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Trace element distribution in Neoproterozoic carbonatesas palaeoenvironmental indicator

Hartwig E. Frimmel ⁎Department of Geological Sciences, University of Cape Town, Rondebosch 7701, South Africa

a b s t r a c ta r t i c l e i n f o

Article history:Received 28 May 2008Received in revised form 17 October 2008Accepted 25 October 2008

Editor: D. Rickard

Keywords:Rare earth elementsCarbonateNeoproterozoicGeochemistrySouthern Africa

A first study of REE+Y distribution in a variety of Neoproterozoic (Cryogenian and Ediacaran) carbonates fromdifferent settings in the Saldania, Gariep, Damara and West Congo Belts in southwestern and central Africarevealed systematic differences that can be explained by varying palaeoenvironmental factors. Themajority ofsamples display relatively unfractionated, flat shale-normalised REE+Ypatterns that cannot be ascribed solelyto shale contamination but are interpreted as resulting from the incorporation of near-shore colloids, possiblyrelated to Fe-oxihydroxide scavenging. Only few carbonate units yielded trace element distributions thatconform to a typical seawater composition. Those carbonates that were affected by stratiform, syn-sedimentary hydrothermal mineralisation are distinguished by Eu anomalies. Considering the similarity inresidence time between REE and carbon, the strong influence of river-born particles on the REE+Y distributionin the analysed carbonates casts considerable doubt over the usefulness of these carbonates for stratigraphiccorrelation of Neoproterozoic sediment successions based on carbon isotopes.

© 2008 Elsevier B.V. All rights reserved.

1. Introduction

A number of recent studies have shown that chemical sedimentaryrocks, such as some carbonates, banded iron formation/cherts andphosphates, can serve as useful proxies for the record of certain traceelement patterns in the water from which these rocks originate.Systematic differences in the properties of the lanthanide series (REE)and Y make it possible to use them for differentiating betweendifferent types ofmineral-precipitatingwaters (e.g. Bolhar et al., 2004;Nothdurft et al., 2004; Bolhar and Van Kranendonk, 2007). Thedistribution of these elements is very sensitive to water depth, salinityand oxygen level. At the same time the REE+Y distribution monitorsalso differences in the input sources, mainly the ratio of continentalinput via rivers and airborne dust versus oceanic hydrothermal input.Marine chemical sediments typically reflect a seawater REE+Ydistribution that appears to be independent of age (e.g., Shields andWebb, 2004; Bolhar and Van Kranendonk, 2007). They are charac-terised bya uniform light REE depletion, enrichment in La, depletion inCe, slight enrichment in Gd and marked positive Y anomaly in shale-normalised diagrams (Zhang and Nozaki, 1996). The extent of the Cedeficiency is related to oxygen level, with oxidised Ce4+ being lesssoluble and thus more readily adsorbed onto particles (De Baar et al.,1991; Möller et al., 1994; Alibo and Nozaki, 1999). Not surprisingly, the

negative Ce anomaly is absent in carbonates and banded ironformation of Archaean to earliest Proterozoic age when seawater wasnot sufficiently oxidising to form CeIV (Bau and Dulski, 1996; Kamberand Webb, 2001). Very different REE+Y patterns have been obtainedfrom reducing, acidic hydrothermal fluids, which typically display adistinct positive Eu anomaly in otherwise uniform, light tomiddle REE-enriched patterns when normalised to a shale composition. This isevident, for example, from analyses of hydrothermal discharge on theseafloor (Michard et al.,1983; Bau andDulski,1999;Wheat et al., 2002).In contrast, river water is characterised by relatively flat REE+Ypatterns with slight uniform light REE depletion and no distinctelement anomalies (Goldstein and Jacobsen, 1988; Lawrence et al.,2006; García et al., 2007), which has been used successfully todistinguish betweenmarine and lacustrine carbonates (Bolhar andVanKranendonk, 2007).

The overall REE concentration in most waters, including seawater,and their respective precipitates, is generally very low and close to thelower limit of detection of methods traditionally employed to analysefor these elements in the past. The increasing availability of inductivelycoupled mass spectrometric (ICPMS) methods, which have a suffi-ciently low detection limit for REE, and the distinct differences in thetrace element geochemistry between different types ofwater and theirrespective precipitates offer an opportunity to gain new insights intolikely ancient palaeoenvironmental conditions through the analyses ofREE+Y distributions in chemical sedimentary rocks.

The Neoproterozoic Era is famous for dramatic climatic changesthat are recorded by glacial and intercalatedwarm-water deposits. Thecarbonate successions that typically overlie glaciogenic diamictite

Chemical Geology 258 (2009) 338–353

⁎ Geodynamics & Geomaterials Research Division, Institute of Geography, University ofWürzburg,AmHubland,D-97074Würzburg, Germany. Tel.: +49 9318885420; fax: +49 9318884620.

E-mail address: hartwig.frimmel@uni-wuerzburg.de.

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deposits (cap carbonates) are, inter alia, critical for the palaeoclimaticinterpretation of the Neoproterozoic rock record. One of the under-lying assumptions for the palaeoclimatemodels is that the informationobtained from chemical sediments of the post-glacial time slices isindeed representative of the global ocean water chemistry. Ourunderstanding of these carbonate deposits is, however, still verylimited, in parts because of severe problems in extrapolating theseawater composition back in time beyond the Phanerozoic Eon, inparts because of poor palaeogeographical and thus palaeoenviron-mental control. A fundamental problem in our understanding ofNeoproterozoic glaciations lies in the difficulty of correlation. In theabsence of precisely datable units in many Neoproterozoic glacial andpost-glacial successions, C isotope chemostratigraphy has become apopular method to correlate stratigraphic units on both a regional andglobal scale (e.g. Halverson et al., 2005).

