorphic forma-tion of grunerite from ankerite and quartz. Both dolomiteand calcite are also found with hydrothermally altered iron

formation associated with high-grade iron ore deposits(Beukes et al., 2002a; Clout and Simonson, 2005).

Iron formations uncontaminated by endogenetic siliciclas-tic or pyroclastic material are depleted in virtually all of themajor elements except iron and silica. Iron contents typicallyrange between ~20 and 35 wt percent and SiO2 between 40and 55 wt percent (Klein, 2005). Iron formations are furthermarked by very low Al2O3 contents (usually

of the Pilbara craton, and the Shushong Group (Beukes,1973) along the northern margin of the Kaapvaal craton (Fig.4). Only two occurrences of Mesoproterozoic Superior-typeiron formations are known, namely in the middle Mesopro-terozoic Bushmanland Group of the Namaqua terrane of theKalahari craton ( Strydom et al., 1987; McClung, 2006) andlate Mesoproterozoic shelf carbonate succession of the Pen-ganga Group (Gutzmer and Beukes, 1998a) of PeninsularIndia (Figs. 3B, 4).

Although they are not to be discussed further in this paper,it is worth noting that Algoma- and Rapitan-type iron forma-tions are distinctly different from Lake Superior-type ironformations in their abundance and distribution. Algoma-typeiron formations, for example, appear much more abundantboth in number (Fig. 3B) and geographic distribution (Fig.4). However, much of this merely reflects the less stable tec-tonic environment in which iron formations in greenstonebelts were deposited and the effects of intense postdeposi-tional deformation. It does not imply that iron formations ingreenstone belts are thinner and were initially laterally lessextensive than those of the Lake Superior type. Iron forma-tions in greenstone belts can be up to several hundred metersthick (Fedo and Eriksson, 1996) and laterally extensive forseveral hundred kilometers in folded greenstone belt succes-sions (Klein, 2005). Indeed, many Algoma-type iron forma-tions display attributes very similar to that of Lake Superior-type iron formations (Eriksson, 1983; Eriksson et al., 1994).

Neoproterozoic Rapitan-type iron formation occurrencesare highly localized, relative to the very wide distribution ofglacial diamictites with which they are associated. By far thelargest and best preserved Rapitan-type iron formationsoccur in the Rapitan Group of the Neoproterozoic inlier inthe McKenzie and Ogilvie Mountains in North America (e.g.,Klein and Beukes, 1993) and the Urucum district (e.g., Kleinand Ladeira, 2004) along the western margin of the Amazoncraton in South America (Fig. 4). Smaller Neoproterozoiciron formations are also known from the Damara-Gariep suc-cession in Namibia and South Africa, the Flinders Range inAustralia (Lottermoser and Ashley, 2000), and the Maliya-Khingun Proterozoic inlier (James, 1983) in far southeasternSiberia (Fig. 4).

Stratigraphic Setting and Correlation between theNeoarchean-Paleoproterozoic Successions of the

Transvaal and Hamersley Basins

Stratigraphic setting

In the Transvaal basin on the Kaapvaal craton the giant ironformations along the Archean-Proterozoic boundary consti-tute the Asbesheuwels Subgroup and correlative Penge IronFormation of the Ghaap and Chuniespoort Group, respec-tively (Fig. 1A; Beukes, 1983, 1984). On the Pilbara cratonthey are represented by the Marra Mamba, Brockman, WeeliWolli, and Boolgeeda Iron Formations of the Hamersley


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Group (Fig. 1B; Trendall and Blockley, 1970; Trendall, 1983;Trendall et al., 2004). The Asbesheuwels succession, com-prised of the Kuruman and conformably overlying Griqua-town Iron Formation, reaches a maximum known thickness of950 m and together with the correlative Penge Iron Forma-tion (Fig. 1A), covers an area of ~110,000 km2 on the Kaap-vaal craton (Beukes, 1983). The combined thickness of theBrockman, Weeli Wolli, and Boolgeeda Iron Formations onthe Pilbara craton is ~1,050 m covering an area of ~60,000km2 (Trendall, 1983).

On both cratons the iron formations conformably overliethick late Neoarchean carbonate platform deposits known asthe Campbellrand-Malmani succession in the Tranvaal basinand the Wittenoom Dolomite in the Hamersley basin (Fig.1). However, it is important to note that the Asbesheuwels-Penge succession comprises well-exposed iron formationsthat extend from the basinal facies of the underlying Camp-bellrand carbonate platform succession onto the shallowshelf (Fig. 5). This basin to shelf transition provides excellentopportunities for understanding the influence of paleoba-thymetry on the nature of iron formation lithofacies (Beukes,1983, 1984). This is not the case in the Hamersley basin ofWestern Australia, where only the basinal facies are well ex-posed and the transition from basin to shelf on the underly-ing Wittenoom-Carawine carbonate platform succession iseither covered by younger rocks or not preserved by erosion(Fig. 5), giving a depositional picture biased toward deeperwater lithofacies. The present contribution will first docu-ment the excellent lithostratigraphic correlation of the suc-cessions of the Transvaal and Hamersley basins. It will then

focus on the lithofacies architecture and depositional envi-ronments of iron formations in the Transvaal basin only,mainly because such data are currently not available for theHamersley succession on a basinwide scale.

Lithostratigraphic correlation

Button (1976a) was arguably the first to point out strikinglithostratigraphic similaritiesand a possible link betweenthe Transvaal and Hamersley basins. Cheney (1996) formal-ized the concept that the two successions may have formedpart of one large basin along the trailing margin of an ancientcontinent referred to as Vaalbara. Since these early contribu-tions, abundant new data have become available that supportthe concept in general but neccessitate revision of some ofCheneys (1996) sequence stratigraphic correlations and hispaleogeographic reconstruction of the relative positions ofthe Kaapvaal and Pilbara cratons.

The correlation between the two successions proposedhere (Fig. 6) is equally founded on stratigraphic and sedi-mentological similarities aided by correlation of unconformi-ties and event beds and first-hand field observations by thesenior author. The correlation was then validated by availableisotopic age data (Fig. 6). It is important to note that correla-tion is restricted to the basinal facies of the Transvaal Super-group (to the south of the Griquatown fault zone; Fig. 5) andthe basinal facies of the Hamersley Group. This is justified bythe poor preservation of the shallow-platform facies of theHamersley Group (Fig. 5). The lithostratigraphic correlationdiagam takes into account only strata from the base of theTransvaal Supergroup to some stratigraphic distance abovethe Koegas Subgroup (Fig. 6). Strata below and above thesebeds may also correlate (Cheney, 1996), but this has little rel-evance to this paper.

In the proposed correlation the basal quartzite of the Vry-burg Formation of the Transvaal Supergroup, which discon-formably overlies basaltic andesites of the Allanridge Forma-tion of the Ventersdorp Supergroup, correlates with theWoodiana sandstone that overlies basalts of the Maddina For-mation and forms the base of the Jeerinah Formation in theupper part of the Fortesque Group in Western Australia (Fig.6). Conformably overlying the Vryburg Formation is a car-bonaceous shale unit known as the Lokammona Formationthat correlates with the upper part of the Jeerinah Formation(Fig. 6). To the north of Balfour Downs in the Carrawine area(Fig. 1B), the upper part of the Jeerinah Formation is knownas the Roy Hill Shale; it contains carbonates with contortedmicrobial mats similar to those of the lower part of the Mon-teville Formation at the base of the Campbellrand Subgroup(Fig. 6). An impact spherule layer (tectite) in the lower partof the Monteville Formation (Simonson et al., 1999) is corre-lated with a spherule layer in the Jeerinah Formation (Fig. 6;Kohl et al., 2006).

In the Hamersley basin, the Jerrinah Formation, includingthe Roy Hill Shale, is conformably overlain by the MarraMamba Iron Formation (Fig. 6). In the Transvaal basin thereis no such prominent iron formation developed but similari-ties between carbonates in the Roy Hill Shale and that of thelower part of the Monteville Formation indicate that the low-ermost ankerite-banded chert unit in the Campbellrand suc-cession may correlate with the Marra Mamba Iron Formation


0361-0128/98/000/000-00 $6.00 12

Younger unconformity

CarawineDolomite

JeerinahMarra Mamba

Wittenoom

Brockman -Boolgeeda

Unconformity

Basin Shelf

WSW ENE

Pilbara Craton (Western Australia)Basin Shelf

LokammonaMonteville BIF

Nauga

Asbesheuwels

Koegas

CampbellrandDolomite

UnconformitySW NE

Kaapvaal Craton (South Africa)

Shale

BIF

Dolomite

Shale

Shale

BIF

DolomiteBIFShale

Cover

FIG. 5. Schematic comparison between late Neoproterozoic to very earlyPaleoproterozoic stratigraphic successions hosting giant iron formations onthe Kaapvaal and Pilbara cratons. Note that on the Pilbara craton only thebasinal facies of the Brockman-Boolgeeda Iron Formation succession is pre-served, whereas on the Kaavaal craton both basinal and shelf facies of the As-besheuwels succession is preserved.



0361-0128/98/000/000-00 $6.00 13

Mn

S S

S S S

S

S SS

Klein Naute

Monteville

LokammonaVryburg

Allanridgelava

VentersdorpSupergroup

Machadodorp SilvertonDaspoort

Ongeluk/HekpoortMakganyene

Timeball Hill

Upper

Lower

Ko

egas

GriquatownIF

Kur

uman

IF

Middelwater

PannetjieDoradaleNaragas

RooinekkeKlipputsNelani

Duitschland

2440 - 2446

232215

222013

ca. 2440

24806

OrangeView

2521325497

Tuff 2

Tectite

Tectite

26423

2597526295

26846

Jeerinah

Marra Mamba

MaddinaBasalt

Woodiana SS

Roy Hill Sh

Fort

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om Main Tuff25659Tectite

Bruno s Band

Colonial Chert25018

249516GorgeDales 24814Whaleback 24635

Mt McRae

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Yandicoogina

24493

Woongarra

LowerBoolgeeda

Upper

2445524515

Bro

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24615

Tu

re

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Folded Sill2208 Ma

Meteorite Bore

OphthalmianOrogeny

Post-MooidraaiFolding

Cheela SpringsBeasley

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NumanaW

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rou

p

KliphuisTectite

Stofbakkies

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5

6

7

9

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8

11

2464624657

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ley

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Mt Sylvia

West Angela

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2588612

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7

FIG. 6. Correlation diagram between late Neoarchean to early Paleoproterozoic strata in the Transvaal and Hamersleybasins. Age references are as follows: 1 = Trendall et al. (2004), 2 = Barley et al. (1997), 3 = Mller et al. (2005), 4 = Anbaret al. (2007), 5 = Pickard (2003), 6 = Altermann and Nelson (1998), 7 = Nelson et al. (1999), 8 = Gutzmer and Beukes(1998b), 9 = Sumner and Bowring (1996), 10 = Hannah et al. (2004), 11 = Cornell et al. (1996) and Dorland (2004), 12 =Martin et al. (1998).


(Fig. 6). This correlation is supported by the fact that the cor-relative Monteville-Jeerinah impact spherule layer is devel-oped in similar position below both the ankerite-bandedchert unit of the Monteville Formation and the MarraMamba Iron Formation (Fig. 6).

Above the Marra Mamba Iron Formation the transitionzone into the overlying Wittenoom carbonate succession iscomposed of carbonaceous shale and carbonates. This unit,known as the West Angela Member of the Wittenoom For-mation (Blockley et al., 1993), bears resemblance to the Mon-teville Formation (Fig. 6) above the ankerite-banded chertmarker in the basinal area of the Campbellrand succession tothe south of the Griquatown fault zone (Beukes, 1987).

The Wittenoom Dolomite of the Hamersley basin (Simon-son et al., 1993a, b) displays many features that can also beobserved in the basinal Nauga Formation (Beukes, 1980c,1987) of the Campbellrand carbonate platform (Fig. 6). Twotuff-rich zones, known as the Main and Crystal-rich tuff in-tervals in the Wittenoom succession (Simonson et al., 1993a,b) and Tuff zones 2 and 3 in the Nauga Formation (Beukes,1980c), respectively, can also be correlated based on theirclose stratigraphic spacing in both basins. However, perhapsthe stratigraphic interval with the most remarkable similarityis the one that caps the carbonate succession in both basins.In the Hamersley basin, where it is known as the MountSylvia Formation, it is only about 30 to 40 m thick and com-posed of two distinct and persistent lower ferruginous chertbeds overlain by carbonaceous shale with interbeds of car-bonate and capped by a very persistent thin BIF informallyknown as Brunos Band (Fig. 6; Simonson et al., 1993a, b). Inthe Transvaal basin a virtually identical succession caps theNauga Formation in the basinal facies of the CampbellrandCarbonate platform. It has not been given formal strati-graphic status, but Beukes (1980) referred informally to theupper iron formation as BIF-1 and the two ferruginous chertbeds as BIFs 2 and 3. It is this upper BIF-1 that is in thispaper informally referred to as Brunos BIF after its proposedcounterpart in Western Australia (Fig. 6). One prominent andlaterally persistent event bed in the Wittenoom succession ofwhich no counterpart has yet been found in the Nauga suc-cession, is the so-called Wittenoom impact spherule layer thatis present a few tens of meters above the Main Tuff intervalin the Hamersley basin (Fig. 6).

Above Brunos Band, black carbonaceous shale of theMount McRae Formation in the Hamersley basin would cor-relate with the Klein Naute Shale in the basinal area of theCampbellrand carbonate platform succession (Fig. 6). InWestern Australia the chert- and shale-rich unit that under-lies the Brockman Iron Formation, is referred to as the Colo-nial Chert Member of the Mount McRae Shale (Fig. 6). InSouth Africa this unit is represented by the Kliphuis Memberof the Kuruman Iron Formation (Fig. 6).