Microbial carbonates have been shown to yield some of the mostreliable proxy data for the REE+Y distributions in ancient seawater(e.g. Kamber andWebb, 2001). Such microbialites are abundant in thevarious Neoproterozoic carbonate successions and they form, there-fore, the main focus of this study. Here results are reported of traceelement, including REE, distributions in carbonate samples from thePan-African Gariep, Saldania, Damara and West Congo Belts insouthwestern and central Africa (Fig. 1) in order to test the usefulnessof these trace elements for better assessing the palaeoenvironmentalconditions during carbonate formation. The studied examples ofcarbonate rocks were selected in such a way as to cover a variety of

Fig. 1. Locations of the studied carbonate samples (circles) in the various Pan-Africanorogenic belts in southwestern Africa, shown in an Upper Cretaceaous Gondwana break-up position.

Fig. 2. Stratigraphic correlation of the studied carbonate units across thePan-African belts of southwestern and central Africa; sourceof age data: 1— Tack et al. (2001), 2— Junget al. (2007),3 — Hoffman et al. (1996), 4 — Frimmel et al. (2001b), 5 — Frimmel et al. (1996b), 6 — Fölling et al. (2000), 7 — Grotzinger et al. (1995), 8 — Barnett et al. (1997), 9 — Naidoo (2008).

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stratigraphic positions (i.e. Cryogenian versus Ediacaran) and deposi-tional environments as known from regional geological and previousisotopic studies. Thus the following examples were included for thisstudy: (i) Cryogenian carbonate rocks of the c. 750 Ma Rosh PinahFormation in the Gariep Belt in southwestern Namibia, a post-glacialcarbonate that hosts syn-sedimentary hydrothermal mineralisation;(ii) Ediacaran examples of pre-glacial limestone of the Dabie RiverFormation and the post-glacial Bloeddrif Member, which displays allthe features that are typical of a post-glacial (“Snowball Earth”) capcarbonate in the same belt; (iii) limestones from the Cango CavesGroup of the Saldania Belt in southern South Africa, representing theprobably Cryogenian Nooitgedagt Member and the younger KombuisMember, which has been interpreted as a stratigraphic equivalent ofthe Bloeddrif Member; (iv) carbonates from the Otavi platform innorthern Namibia, which include both an older post-glacial carbonate,possibly related to the aftermath of the Sturtian glaciation, and ayounger post-glacial succession above a syn-Marinoan (c. 636 Ma)glacial unit; and (v) carbonates from the West Congolian Group in theDemocratic Republic of Congo, in which again carbonates above anolder and a younger glaciogenic unit can be distinguished.

2. Geological settings

The choice of samples investigated was dictated by the desire toobtain data from Neoproterozoic carbonates of different age, differentdepositional setting (proximal versus distal) and different palaeogeo-graphic position. Thus, the analysed samples come from a variety ofNeoproterozoic units from awide area spanning from the southern tipof Africa to Central Africa (Figs. 1, 2). Although some of the details inthe correlations shown in Fig. 2 may be contentious, both theCryogenian and Ediacaran Periods are covered by the chosen samples.

2.1. Cryogenian carbonates in the Gariep Belt

The Rosh Pinah Formation forms part of the Hilda Subgroup in thePort Nolloth Group, which constitutes the Neoproterozoic strata in theexternal part of the Pan-African Gariep Belt, the Port Nolloth Zone, insouthwestern Namibia and western South Africa (Fig. 3). This part ofthe belt, although intensely deformed and metamorphosed atgreenschist-facies conditions, still rests on its original Palaeo- toMesoproterozoic basement. The Neoproterozoic sediment successionstarts with continental siliciclastic rift deposits, the StinkfonteinSubgroup (Lekkersing and Vredefontein Formations). First marinesediments appear in the uppermost part of the subgroup as minorintercalated dolostone and limestone. Flooding of the rift shoulderswas interrupted during a first glacial period, which resulted in thedeposition of the largely diamictitic Kaigas Formation (Fig. 4). Rift-related, mainly felsic volcanism occurred at the eastern basin margintowards the end of the glaciation. A horst-and-graben structure thatdeveloped during rifting was accentuated by a syn-glacial eustatic sealevel fall during which an eastern failed graben, the Rosh PinahGraben, became separated from a western half-graben. The lattereventually evolved into the Gariep Basin proper (Frimmel andJonasson, 2003). While the volcanic activity led to the localaccumulation of a thick volcanic to volcaniclastic succession, thebackground sedimentation further away from the volcanic centre(s)took the form of relatively thin carbonaceous argillite beds in the lessventilated parts of the Rosh Pinah Basin and of thinly laminated, partlyallodapic and variably argilleceous limestone (Pickelhaube Formation)elsewhere. A series of arenitic intercalations, largely arkose andfelspathic sandstone, in all of these facies reflect the overall proximityto the palaeoshore line, defined by a basement high whose outline isapproximated by the current margin of the orogenic belt.