The overlying major iron formation succession displaysmany similarities that permit correlation. As pointed out byCheney (1996), the most convincing correlation is betweenthe Dales Gorge Member of the Brockman Iron Formationand the Stofbakkies Member of the Kuruman Iron Formation(Fig. 6). It has long been recognized that the Dales GorgeMember is composed of 16 so-called shale macrobands (Sbands) each overlain by a BIF macroband (Trendall and

Blockley, 1970). These shale-BIF macrocycles correspond tostilpnomelane lutite-banded micritic iron formation macrocy-cles as defined by Beukes (1978, 1980b) in the Kuruman IronFormation. It may be pure chance, but it is interesting that inthe type profile of the Kuruman Iron Formation near Kuru-man, the Stofbakkies Member also comprises 16 such macro-cycles, starting with the first stilpnomelane lutite bed that di-rectly overlies the Kliphuis Member (refer to fig. 5 in Beukes,1980b). Although Cheney (1996) also noted this correlation,he recognized 17 BIF units, because he included the upperBIF of the Kliphuis Member. This, however, would corre-spond to BIF-O of Trendall and Blockley (1970) and Trendall(1983) at the base of the Dales Gorge Member of the Brock-man Iron Formation. For reference purposes, it should benoted that the Stofbakkies Member, as defined in this paperfor simplicity (Fig. 6), corresponds to the Matlipani andWhitebank members of the type profile of the Kuruman IronFormation as originally defined by Beukes (1980).

In addition to the detailed correlation of macrocycles, theequivalent of an impact spherule layer from S-4 in the DalesGorge Member (Hassler and Simonson, 2001) has been recog-nized in equivalent stratigraphic position in the StofbakkiesMember of the Kuruman Iron Formation (Fig. 6). The spherulebed was first found by Van Wyk (1987) in drill core from nearPomfret but interpreted at the time as volcanic lapilli. More re-cently, the layer was intersected in core from the AgouronDrilling Project and interpreted as of possible meteorite impactorigin (Schroeder et al., 2006; Simonson et al., 2006).

Above the Stofbakkies Member, the stilpnomelane-richBuisvlei Member is correlated with the Whaleback Shale andthe Orange View Member of the Kuruman Iron Formationwith the Joffre Member of the Brockman Iron Formation ofthe Hamersley Group (Fig. 6). The Orange View and JoffreMembers both represent rather monotonous successions ofmagnetite-siderite banded micritic iron formation with somestilpnomelane lutite interbeds (Beukes, 1980b; Trendall andBlockley, 1970).

The Yandicoogina Shale Member that forms the top of theBrockman Iron Formation is considered correlative to thestilpnomelane-rich unit that marks the base of the Wester-berg Member in the upper part of the Kuruman Iron Forma-tion (Fig. 6). Above that the finely laminated micritic iron for-mation of the Westerberg and Geduld Members correspondsto the Weeli Wolli Iron Formation of the Hamersley Group.Similar to the Westerberg and Geduld Members, the WeeliWolli Iron Formation is also characterized by a lack of distinctchert mesobands (Trendall, 1973). However, it is typically afinely microbanded magnetite-hematite micritic iron forma-tion, whereas the Westerberg and Geduld Members aregreenalite rich. The Weeli Wolli Iron Formation is furthercharacterized by the presence of several diabase sills and thesame applies to the Westerberg and Geduld Members. Lo-cally, a basaltic pillow lava unit is present in the Weeli WolliFormation (Barley et al., 1997) that has no obvious correlativein the Transvaal basin (Fig. 6).

The Weeli Wolli Formation is overlain by the ~400-m-thickWoongarra rhyolite which is considered to comprise both ex-trusive and intrusive phases (Doyle et al., 2001), althoughTrendall (1995) argued for it to be an intrusion. If the corre-lation of the Geduld Member with the upper part of the


0361-0128/98/000/000-00 $6.00 14


Weeli Wolli Formation is accepted, then the Woongarra rhy-olite should have extruded coevally with the lower part of theMiddelwater Member of the Griquatown Iron Formation(Fig. 6). Although no equally prominent volcanic unit hasbeen identified in the Transvaal basin, it is interesting to notethat the Griquatown Iron Formation at this stratigraphic levelcontains abundant stilpnomelane lutite beds in the Pomfretarea (Van Wyk, 1987). These stilpnomelane beds most prob-ably represent pyroclasic beds derived from acid volcanism(Van Wyk, 1987) and may be tentatively correlated with theWoongara igneous event.

The Woongarra rhyolite is overlain by the Boolgeeda IronFormation that is microbanded but does not contain well-de-fined chert mesobands (Trendall and Blockley, 1970). It is di-vided into a Lower and Upper unit by a coarsening-upwardsiliciclastic shale-fine quartzite unit that is similar to the Pan-netjie Formation of the Koegas Subgroup (Fig. 6). Based onthis observation, the lower Boolgeeda Iron Formation is ten-tatively correlated with the upper part of the basinal Middel-water Member of the Griquatown Iron Formation and theupper Boolgeeda Iron Formation with the Doradale IronFormation of the Koegas Subgroup (Fig. 6).

The Kungarra Formation, which overlies the BoolgeedaIron Formation with sharp contact and forms the base of theTuree Creek Group of the Mount Bruce Supergroup in theHamersley basin (Fig. 6), is composed of fine siltstone,greywacke, and quartzite (Martin et al., 1998), which is simi-lar in character to the Naragas Formation of the Koegas Sub-group (Beukes, 1983) and could be correlative (Fig. 6). Aglacial diamictite, known as the Meteorite Bore Member(Trendall and Blockley, 1970), is developed in the KungarraFormation (Fig. 6). The classification of the diamictite as amember within the Kungarra Formation may, however, bevery misleading as the diamictite contains boulders of the un-derlying Woongarra rhyolite (Trendall, 1983). The MeteoriteBore diamictite thus probably overlies a major low-angle un-conformity in the succession and should mark the base of anew sequence. In the Transvaal area of the Transvaal Super-group, a glacial diamictite is developed near the base of theDuitschland Formation (Martini, 1979; Bekker et al., 2001)and the Meteorite Bore Member is considered a correlativeof that diamictite (Fig. 6; Cheney, 1996). The presence of anerosional unconformity at the base of the Meteorite Bore di-amictite may explain why iron formations equivalent to theRooinekke-Nelani succession of the Koegas Subgroup is notpresent in the Hamersley basin (Fig. 6).

Correlation of the succession above the diamictite falls out-side the focus of this paper. However, it is worthwhile notingthat in contrast to comparisons by Button (1976a) and Ch-eney (1996), the Cheela Springs Basalt is considered equiva-lent to the Machadodorp Basalt of the Pretoria Group and notthe Ongeluk/Hekpoort succession (Fig. 6). The erosional un-conformity at the base of the Beasley River Quartzite is thuscorrelated with the unconformity below the DwaalheuwelFormation of the Pretoria Group along which the well-knownHekpoort paleosol (Rye and Holland, 1998; Beukes et al.,2002b) is developed. A thick succession of strata, includingthe Makganyene diamictite (Evans et al., 1997), has thus ap-parently been removed in the Hamersley basin during the~2200 Ma Opthalmian orogeny which corresponds to the

post-Mooidraai and/or pre-Dwaalheuwel fold event (Beukeset al., 2002b) in the Transvaal basin (Fig. 6).