The Rosh Pinah Formation contains, apart from felsic volcanic tovolcaniclastic and arenitic rocks, intercalated carbonate beds that aregenerally only several decimetres thin. Oxygen and carbon isotopic

evidence points to dolomitisation of original limestone by syngeneticto early diagenetic hydrothermalfluids (Frimmel and Lane, 2005) someofwhich formedmassive, broadly synsedimentary, stratiform sulphideore bodies (Rosh Pinah Pb–Zn–Cu-sulfide deposit and protore of thesecondary Scorpion Zn-deposit). The carbonate beds are in placesmicrobialites. Elsewhere large slump breccias occur that indicatedeposition in a seismically active region, probably the principal growthfault along the eastern margin of the Rosh Pinah Graben.

Deposition of the Rosh Pinah Formation and the largely coevalPickelhaube Formation was terminated by a hiatus of poorlyconstrained duration. The overlying stratigraphic unit is the Wallek-raal Formation. It consists mainly of conglomerates, dolomite breccias,

Fig. 3. Positions of the studied carbonate sample localities for the Rosh Pinah Formation(1), Dabie River Formation (2), and Bloeddrif Member (proximal — 3, distal — 4) in thePort Nolloth Zone of the Gariep Belt.

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intercalated arenites (immature greywacke and arkose) and argillite.The contact with the underlying strata is a regionally extensiveerosional unconformity, with the Wallekraal Formation sedimentscutting through the entire older stratigraphy down to the pre-Gariepbasement in places (Fig. 4). It is speculated that this erosion surfaceformed during the global Marinoan glaciation.

The maximum age of rift sedimentation is given by the youngestage obtained on a basement rock, a U–Pb single zircon age of 771±6Mafor the Lekkersing granite (Frimmel et al., 2001b). U–Pb and Pb–Pbsingle zircon age data of 752±6 (Borg et al., 2003) and 741±6 Ma(Frimmel et al., 1996b), respectively, were obtained on Rosh PinahFormation felsic volcanic rocks and they provide a minimum age forthe Kaigas Formation diamictite. A minimum age for the Pickelhaubeand contemporaneous Rosh Pinah Formation carbonate rocks is givenbya double-spike Pb–Pb carbonate datumof 728±32Ma obtained on alimestone bed in the lower Pickelhaube Formation (Fölling et al.,2000), which has been interpreted to date early diagenesis.

2.2. Ediacaran carbonates in the Gariep Belt

A younger glaciogenic diamictite (Numees Formation) within thePort Nolloth Zone is both underlain and overlain by carbonates (Fig. 4).The age of this unit is controversial. Correlation with the globalMarinoan (c. 636 Ma) glaciation has been suggested (Frimmel et al.,2002) because of strong similarities in the C isotopic record andlithofacies of the bounding carbonate units. Several arguments speak,however, for a younger, Ediacaran age. These include relatively highnear-primary Sr isotope ratios consistently between 0.7080 and 0.7085(Fölling and Frimmel, 2002) and a double-spike Pb–Pb carbonate ageof555±28 Ma for the overlying cap carbonate of the Bloeddrif Member(Fölling et al., 2000). More recently, micropalaeontological evidenceprovided further support for such an age (Gaucher et al., 2005) andcorrelation with the approximately 582 Ma Gaskiers (or possibly theslightly younger Moelv) glaciation is therefore favoured.

The reef facies carbonates of the Dabie River Formation beneaththe Numees Formation diamictite contain a marked positive δ13Canomaly but are progressively depleted in 13C with proximity to thecontact with the Numees Formation (Fig. 4). This reversal in the C

isotopic trend towards lower δ13C ratios has been interpreted assignalling a change in global climate in preparation of the Numeesglaciation (Fölling and Frimmel, 2002). No significant hiatus istherefore suspected between the Dabie River and Numees Formations.

The Dabie River Formation, which attains a maximum thickness of160 m, is lithologically distinguished from the other carbonate-bearing successions of the Hilda Subgroup by the presence ofstromatolites displaying Conophyton-like forms, several centimetresto decimetres in height. Pisolites, oolites and oncolites are alsopresent. The formation is almost exclusively calcareous, with originallimestone variably dolomitised. The carbonate rocks are typicallymassive, light to medium grey and, in places, intensely brecciated.Some of the carbonate breccias are interpreted as debris flow deposits,whereas others are ascribed to gravitational slumping. Cyclicalemergence and submergence is indicated by desiccation cracks andby the interbedding of limestone and dolostone. A shallow-water,rimmed shelf environment, such as a barrier bar or shelf lagoon,passing seaward (westward) into a shelf margin, comprising reefbuild-ups and oolitic to pisolitic shoals is envisaged for the deposi-tional environment. Reef rocks formed, or are preserved, particularlyin those areas that escaped the pre-Wallekraal erosion. Consequently,the Dabie River Formation carbonates rest in many places para-conformably above dolostone of the Pickelhaube Formation and notnecessarily on top of the clastic Wallekraal Formation rocks.

The Bloeddrif Member (lower Holgat Formation), which attains athickness of 100m but thins out to only a fewmetres along the easternmargin of the Port Nolloth Zone, is light grey, cream or pale pink incolour and poor in organic matter. It displays the characteristics of atypical cap carbonate, i.e., vertical tube-like structures of infilledmicritic sediment and cement. They are usually a few centimetresacross and several decimetres high. Similar structures have beendescribed from many other Neoproterozoic post-glacial cap carbo-nates, including the post-Ghaub Keilberg Member (Maieberg Forma-tion) in the Otavi platform (Hoffman and Halverson, 2008). TheBloeddrif Member consists mainly of clean, thinly laminated lime-stone, but dolostone is abundant particularly in the basal and proximalsections (Frimmel and Fölling, 2004). The limestone is distinguishedby very high Sr contents, reaching several thousand parts per million,

Fig. 4. Stratigraphic subdivision of the Port Nolloth Group in the Port Nolloth Zone, external Gariep Belt; C isotope profiles fromFölling and Frimmel (2002) and Frimmel and Fölling (2004).