Isotopic ages and rock accumulation rates

Available U-Pb ages obtained on zircons in tuffaceous sed-iments (usually by SHRIMP) generally show good agreementbetween stratigraphic units in the Transvaal and Hamersleybasins that are correlated based on lithologic grounds (Fig. 6).However, inherited zircons abound in some of the tuffaceousbeds and in some cases produce ages evidently older thansediment deposition (Alterman and Nelson, 1998; Pickard,2002, 2003; Trendall et al., 2004). This renders interpretationof depositional age somewhat inconclusive. An example is theSHRIMP U-Pb age obtained by Pickard (2003) on zircon sep-arated from stilpnomelane lutite beds from the lower part ofthe Geduld Member of the Kuruman Iron Formation (Fig.6). The preferred age reported for the stilpnomelane beds byPickard (2003) is around 2460 Ma. However, calculation ofthe ~2460 Ma age ignores what are referred to as youngerstatistical outliers by Pickard (2003). If these are taken intoconsideration the depositional age could be on the order of2440 to 2446 Ma, which corresponds very well with that ofthe Weeli Wolli Iron Formation in the Hamersley basin (Fig.6).

Although in the correlation diagram strata in the basinal fa-cies if the two successions are correlated, it should be notedthat samples of tuffaceous beds for which U-Pb ages of 2588 6 Ma (zircon by SHRIMP) and 2521 3 Ma (zircon byTIMS) were obtained (Fig. 6) are from the shallow shelf andnot the basinal facies of the Campbellrand succession. Simi-larly, a 2488 6 Ma U-Pb age on zircons reported in Figure6 was obtained on a stilpnomelane lutite sample from the cor-relative unit of the Stofbakkies Member in the Penge IronFormation (Transvaal region, shallow-platform facies). Fur-thermore, three of the ages reported in Figure 6 are not U-Pb ages obtained on zircon. These are (1) a 2501 8 Ma Re-Os age on pyrite for the Mount McRae shale (Anbar et al.,2007), (2) a 2322 15 Ma age Re-Os age on pyrite for theTimeball Hill Formation (Hannah et al., 2004), and (3) a 2220 13 Ma Pb-Pb on whole-rock age for the Ongeluk Lava(Cornell et al., 1996). The latter age has recently been con-firmed by U-Pb analyses of zircons from tuffaceous beds inthe Hekpoort Lava (Dorland, 2004).

The ages provide a means of estimating rock accumulationrates, also referred to as compacted sedimentation rates(Pickard, 2002, 2003) or characteristic depositional rates(Trendall et al., 2004). Several publications deal with this as-pect of deposition of the iron formations of the Asbesheuwelsand Hamersley Groups (Arndt et al., 1991; Barley et al., 1997;Alterman and Nelson, 1998; Pickard, 2002, 2003; Trendall etal., 2004). The calculated rock accumulation rates for BIFunits vary greatly from 3 to 4 to 180 m/m.y. Of course, in suc-cessions with major differences in stratigraphic thickness be-tween shallow shelf and basin, such as the Campbellrand andAsbesheuwels Subgroups, rock accumulation rates must havevaried accordingly. Based on the available ages of 2588 Ma atthe base and 2521 Ma near the top of the Campbellrand suc-cession (Fig. 6) an overall rock accumulation rate of ~33m/m.y. is calculated for the shelf and 8 m/m.y. for the basinalfacies. Similarly, for the Kuruman Iron Formation, using an


0361-0128/98/000/000-00 $6.00 15


age of 2495 Ma near the base and 2464 Ma near the top of theOrange View Member (Fig 6) a rock accumulation rate of 16to 17 m/m.y. is calculated for the basin and 4 to 5 m/m.y. forthe shallow shelf.

Reconstruction of the Transvaal-Hamersley basin

The excellent correlation between the Hamersley andTransvaal successions strongly suggest that deposition tookplace in a single basin along the margin of Vaalbara, as origi-nally proposed by Cheney (1996). This raises the questionabout the relative and absolute paleogeographic positions ofthe Kaapvaal and Pilbara cratons at the time of deposition. Inthe reconstruction of Cheney (1996), the Pilbara craton isplaced to the south of the Kaapvaal craton. However, recentpaleomagnetic analyses by De Kock (2007) and De Kock etal. (2007) place the Pilbara craton immediately to the north-northwest of the Kaapvaal craton. De Kock (2007) stressesthat several structural and depositional elements in the twobasins support this configuration.

According to the reconstruction put forward by De Kock(2007) and De Kock et al. (2007) the basin in which theHamersley-Asbesheuwels Iron Formation succession was de-posited consisted of a shallow-shelf area to the east and adeep basin to the west. A large embayment is envisaged be-tween the two cratons (Fig. 7) because of the known arcuateshelf margin along the western side of the Kaapvaal craton(Beukes, 1983, 1984). If a volcanic arc situated some distanceoff Vaalbarra to the west is invoked, iron formation depositionwould have taken place in a back-arc environment, as sug-gested for both the Hamersley (Blake and Barley, 1992) andthe Asbesheuwels successions (Klein and Beukes, 1989). Thisvolcanic arc could also have been a prolific source of hy-drothermal fluids needed for deposition of the iron forma-tions. Also indicated in the basin reconstruction is a proposed

source for mafic tuffs during deposition of the Wittenoom-Campbellrand carbonate platform, based on results of Has-sler (1993), and a volcanic center responsible for outflow ofpillow lavas during deposition of the Weeli Wolli Iron For-mation and formation of the Woongarra rhyolite (Barley et al.,1997) shortly thereafter (Figs. 6, 7).

The successful reconstruction of the combined Hamersley-Transvaal basin implies that the two largest known iron for-mation successions in the world were deposited in one large,possibly restricted (Hassler, 1993) oceanic basin at ~2,44 to2,50 Ga. Similarities in secular variations in depositional envi-ronments in the Hamersley and Transvaal successions maythus, for the most part, be only of regional significance. Localinfluences on deposition would have to be carefully evaluatedbefore any interpretations about possible global evolutionarychanges can be made from studies of correlative rock units inthe two areas.

Iron Formations in the Campbellrand Carbonate Platformand Transition to the Kuruman Iron Formation

The Neoarchean Cambellrand carbonate platform succes-sion of the Transvaal basin (Figs 1, 5, 6) contains several thinankerite-banded chert units, so-called proto-iron formationsof Button (1976b), and two thin banded micritic iron forma-tion units (Fig. 8; Beukes, 1987; Sumner and Beukes, 2006).The two iron formations, known as the Kamden Member inthe middle and Brunos BIF near the top of the carbonateplatform succession (Fig. 8), are believed to hold fundamen-tal information about the depositional systems in lateNeoarchean oceans that eventually led to deposition of theoverlying giant iron formations of the Asbesheuwels Sub-group (Figs. 1, 8). The depositional setting of these iron for-mations and the transition into the basal units of the Kuru-man Iron Formation (Fig. 8), thus needs to be discussed indetail.