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from which an aragonite precursor has been inferred (Fölling andFrimmel, 2002). The calcareous Bloeddrif Member is followed by asiliciclastic metasedimentary succession (upper Holgat Formation) thatcomprises upwards-fining cycles of medium-bedded sandstone, grey-wacke and arkose with minor siltstone, mudstone and intraformationalconglomerate. These rocks are interpreted to represent turbidites thatwere laid down in a foredeep (Frimmel and Fölling, 2004).

2.3. Carbonates of the Cango Caves Group in the Saldania Belt

Neoproterozoic rocks are exposed in a number of anticlinalerosional windows within the Permotriassic Cape Fold Belt in thesouthwestern and southern tip of South Africa. The pre-Cape rocksuccessions form the basement of the Palaeozoic Cape Supergroup andwere deformed into folds and thrusts in the course of a latestNeoproterozoic to Cambrian accretionary orogeny that led to theformation of the Saldania Belt. Carbonate rocks occur at severallocalities, but in most cases their stratigraphic position is unclearbecause of poor exposure of stratigraphic contacts. Reasonablestratigraphic control exists only for carbonates in the Kango Inlier inthe southern branch of the Saldania Belt (Fig. 5). There, theNeoproterozoic rocks are grouped together as Cango Caves Group,which is predominantly a carbonate-clastic turbidite succession. Thegroup is subdivided into the Matjies River Formation, GroenefonteinFormation and Huis Rivier Formation (Le Roux and Gresse, 1983). Ofthese, only the lower, former formation contains carbonates in theNooitgedagt Member and the Kombuis Member (Fig. 2). TheNooitgedagt Member consists of shale, greywacke and limestone thatrepresent a coarsening upward deltaic to shallow marine succession,whereas the latter is a predominantly calcareous shelf deposit.Although both members have been grouped together into the sameformation (Le Roux and Gresse, 1983) and considered to represent acontinuous sedimentary succession, marked differences in C and Srisotope chemostratigraphy (Fölling and Frimmel, 2002) and adifference of 100° C in thermal overprint between the Nooitgedagtand Kombuis Members (Frimmel et al., 2001a) point to a major hiatusbetween the two. A distinct positive δ13C excursion (to a maximum of+10‰) and relatively low 87Sr/86Sr ratios (as low as 0.7074 in Sr-richlimestone) in the Nooitgedagt Member carbonates compare well withsimilar data for other Cryogenian, post-Kaigas carbonates in the GariepBelt, whereas consistently lower δ13C (between +2 and −4‰) andsignificantly higher 87Sr/86Sr ratios from0.7080 to 0.7087 in the Sr-rich(Rb/Srb0.0001) Kombuis Member limestone are comparable to thoseof the Bloeddrif Member. The latter correlation is further supported byan identical double-spike Pb–Pb carbonate age of 553±30Ma obtained

for the Kombuis Member limestone (Fölling et al., 2000) and a verysimilar microfossil assemblage (Gaucher and Germs, 2006).

2.4. Carbonates in the Otavi Platform (Otavi Group)

The Otavi Platform in the Northern Foreland of the Damara Orogenin northern Namibia formed along the southern fringe of the CongoCraton and abuts against the continental slope facies further southandwest. Thus its predominantly calcareous sedimentary successions,unified as Otavi Group, are located in a foreland position relative to theKaoko Belt in the west and the Damara Belt in the south (Fig. 1). TheOtavi Group is subdivided into three subgroups, which are separatedfrom each other by two glaciogenic diamictite units, the lower ChuosFormation and the upper Ghaub Formation (Fig. 6). First drowning ofthe continental rift shoulders is evident in carbonates of the lowestsubgroup, the Ombombo Subgroup. These overlie and interfingerfelsic, rift-related volcanic rocks (Naauwpoort Formation). Severalprecise U–Pb zircon and titanite ages between 759 and 746 Maconstrain the period of magmatism (Halverson et al., 2005; Hoffman

Fig. 5. Simplified geology of the Saldania Belt with location of the carbonate samples from the Cango Caves Group (arrow).

Fig. 6. Schematic cross-section through the Otavi platform (modified after Hoffman andHalverson, 2008). For sources of the indicated U–Pb single zircon ages see text.

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et al., 1996; Jung et al., 2007) with the latter setting a maximumconstraint on the timing of the Chuos glaciation.

Different stratigraphic schemes have been used for the AbenabSubgroup above the diamictitic Chuos Formation in the west (EasternKaoko Zone) and the east of the platform (Otavi Mountainland). As thesamples for this study come from the latter region, only thestratigraphic subdivision as used for the Otavi Mountainland will besummarised here. Dark grey to black, laminated dolostone or lime-stone rhythmite that grades up from abiotic laminate, not more than15 m in thickness, into sublittoral microbialaminite follows above theChuos Formation. It forms the Berg Aukas Formation. A firstgeneration of largely stratiform Pb–Zn sulphide mineralisationaffected the formation and evidences late-rift hydrothermal activityin the basin (Pirajno and Joubert, 1993; Frimmel et al., 1996a).Overlying massive to bedded dolostone constitutes the GaussFormation, which is followed, in turn, by stromatolite and oolite,alternating with thin-bedded limestone and argillite beds of the AurosFormation. Fault-controlled rapid changes in thickness mark the lattertwo formations, but the original basin configuration at that stage is notconstrained in sufficient detail.