Regional facies architechture

The regional facies architecture of the Campbellrand car-bonate platform has been documented by Beukes (1980c,1983, 1987), Sumner (1997a, b), Sumner and Grotzinger(2004), Schroeder et al. (2006), and Sumner and Beukes(2006). The platform comprises a 2,400-m-thick stromatoliticshallow-shelf succession that sharply thins along the Griqua-town fault zone, along the southwestern margin of the Kaap-vaal craton, into a 550-m-thick finely laminated, nonstroma-tolitic basinal (i.e., deeper shelf) carbonate succession (Fig.8). Deposition of the succession took place over a period ofabout 70 m.y. between 2.52 and 2.59 Ga and is constructed of12 third-order transgressive-regressive sequences, each withan average duration of ~6 m.y. (Sumner and Beukes, 2006).The lower part of the platform consists of a carbonate ramp(Beukes, 1987) comprising supratidal and intertidal chertydolostones that interfinger toward the ramp margin with shal-low-subtidal elongate giant stromatolite mounds (Fig. 8). Bas-inward of the ramp margin deep subtidal fenestrate micro-bialites, often with steep conophyton-like stromatolites(Beukes, 1987), are interbedded with muddy and grainy slopecarbonates (Schroeder et al., 2006; Sumner and Beukes,2006). With time, this ramp developed into a rimmed shelfmargin that characterizes the upper part of the succession


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Hydrothermal Plum

e

BASIN

VAALBARA

PilbaraCraton

0 km 500

Koegas

Newman

Post DepositionalUplift

SHELFNullagine

Post Depositional Uplift

Griquatown

Pomfret

KurumamJohannesburg

KaapvaalCratonVolcanic

CentreWeeli WolliTimes

PyroclasticsWittenoomTimes

N

FIG. 7. Reconstruction of depositional environments in the original com-bined Transvaal-Hamersley basin. Modified from the paleogeographic re-construction by De Kock (2007) based on paleomagnetic results. Position ofsource of pyroclastics during deposition of the Wittenoom carbonates is afterHassler (1993) and that of mafic volcanics in the Weeli Wolli Iron Formationafter Barley et al. (1997).


(Beukes, 1987; Sumner and Beukes, 2006). The rimmed shelfmargin is mainly constructed of high-relief columnar stroma-tolites and giant stromatolite mounds. Stratiform stroma-tolitic shelf lagoonal deposits are developed behind and slopedeposits in front of this rim. Steep conophyton-like stromato-lites are well developed in the upper part of the slope (Fig. 8)and mark highstands when very little carbonate detritus wasexported from the shallow shelf (Schroeder et al., 2006 Sum-ner and Beukes, 2006). Distal from the slope the carbonatesuccession is essentially composed of finely laminated fineturbiditic carbonate wackestones, thin grainstones, and strat-iform microbialites. These microbialites typically comprisevery organic-rich microbial laminae that are locally contortedor disrupted to form roll-up structures (Beukes, 1987; Kleinet al., 1987; Sumner, 1997a, b). Thin carbonaceous shale beds

are interbedded with both shelf and basinal carbonates; in-terestingly those of the basin appear more abundant along thetoe of slope than farther offshore (Schroeder et al, 2006). Sev-eral mafic tuff intervals are present in the deep-water basinalfacies but absent or not preserved among shallow-shelf car-bonates (Beukes, 1980c, 1987).

Within this carbonate platform setting, the ankerite-banded cherts and two thin BIF units are all interbeddedwith some of the most distal, deep-shelf carbonates in thebasin (Fig. 8; Beukes, 1980c, 1983, 1987). There are at leastsix ankerite-banded cherts in the most distal part of thebasin. Characteristically, most of them pinch out toward theshelf slope and only one unit, near the base of the succession,actually extends into the shelf slope environment (Fig. 8).Relative to the ankerite-banded cherts, the two BIF units


0361-0128/98/000/000-00 $6.00 17

ch

chch

chchch

chchch

chch

ch

RAMP

LAGOON

2000 m

1000

0

North

~ 450 km

Kuru

man

BIF

Fe

Fe

FeFeFeFeFeFeFeFe

FeFeFe

Fe

FeFeFeFeFe Fe Fe Fe Fe FeFe Fe Fe

Fe Fe Fe FeFe

FeFe

FeFe

Fe

South

FAULT

Deep Shelf

SHELFMARGIN

SLOPE

BrunosBIF

KamdenBIF

Kuruman BIF

BASIN SLOPE MARGIN CARBONATE SHELF INTERIOR

Monteville Establishment of Platform

Conophyton-like microbialites

Griquatown Growth Fault

RAMP

INTERTIDAL

INTERTIDAL

Koegas Griquatown Whitebank Pomfret

Intertidal dolomite =Cherty, light grey, highly depleted in FeO with MnO : FeO >> 1Lagoonal dolomite = Manganiferous with MnO : FeO > 1Ramp and shelf margin dolomite = Manganiferous with MnO : FeO > 1Slope and basinal dolomite = Ferruginous with MnO : FeO < 1

etuaNnielK

Gamohaan drowning of platform

Oolite Banks

FIG. 8. South to north stratigraphic section of the Neoproterozoic Cambellrand carbonate platform succession, illustrat-ing facies relationships between basinal and shallow-shelf environments and the stratigraphic setting of two persistent ironformation units known as the Kamden Member in middle and Brunos BIF near the top. Both iron formations are associatedwith major transgressions in the succession with Brunos BIF marking final drowning of the carbonate platform. The Kam-den BIF in the middle of the succession thickens from a few meters to almost 30 m along the slope of the shelf in front ofthe shelf margin. Deeper into the basin it is represented by ankerite-banded chert (proto iron formation). Note that ankerite-banded chert (proto iron formation) units are essentially restricted to deep basinal environments with one unit near the baseof the carbonate platform extending into the shelf environment to the north of the Griquatown growth fault that determinedthe position of the shelf margin for most of the time. This lower ankerite-banded chert was deposited during the initial trans-gression that led to establishment of the carbonate platform on the Kaapvaal craton (modified after Beukes, 1987). Refer toFigures 1 and 15B for position of section line.


(Kamden Member and Brunos BIF) have much wider lateraldistribution and extend from the basin well onto the shelf(Fig. 8). All of the iron-rich units mark transgressions or max-imum flooding surfaces in the basin (Beukes, 1987; Sumnerand Beukes, 2006). Transgressions associated with the twoBIF units (Kamden Member, Brunos Band, Fig. 8) were,however, most severe and led to drowning of the shallow-shelf carbonate platform. Drowning of the platform duringdeposition of the Kamden BIF in the middle of the succes-sion was short-lived, whereas transgression associated withBrunos BIF initiated the final demise of carbonate buidupand eventually led to deposition of the overlying very thickKuruman Iron Formation (Fig. 8).

In contrast to the iron-rich units, carbonaceous shale bedsin the basin appear to have been deposited at times of maxi-mum regression when the carbonate shelf was either exposedor produced very little carbonate so that clays from the plat-form interior could bypass the shelf and be deposited in thebasin. This would explain why shale beds are more prominentimmediately in front of the shelf slope than farther into thebasin (Schroeder et al., 2006). On the shallow-shelf, carbona-ceous shale beds often overly karstic carbonate exposure sur-faces. Here, they appear to mark initial phases of transgres-sion when newly created accommodation space led toretrogradation of siliciclastic muds into the shelf interior andrenewed carbonate production was established.

Carbonates, both limestone and dolomite, of the deep-basinand shallow-subtidal environments are generally dark and car-bonaceous, whereas those of the inter- to supratidal environ-ments are light gray in color. The limestones and dolomites areon average 10 to 100 times enriched in MnO and FeO relativeto Phanerozoic carbonates (Veizer, 1978; Beukes, 1987). Mostinterestingly dolomite and limestone of the Campbellrandsuccession display a distinct trend from being enriched in ironover manganese in the basin to just the opposite in shallow

sub- and inter- to supratidal carbonates (Fig. 8; Beukes, 1987).Light gray inter- and supratidal dolomites in the interior of theshelf have rather similar concentrations of MnO relative tocarbonates of other depositional environments but are highlydepleted in FeO (Beukes, 1987).

Carbonate to iron formation depositional systems tracts

Ankerite-banded chert: The ankerite-banded cherts in thebasinal facies of the Campbellrand carbonate platform suc-cession are composed of alternating bands of highly ferrugi-nous to ankeritic dolomite and chert. The highly ferruginousto ankeritic dolomite bands are regarded as ferruginizedequivalents of deep-water basinal limestone and dolostonethat are similar in texture (Beukes, 1984). Two types of chertbands are present, namely intraclastic chert, representingchertified deep-water muddy carbonate turbidites and mi-crobanded primary sedimentary chert bands (Beukes, 1984).The ankerite-banded cherts indicate that the deep-basinalcarbonate shelf was occasionally encroached by a water massthat must have been enriched in iron and silica to have al-lowed for precipitation of microbanded primary cherts andferruginization and chertification of surrounding deep-shelfcarbonates.

Kamden BIF: The Kamden BIF (Fig. 8) represents themaximum flooding stage of the sixth third-order depositionalsystems tract or sequence (Sumner and Beukes, 2006) in theCampbellrand carbonate platform succession (Fig. 9). TheBIF is sideritic and reaches a maximum known thickness of~30 m along the slope of the shelf margin from where it pe-ters out to thin ankerite-banded chert deeper into the basin(Fig. 8). On the shallow shelf it is represented by a persistentzone of ferruginous dolomite and chert (Fig. 8) associatedwith deep-water carbonates traceable for hundreds of kilo-meters along strike (Beukes, 1987; Sumner and Beukes,2006).


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B I F

DolareniteStromatoliteLaminated dolmicriteSiderite BIFLaminated dolmicrite

Conophytic dolomite

Giant stromatolitemounds

Carbonaceous shale anddolomite rip-ups

LSST

ETSST

TSSTLTSS

EHSSTLHSST

SB Exposure

Maximum Flooding

Stage 1=Exposure and lowstand wedge of shale

Stage 2 = Transgression

Stage 3 = Maximum flooding. Plume upwelling

Stage 4 = Late highstand

PL

NWB

Carbonate sands into basinPlume retreat

Hydrothermal Plume

Turbidite

SB Exposure(Clay bypassing)

~25

0m

CM6

Chert

BIF

++

Si0 + Fe2

Fe-dolomite

BIF

Shale

FIG. 9. Schematic illustration of the depositional systems tract that holds the Kamden Iron Formation in sequence 6 ofSumner and Beukes (2006) on the shelf of the Cambellrand carbonate platform. See text for details. Symbols in diagrams il-lustrating lateral facies relationships at different stages of development of the sequence are the same as that in the profile.EHSST = early highstand systems tract, ETSST = early transgressive systems tract, LHSST = late highstand systems tract,LSST = lowstand systems tract, LTSST = late transgressive systems tract, SB = sequence boundary, TSST = transgressivesystems tract.


Examination of the sequence hosting the Kamden Memberallows rconstruction of depositional systems tracts, commenc-ing with initial exposure of the carbonate platform, at thebasal sequence boundary, through flooding of the platformand BIF deposition to regression and development of anupper exposure surface or sequence boundary (Fig. 9). Onthe shallow shelf the lower sequence boundary is representedby a marked dissolution and erosional surface capping high-stand carbonates of the underlying sequence (Fig. 9). It is thisstage that would have been most favorable for bypassing ofthe carbonate shelf by siliciclastic clays and deposition in thebasin immediately in front of the shelf margin (Stage 1, Fig.9; Schroeder et al., 2006). Following exposure, during earlylowstand, slightly increased accommodation space led to de-position of carbonaceous shale with carbonate ripup clastsrepresenting the basal lithologic unit of the sequence on theshelf (Fig. 9). Subsequent transgression accompanied by in-creased accommodation space led to rapid carbonatebuildup, trapping and precipitation of carbonate in stromato-lites on the shelf, and very little carbonate bypassing the shelfmargin (Stage 2, Fig. 9). Steep conophyton-like stromatolites,essentially made up of precipitated carbonate cement andfree of transported carbonate mud and sands, grew along theshelf slope during this time. At greater water depth, thick or-ganic-rich microbial mats with contorted laminae and someroll-up structures and subordinate, thin, micritic turbiditebeds formed (Stage 2, Fig. 9). Because of abundant accom-modation space, the deep basin became starved of influx ofcarbonate mud and sand from the shallow shelf, creating anideal setting for deposition of ankerite-banded chert and/oriron formation.

Maximum flooding and drowning of the carbonate shelfwas followed by deposition of the Kamden BIF (Stage 3,Fig. 9). BIF deposition was most pronounced along theslope of the shallow carbonate platform, in an area thatwould have favored enhanced upwelling of deep oceanwater (Stage 3, Fig. 9). In the central part of the shallowplatform, influx of this deep ocean water is recorded by fer-ruginization and silicification of deep-water carbonates.This indicates that shallow-marine carbonate productionwas not completely shut off by the marine flooding eventbut must have persisted deep into the interior of the plat-form. This is corroborated by the fact that the KamdenMember pinches out toward the more interior part of thecarbonate platform (Fig. 8).

A progradational systems tract followed on deposition ofthe Kamden Member (Stage 4, Fig. 9). It implies retreat ofthe iron-enriched deep-water column from the shelf. Withdecreasing accommodation space, especially during late high-stand, carbonate was exported from the shallow shelf leadingto deposition of turbiditic carbonate in the basin. Shallowsub- to intertidal stromatolitic carbonates and cross-lami-nated dolarenites formed on the shelf immediately prior todevelopment of a subsequent exposure surface at the next fallinflection point (Stage 4, Fig. 9).