Lithology and carbon isotopes presage the impending glaciationand the accompanying sea-level fall that eventually resulted in theyounger glacial unit, the Ghaub Formation. The δ13C ratios range from+1 to +8‰ throughmost of the Abenab Subgroup but decrease to −6‰in the uppermost part (Halverson et al., 2005). The predominantlydiamictitic Ghaub Formation reaches as much as 2000 m in thicknessin the western Otavi Mountainland but is laterally discontinuous.Based on a proposed correlation with a diamictite in the centralDamara Belt for which Hoffmann et al. (2004) obtained a precise U–Pbzircon age of 635.5±1 Ma for an intercalated ash bed, a syn-Marinoanage has been postulated for the Ghaub Formation.

The cap carbonate to the Ghaub Formation glaciogenic rocks arerepresented by the tan to pinkish or pale grey, up to 40m thick KeilbergMember of the Maieberg Formation (Hoffman and Halverson, 2008;Hoffmann and Prave,1996). In themost instructive exposures, the basalmeter is composed of recrystallised and cemented dolomite siltite orgrainstone laminated by low-angle, metre-scale cross-beds. Above this,narrow, vertical, convex-up, stromatolitic columns about 2 to 5 cmacross and with a similar spacing are characteristically developed inmany places. Themicrobial lamination of the columns is usually poorlyvisible. A laminated, concave-up dolomicrite infill between thecolumns is generally far more apparent than the stromatolitic columnsand appears as evenly spaced ‘tubes’ in outcrop. The stromatolites areoverlain by a zone of giant wave ripples in peloidal dolomite withindividual ripple sets reaching 150 cm in thickness and crestal spacingof 1 to 1.5 m (Allen and Hoffman, 2005). Locally, vertically standingcrystal fans of calcite pseudomorphs after sea-floor aragonite aredeveloped just above the Keilberg Member. Overall, the MaiebergFormation, which reaches a thickness of 1800 m in the OtaviMountainland, forms a single, thick and extensive depositionalsequence, initially transgressive and thengradually upward shallowing.

The Elandshoek Formation (up to 1500 m thick) above theMaieberg Formation is made up of cherty, grainstone-dominated,dolomite. Several metres thick depositional cycles contain micro-bialaminites and are terminated by flooding surfaces. Columnarstromatolites are common in the upper half of the formation, whichis better bedded. Bedding-parallel silicification is a common feature.

The overlying Hüttenberg Formation begins with up to 1000 mthick, light to medium grey, bedded, in places stromatolitic dolomitewith numerous chert layers. This is followed by 290 m thick light anddark grey grainstone and mudstone beds with black carbonaceouslimestone intercalations and black chert layers. Above that followmassive to bedded and cyclically graded dolomitic grainstone andmudstone layers with interbedded silicified oolite beds and columnarstromatolites, altogether reaching 300 m in thickness. Evidence of arestricted basin with elevated evaporation rates exists near Tsumeb.

There stratiform anhydrite and gypsum were intersected in explora-tion drill holes through the upper Hüttenberg Formation. Further-more, a distinct stratiform breccia zone, the so-called North BreakZone, in the Tsumeb mine has been re-interpreted as a palaeo-aquiferconfined to a former evaporite bed, and fluid inclusion leachate dataindicate post-diagenetic circulation of evaporitic residual brines(Chetty and Frimmel, 2000).

2.5. Carbonates of the West Congolian Group

The Pan-AfricanWest Congo Belt that stretches from southwesternGabon across western Congo-Brazzaville and the westernmostDemocratic Republic of Congo into northwestern Angola containsNeoproterozoic carbonates within the West Congolian Group (Fig. 2).The group is particularly well developed in the foreland basin to theeast of the belt where it is only gently folded (Fig. 7) and hardlymetamorphosed. For a detailed description of the lithostratigraphy ofthe group see Cahen (1978). Siliciclastic continental rift deposits makeup the Sansikwa Subgroup at the base of the West Congolian Groupand are overlain by an older glaciogenic diamictite unit, the LowerMixtite Formation. This is followed by a varied succession ofconglomerate, argillite, calcpelite, quartz arenite, calcarenite andlimestone (Haut Shiloango Subgroup) and eventually a seconddiamictite, the Upper Mixtite Formation. The latter is overlain by acap carbonate sequence that develops into carbonate ramp andplatform deposits with abundant stromatolite bioherms and thefilamentous cyanobacterium Obruchevella (Alvarez et al., 1995). Thisunit is known as the Schisto-Calcaire Subgroup whose sequencestratigraphy has been described in greater detail by Alvarez (1995).

Eight different lithofacies have been recognised within the HautShiloango Subgroup (Cahen, 1978): Conglomerate and quartzite at thebase (Sh1), argillite (Sh2), partly calcareous argillite (Sh3), quartziteand argillaceous limestone (Sh4), and quartz phyllite (Sh5) comprisethe Little Bembezi Formation (450 to 650 m in thickness). This isfollowed by the 200 to 250m thick Sekelolo Formation, which consistsof feldspathic quartzite (Sh6), argillite (Sh7), and amixed succession ofargillite, limestone (partly nodular with intercalated calcarenite),stromatolites and calcareous breccias (Sh8). Similarly, the Schisto-Calcaire Subgroup is subdivided into five lithofacies: a post-glacial capdolomite (C1), calcareous argillite and quartz arenite (C2), and partlyoolitic limestone (C3) make up the Kwilu Formation (approximately500 m in thickness). Lithofacies C3 is characterised by very high Srconcentrations (N5000 ppm; Frimmel et al., 2006). Considering thelocal abundance of diagenetic albite in this unit, evaporitic anhydritehas been inferred as likely precursor mineral. This is also supported bythe presence of diagenetic celestite, polyhalite and clinochlore(Delpomdor, 2007). Above C3 follow limestone and dolostone, partlystromatolitic, with minor intercalated calcpelite and chert of theLukunga Formation (C4). Lithofacies C5 (Bangu Formation, up to270 m in thickness) comprises variably dolomitic and partly verycarbonaceous dark limestone, abundant oolite, minor stromatolitesand chert, with intercalations of calcpelite and talc schist. Obruche-vella is present in silicified oolite. For a detailed description of thevarious stromatolites present in theWest Congolian Group carbonatessee Bertrand-Sarfati (1972).

Near-primary 87Sr/86Sr ratios consistently around 0.70715 wereobtained for theHaut ShiloangoSubgroup limestoneand these support aCryogenian, possibly post-Sturtian age (Frimmel et al., 2006). High δ13Cvalues of asmuch as +8‰obtained in the same studyare consistentwithsuch a correlation. In contrast, limestonewith very low Rb/Sr (bb0.003)of the C3, C4 and C5 lithofacies in the Schisto-Calcaire Subgroup hasnear-primary 87Sr/86Sr ratios that are slightly higher (0.70740–0.70753;Frimmel et al., 2006) and that correspond to those found elsewhere inpost-Marinoan carbonates (Halverson et al., 2007). The δ13C ratios arenegative throughout the C3 limestones with values around −1‰ but arehighly variable and erratic, without noticeable trends, in the C4 and

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especially the C5 carbonates (range between −3 and +9‰). The lattervariability has been explained by differences in the depositionalenvironment, more specifically by temporarily elevated evaporationrates. Consequently, C isotopes appear not suitable for chemostrati-graphic correlation of these units.

3. Samples and analytical techniques

A total of 148 carbonate samples were selected for geochemicalanalysis. The rock samples (c. 0.5–2 kg) were reduced in size by a steelpress and c. 100 g of alteration-free rock chips were handpicked andpulverised (300–400 mesh). Major element concentrations weremeasured by conventional X-ray fluorescence (XRF) spectrometry offusion disks on a Phillips X'Unique II PW1480 spectrometer and traceelement concentrations were determined by ICPMS using a PerkinElmer/Sciex Elan 6000 mass spectrometer at the Department ofGeological Sciences, University of Cape Town. For the latter technique,50mg of sample powder was first cleaned in ultrapurewater and thendissolved in HF and HNO3 and eventually analysed against five-pointcalibration curves. For the analytical details of these techniques seeFrimmel et al. (2001b). Typical lower limits of detection for the traceelement concentrations reported here are b0.01 ppm. Contaminationby oxides, sulphides or silicates can be a potential problem in theinterpretation of the results. This problem was minimised by carefulsample selection and handpicking of chips. It should be noted,however, that Nothdurft et al. (2004) tested dissolution of carbonaterocks in acids of variable strength (1 N acetic acid and 15 N HNO3) andthey could not observe any significant differences in the REE recoveryand patterns. The procedure followed in this study, i.e. rockdissolution in HF, invariably liberated more silicate-bound traceelements than dissolution in HNO3 only or in acetic acid. Dissolutionin HF was preferred because it keeps Th and Zr in solution, whereas

HNO3 may release REE from clastic components but not Zr. As will bediscussed later, Zr serves as important monitor of contamination andobtaining a correct concentration is thus crucial for the subsequentinterpretation.

The Rosh Pinah Formation carbonates are all dolomitised micro-bialaminite. Based on O and C isotopic evidence, the dolomitisationhas been related to hydrothermal activity along a growth fault at themargin of a volcanic rift graben, which, in places, led to economicstratiform sulphide mineralisation (Frimmel and Lane, 2005). Theanalysed samples come from the farm Spitzkop 111, some 10 to 18 kmnorth–northwest of Rosh Pinah. There five dolomite beds, each a fewdecimetres in thickness, occur at the top of upward-fining cycles andone such bed at the bottom of an upward-coarsening cycle (Frimmeland Lane, 2005). In each cycle, the dolomite is associated with arkose,calcareous sandstone, siltstone and mudstone. Ripple marks arecommon in the siliciclastic parts of the cycles as well as in tuffite. Closeto a hydrothermal breccia in felsic volcanic rocks, representing ahydrothermal vent site, the dolomite is enriched in Fe and Mn(Samples KL01-36 and KL01-38 as opposed to KL01-39 to KL01-41).Magnetite is the principal Fe-phase in all samples except for one(KL01-30), which contains abundant pyrite and is also enriched in anumber of minor and trace elements, such as Ti, K, Rb, Cu, Y, Zr, Nb, Cs,and Th. The pyrite-rich sample is explained as reflecting stronghydrothermal contamination by nearby felsic volcanic degassing(Frimmel and Lane, 2005).

The nine analysed carbonate samples from the Dabie RiverFormation come from the Helskloof Pass in the Richtersveld NationalPark in westernmost South Africa (16.979°E, 28.3124°S). Theycomprise dolomitic microbialaminite (PF2, 10), syn-sedimentarylimestone breccia (PF17, 18), syn-sedimentary dolomitic breccia(PF8), dolomitic stromatolite (PF4, 31, 34), and limestone from thelower part of the formation (PF121). The stromatolite sample PF34

Fig. 7. Distribution of the main stratigraphic units of the central West Congo Belt and SW–NE profile (from Frimmel et al., 2006).