Brunos BIF: A transgressive systems tract similar to theone described above also led to deposition of Brunos BIFnear the top of the Gamohaan Formation in the uppermostpart of the Campbellrand-Malmani carbonate platform suc-cession (Fig. 8). However, in the latter case the carbonate

platform did not recover from the flooding event and deposi-tion of iron formation persisted into the overlying As-besheuwels Subgroup (Figs. 8, 10A).

The systems tract for Brunos BIF, sequence CM12 of Sum-ner and Beukes (2006), has been described in detail byBeukes (1980c, 1987) and Sumner (1997a, b). Here we pro-vide only a brief summary. The systems tract commences withlight gray cherty intertidal carbonates overlain by laminoidfenestral lagoonal limestone and shallow-subtidal elongatedgiant stromatolitic mounds that constructed the shelf margin(Fig. 10A). The shelf margin deposits are overlain by rippledmicrobial carbonate sands that were washed down from envi-ronments above normal wave base. These are, in turn, over-lain by a prominent zone of carbonaceous fenestral limestonecharacterized by a variety of conophyton-like stromatolites,also referred to as the conoform microbialite assemblage bySumner (1997a, b). The lower part of this zone consists of asuccession of cuspate small conophytic microbialites alternat-ing with thin layers of contorted microbial mats. Some of thecontorted mats appear to have been reworked by currents.Above the bedded cuspate microbialites there is no indicationof current reworking so that the top of the bedded cuspatesubzone is considered the absolute limit of the storm-wavebase (Fig. 10A). The bedded cuspate microbialites are suc-ceeded by a subzone with abundant conoform columnar stro-matolites set among highly carbonaceous stratiform to con-torted and rolled-up microbial mats (Fig.10A). A fewpersistent marker beds composed of delicate plumose micro-bialite structures (Sumner 1997a, b) are interbedded with theconoform stromatolite subzone. Conophyton-like micro-bialite structures characteristically start small and closelyspaced in lower parts of the succession and become taller andmore widely spaced upward, until they disappear abruptly.This is taken as an indication of increasing water depth, withthe disappearance of conophytic microbialites most probablymarking the photic limit (Fig. 10A). Above the conoform stro-matolite zone the succession is essentially composed of deep-water, highly carbonaceous stratiform microbialites withabundant contorted bedding, roll-up structures, and fram-boidal pyrite nodules (Fig. 10A). This facies is considered torepresent a proximal deep-shelf environment.

Farther up the sequence into distal deep-shelf strata, thecarbonates are essentially composed of carbonaceous strati-form microbialites with lesser contorted bedding and roll-upstructures but with more frequent thin black carbonaceousshale partings and some thin-graded low-density turbiditebeds made up of tiny fragments of dark carbonaceous micro-bial laminae (Fig. 10A; Beukes, 1987). This facies is in directcontact with ankerite-banded chert and microbanded bandedmicritic iron formation of Brunos BIF that marks the base ofthe Tsineng Member of the Gamohaan Formation (Fig. 10A).As the BIF unit is approached, nodular pyrite disappearsfrom the succession with only trace amounts of very fine dia-genetic pyrite present in carbonaceous carbonate and chertnext to banded micritic iron formation units. Banded micriticiron formation beds proper contain no evidence of any earlydiagenetic pyrite and the trace amounts of pyrite that are pre-sent can normally be seen to be either very late diagenetic orpostdiagenetic in age, having formed after lithification ofthe banded micritic iron formation. The banded micritic


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iron formation is typically very chert rich and sideritic imme-diately adjacent to the dark carbonaceous limestones. How-ever, at some localities the center of Brunos BIF may becomposed of hematite-facies banded micritic iron formation(Fig. 10A). Brunos BIF is usually overlain by a thin unit ofdeep-water carbonate that in turn grades upward into car-bonaceous shale that consistently underlies the KurumanIron Formation (Figs. 8, 10A).

Dolostones in the depositional systems tract that leads upto Brunos BIF (Fig. 10A) display compositional variations re-lated to depositional settings that are similar to those observed

on a basinal scale (Fig. 8). Dolomites in upper slope, shallow-shelf margin lagoonal and intertidal settings are enriched inmanganese over iron with the opposite true in lower slopeand deep-shelf dolomites. Dolostones immediately adjacentto ankerite-banded chert and banded micritic iron formationare highly ferruginous and even ankeritic to sideritic (Fig.10A; Klein and Beukes 1989).

It is important to note that in the basin proper, BrunosBIF is represented by a single, well-defined banded micriticiron formation unit. However, on the shelf it typically splitsinto several layers separated by carbonate and/or black shale


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Kuruman

BrunosBIF

Abundantlaminatedmat

Deep shelfmargin

Distaldeep shelf

Proximaldeep shelf

AnkeriticDolomite

Lowerslope

Upperslope

Shelfmargin

Shelflagoon

Mn dol.MinorLimestone

Limestone +Mn dol.Light greydolomite

Dysaerobic

DysaerobicAnaerobic

Aerobic

Sid

Sid

SidHem

PhoticLimit

Pyrit

eout

Py

Py

Koge

lbee

nFm Intertidal

flatsAerobic

Dys

aero

bic

Anae

robi

cD

ysa

ero

bic

GAM

OHA

ANFO

RMAT

ION

chPy

ch ch ch

Cherty

StormWB

NormalWB

MRA Matrix PyriteAnkerite-bandedchertContorted Mat Distal Deep shelf

Nodular limestone

Giant moundsRippled calcarenite

Shallow shelf assemblageBedded CuspatePlumose StructureConoform Columns

Conoform MicrobialiteAssemblage

Carbonaceouschert

Contorted and laminated mat

Siderite BMIFHematite BMIF

Carbonaceous shale

Laminoidfenestral limestoneSmall stromatolitesand calcarenite

Kuru

ma

nG

amoh

aan

BrunosBIF

Oxide Kuruman IF

CarbonaceousLaminaeConoform zone

(Kam

ber

and

Web

b20

01)

(Bau and Dulski,1996)

Contorted MatNodular Pyrite

Pyritic zone(Kamber and Whitehouse,2007)

LagoonalLimestone

Spar CementConoform Zone

LimestoneMinorFe dol.