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comes from the immediate contact with the overlying NumeesFormation diamictite.

In order to test the facies-dependence of the REE+Y distribution incarbonates within a given stratigraphic unit, two different profilesthrough the Bloeddrif Member were considered for this study. Oneprofile (total of 17 samples, PF216–232) represents a proximal sectionsampled in the southern Richtersveld region at 17.07°E, 28.63°S. Thesection is characterised by large-scale early diagenetic dolomitisationand few clastic (quartz arenite) intercalations as well as a positive δ13Canomaly (Frimmel and Fölling, 2004). The second profile (total of 17samples, PF25–30, PF89–102) represents a distal section just south ofthe Orange River at 16.8550°E, 28.3931°S. It comprises only cleanlimestone that is devoid of continental detrital input and displays aδ13C trend from negative ratios to values around 0‰ (Fölling andFrimmel, 2002).

The carbonate samples from the Kombuis Member, Cango CavesGroup, in the Saldania Belt look identical to those of the distal sectionthrough the Bloeddrif Member. They are all limestone devoid ofsignificant detrital input and show neither petrographic nor geo-chemical evidence of significant post-depositional alteration (Föllingand Frimmel, 2002). The sampleswere collected near the Cango Caves,21 km north of Outshoorn (22.2157°E, 33.3938°S). For comparison,three limestone samples from the Nooitgedagt Member were alsotaken from road cuts at Schoemanspoort southeast of the Cango Caves(22.236°E, 33.414°S).

The carbonate samples from the Otavi Group were selected on thebasis of petrographic and cathodoluminescence studies that revealeddifferent generationsof carbonate, especially in the vicinity of epigenetichydrothermal base metal sulphide mineralisation and later supergenealteration in the course of karstification (Frimmel et al., 1996a; Verran,1996; Chetty and Frimmel, 2000). They include the following: onedolomicrite sample, representative of the Berg Aukas Formation, fromthe abandoned Berg Aukas mine at the eastern end of the OtaviMountainland (18.250°E, 19.516°S); five micritic limestone samples(DV23, 32–35) from the Keilberg Member of the lower MaiebergFormation, a dolomicrite from the middle Maieberg, and a micriticlimestone (DV9) as well as a dolomicrite (KH91) from the upperMaieberg Formation, all taken from the Khusib Springs mine (18.022°E,19.426°S); two dolomicrite samples from the upper ElandshoekFormation (2965-300, DCT2) near the abandoned Tsumeb mine(17.715°E, 19.240°S); and dolomicrite from the Hüttenberg Formation,one sample from the Tsumeb mine and another from the Kombat mine(17.702°E, 19.710°S).

Carbonates from both the Haut Shiloango and the Schisto-CalcaireSubgroups of the West Congolian Group were selected for this study.All samples come from the central portion of the foreland to the east ofthe West Congo Belt in the Bas Congo province of the DemocraticRepublic of Congo (Frimmel et al., 2006). Eleven samples from theHautShiloango Subgroup comprise largely limestone. In addition, threecalcpelite samples were included to test the influence of shalecontamination. The different lithofacies of the Schisto-Calcaire Sub-group are represented by three dolomitic limestone samples from theC1, eight limestone samples from the C3, 21 variably dolomitisedlimestone samples from the C4, and 18 limestone, dolomitic limestoneand calcpelite samples from the C5 lithofacies.

4. Results

Trace element concentrations are reported in Tables A1–A6. REE+Yconcentrationswere normalised to the Post-Archaean Australian Shale(PAAS) composite (Taylor and McLennan, 1985). Overall, the total REEcontents in clean carbonate samples are expectedly very low. ElevatedREE contents could be the result of contamination with oxides,sulphides, phosphates or silicates, derived mainly either from hydro-thermal input and/or from the presence of terrestrial particulatematter (shale). Fortunately, these contaminants, though potentially

rich in REE, typically have distinctly different REE+Y patterns.Zirconium was used as monitor element to assess the extent of shalecontamination, because Zr is abundant in shale (210 ppm in PAAS) buteffectively absent in low-temperature waters. Similarly, Al concentra-tions were used as further proxy for small amounts of shalecontamination. This was possible because of the analytical proceduresemployed that ensured that all clays present in the samples were alsodissolved. Shale-normalised (SN) elemental anomalieswere calculatedon a linear scale, assuming that differences in concentration betweenneighbouring pairs are constant, as follows: La/La⁎=La/(3Pr−2Nd), Ce/Ce⁎=Ce/(2Pr−Nd), Eu/Eu⁎=Eu/(0.67Sm+0.33 Tb), and Gd/Gd⁎=Gd/(2Tb−Dy). As an alternative, following Lawrence et al. (2006), theanomalies were also calculated from a geometric average, assumingthat the ratio between near neighbour concentrations is constant, asfollows: La⁎=Pr⁎ (Pr/Nd)2, Ce⁎=Pr2⁎Nd, Eu⁎=(Sm2⁎Tb)0.33, Gd⁎=(Tb2⁎Sm)0.33, and Lu⁎=Yb2/Tm. In most cases, the differences in theresults obtained by these twomethods areminor (b5%) or negligible. Ifnot indicated otherwise, the elemental anomalies are reported asobtained using the linear method.