MORHydrothermalFluidHematiteMagnetiteSideritePyrite

Pyrite fromRouxel et al. (2005)

Tuff2521Ma

Tsin

eng

Mem

ber

Hem

MagSid

Others fromJohnsonet al.(2003)

S Pool

Shaley ContortedMat; Deep shelf

REE

/PAA

S

-310

-210

-110

10

12

8

4

0

Eu

-15 -10 -5 0 5 10 15 20 25 30

La Ce

Y

H H

-2,0 -1,0 0 +1,0

RelictReservoir4

mSR

Klei

na

nd

Beuk

es(19

89)

ch

Sid =

Hem =

33s(%

)

34s(% )

La Ce Pr Nd Sn Eu Gd Tb Dy Ho ErTm LuYbY

56Fe

Legend for B Legend for AA B

C

Conoformmicrobialitezone

E

SO4

MRA12

8

4

-10 10 20 30

33s(%)

34s(% )

D

-2

Samples below TsinengMemberSamples of TsinengMember(Kaufman et al.,2007)

FIG. 10. A. Stratigraphic profile of the upper Kogelbeen to Gamohaan successions that led to drowning of the Camp-bellrand carbonate platform and deposition of Brunos BIF and overlying Kuruman Iron Formation in the area of the shelfto the north of the Griquatown fault zone near Whitebank (see locality in Fig. 8). Also indicated is an interpretation of de-positional and diagenetic environments and changes in the composition of dolomites from shallow-intertidal to deep-basinalsettings and the part of the succession that was studied by Klein and Beukes (1989). Modified after Beukes (1987) andSumner (1997a). ch = chert, dol = dolomite, Hem = hematite-facies BIF, py = pyrite, sid = siderite-facies BIF, WB = wavebase. B. Iron isotope composition of various mineral phases (expressed as 56Fe) in the succession, including samples fromthe overlying Kuruman Iron Formation. C. Secondary iron mass spectrometer multiple sulfur isotope analyses of differenttypes of pyrite in the succession (for location of samples see symbols, which are similar to those in crossplot of sulfur isotopevalues, immediately to left of stratigraphic profile). D. Whole-rock multiple iron isotope analyses of pyrite in the succession.The samples come from one of the drill cores that Klein and Beukes (1989) investigated. E. Examples of ICP-MS REE analy-ses of different lithofacies in the succession, including shallow-water lagoonal limestone. References to the different data setsare given in the figures.


composing the Tsineng Member of the Gamohaan Formation(Fig. 10A; Klein and Beukes, 1989). This suggests that thetransgression of Brunos BIF from deep basin onto the shelftook place in steplike fashion through time and that it doesnot necessarily represent a synchronous time marker in thebasin but rather a protracted time interval in which the trans-gression took place.

Paleoenvironmental reconstruction

The detailed lithostratigraphic data provided above in com-bination with geochemical information allow reconstructionof a depositional model for iron formation in relationship tomarine carbonate sedimentation, primary organic carbon pro-ductivity and paleobathymetry in an oceanic basin toward theend of the Neoarchean. Certain aspects of this model are sim-ilar to that proposed by Klein and Beukes (1989; Fig. 11). Thelatter model envisaged a depositional system in which ironformation was deposited from a chemocline along which hy-drothermally enriched deep water came into contact with theshallow well-mixed surface layer of a more or less perma-nently stratified ocean water column (Fig. 11). Deposition ofiron formation specifically took place during transgressionswhen the carbonate platform became drowned and organiccarbon input into the deep shelf from adjacent highly pro-ductive stromatolitic shallow-shelf settings was at minimum(Fig. 11). However, the Klein and Beukes (1989) model re-quires substantial revision based on the availability of newgeochemical data and concepts. These are summarized belowbefore a revised depositional model is developed.

Rare earth element (REE) data: The iron formations andclosely related ankerite-banded cherts are invariably in-terbedded with the most distal and deepest basinal shelf fa-cies of the carbonate platform succession (Figs. 810A).However, for most of the time, only deep-water limestonewith subordinate carbonaceous shale beds accumulated onthe deep shelf in front of the shallow-carbonate platform mar-gin (Fig. 8). This implies that a more distal or deeper watermass, which must have been enriched in silica and iron, onlyoccasionally transgressed onto the deep-carbonate shelf and

resulted in silification and ferruginization of deep-water lime-stone to form ankerite-banded chert (Beukes, 1983) and/ordeposition of banded micritic iron formation (Fig. 9). From asedimentologic perspective, it would thus appear that twowater masses or columns with different composition, a shal-lower one from which limestone and shale were deposited,and a deeper one from which silica and iron were deposited,interacted in the most distal deep-carbonate shelf environ-ment. This notion is supported by the fact that limestones andassociated carbonaceous shales have geochemical signaturesthat are distinct from those of the more distal ankerite-banded chert and banded micritic iron formation, as was firstnoted by Klein and Beukes (1989). Limestones display ratherflat shale-normalized REE patterns similar to that of modernshallow-marine surface waters with detrital riverine influxwhile iron formations display LREE-depleted patterns simi-lar to that of modern deep-marine water with no terrigenousinput (Klein and Beukes, 1989). The concept finds furthersupport in Al2O3 concentrations that reflect the amount of de-trital alumosilicates (i.e., clay minerals). Very low Al2O3 con-tents indicate that iron formations were deposited in an envi-ronment devoid of siliciclastic detrital input as opposed todeep-shelf limestones that have distinctly greater Al2O3 con-tents (Fig. 12A).

In contrast to modern seawater, shale-normalized REEpatterns of both deep-shelf iron formation and limestone inthe Gamohaan-Kuruman transition zone (Fig. 10A) displaydistinctive positive Eu anomalies (Klein and Beukes, 1989).It is generally accepted that the Eu anomalies were derivedfrom a component of hydrothermal fluids derived from high-temperature (>250oC) alteration of ocean-floor basalts (Bauand Dulski, 1996); the latter are highly enriched in REE withvery marked positive Eu anomalies (Michard et al., 1983).The limestone samples analyzed in the study of Klein andBeukes (1989) were from the distal deep-shelf environmentof the Gamohaan Formation (Fig. 10A), i.e., the REE char-acteristics of shallow-shelf carbonates (and thus the shallowocean-water reservoir) remained unknown. More recently,Kamber and Webb (2001) produced REE data from lower


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Blackchert

or shale

cc

O2O2

2+Fe2+Fe2+Fe

Hydrothermal input

Oxide BIF SideriteBIF

O2 Carbonate shelf

C=Organic Carbon

Stratified ocean system. Production of organic matter low in open marine environment but high on carbonate shelf.

O2

C

FIG. 11. Diagram illustrating the main components of the model that Klein and Beukes (1989) developed for depositionof iron formation in the transition zone between the Campbellrand carbonate platform and the Kuruman Iron Formationsuccessions. This model implies a more or less permanently stratified ocean with deposition of iron minerals taking placealong a chemocline situated at the base of the surface mixed zone of the ocean water column. It also indicates siderite to bea primary preciptate in response to increased pCO2 that developed in water column due to degradation of organic matter inareas of high-organic carbon input from primary production that took place through photosynthesis in shallow-carbonateshelf environments. In open ocean environments little primary production of organic matter took place due to scarcity in nu-trients and thus oxide-facies iron formation formed through reaction of dissolved ferrous iron, derived from upwelling deephydrothermally enriched water, with free oxygen that was available in shallow-surface ocean water. Deposition of iron min-erals and biological activity is totally decoupled in this model.


slope conophytic microbialites and laminoid fenestral la-goonal limestone in the Gamohaan succession (Fig. 10E).They showed that the positive Eu anomaly persists even intothe lagoonal deposits, although it is progressively less pro-nounced into shallower w

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