4.1. Carbonates from the Gariep and Saldania Belts

The dolomitic rocks of the Rosh Pinah Formation are overallrelatively rich in REEwith two samples even approaching PAAS values.The REE+Y distribution deviates, however, from that of PAAS bydisplaying a marked relative enrichment in the middle REE (NdSN/DySN=0.56±0.19; Fig. 8). Two samples that were taken from theimmediate contact with a Fe- and Mn-oxide-rich hydrothermalbreccia in underlying felsic volcanic rocks (KL01-36 and KL01-38)show higher total REE contents and a distinct negative Eu anomaly((Eu/Eu⁎)SN=0.51−0.72) compared to those samples taken fromgreater distance (Fig. 8a). Some of the latter have even a positive Euanomaly with the highest (Eu/Eu⁎)SN ratios (1.7) achieved in pyrite-

Fig. 8. Shale-normalised REE+Y patterns of Rosh Pinah Formation carbonates. (a) Profileaway from a hydrothermal vent; (b) variably contaminated and almost uncontaminateddolomitised microbial limestone.

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bearing carbonates (e.g. KL01-23). Although this could be an analyticalartefact of overlap with the peak for BaO, the noted positive Euanomalies are regarded as real signals because of a lack of a systematicpositive correlation between measured Ba and Eu signals (Table A1).Except for two samples, no Yanomaly can be observed in the analysedcarbonates. All other carbonates from the Port Nolloth and the CangoCaves Groups have REE+Y concentrations that are one to two orders ofmagnitude lower than those of PAAS.

The carbonate samples of the Dabie River Formation displayrelatively uniform shale-normalised REE+Y patterns (Fig. 9a) with aslight light REE depletion (mean (Nd/Yb)SN=0.79). The limestonesample (PF121) is distinguished by the highest REE+Y concentrations.Otherwise there are no systematic differences in the trace elementdistribution between the various carbonate types analysed (lime-stone, dolomitic microbialaminite, stromatolite and syn-sedimentarybreccia). Most samples of each lithotype display a slight positive Yanomaly ((Y/Ho)SN≤1.38), whereas others have Y/Ho ratios that areclose to the PAAS value of 27.3. One stromatolite and one dolomiticbreccia sample show a slight positive Eu anomaly ((Eu/Eu⁎)SN=1.5 and1.7, respectively), andmost samples also feature a positive Gd anomaly(mean (Gd/Gd⁎)SN=1.14).

The two profiles through the Bloeddrif Member carbonates yieldedvery consistent results, but significant differences can be noted betweentheproximal section (Fig. 9b) and thedistal section (Fig. 9c), i.e., betweendolomitic and limestone samples. The only limestone sample in theproximal section (PF229) has the lowest REE+Y concentrations andcompares well with the limestone samples from the distal section. Thedolomitic samples from the proximal section display only very mildlight REE depletion and weak positive Gd and Y anomalies in mostsamples (mean (Gd/Gd⁎)SN=1.1, mean (Y/Ho)SN=1.1). In contrast, thelimestone samples from the distal section show a distinct positive Yanomaly with (Y/Ho)SN consistently around 1.7. They are furthercharacterised by a significant light REE depletion ((Nd/Yb)SN=0.41±0.09). The proximal carbonates lack a La anomaly but the distalcarbonates show a positive La anomaly with a mean (La/La⁎)SN of 1.80(1.47whenusing the geometric, log-linearmethod). Aweak positive Ceanomaly is also common to the limestone samples ((Ce/Ce⁎)SN=1.11±0.10). Very low REE concentrations in some samples prevented acomplete set of analyses to be obtained for the middle and heavy REEand thus the existence of a Gd anomaly could not be verified for allsamples. Most samples, however, display a weak positive Gd anomaly((Gd/Gd⁎)SN=1.14±0.05).

The REE+Y patterns obtained for the Kombuis Member limestonesamples show all the same characteristics as those for the distalBloeddrifMember limestone samples (Fig.10a). The REE concentrationsare so low that they are, for some samples, below the lower limit ofdetection for somemiddle andheavyREE. Overall, a clear trend towardslight REE depletion is evident ((Nd/Yb)SN=0.46±0.20), as well as adistinct positive Y anomaly ((Y/Ho)SN=1.65±0.11). Similarly, a signifi-cant positive La anomaly ((La/La⁎)SN=1.57 or 1.36, depending on thecalculationmethod) and positive Gd anomaly ((Gd/Gd⁎)SN=1.22±0.06)are observed, whereas no significant Ce anomaly is noted ((Ce/Ce⁎)SN=1.07±0.03). The three samples from the Nooitgedagt Member(PF200, 201, 203) show very similar patterns but a smaller La anomaly(Fig. 10a). A distinct positive Ce anomaly ((Ce/Ce⁎)SN=1.32±0.14)distinguishes them, however, from the Kombuis Member limestones.

Fig. 9. Shale-normalised REE+Y patterns of (a) Dabie River Formation carbonates,(b) carbonates of the Bloeddrif Member, lower Holgat Formation, in a proximal sectionsouth of the Kuboos Pluton, and (c) Bloeddrif Member carbonates from a distal sectionnorth of the Kuboos Pluton.

Fig. 10. Shale-normalised REE+Y patterns of carbonates from (a) the Kombuis andNooitgedagtMembers (Matjies River Formation, CangoCavesGroup) and (b) theOtavi Group.

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4.2. Otavi Group carbonates

All analysed carbonate samples fromtheOtaviGrouphaveone to twoorders of magnitude less REE+Y compared to PAAS, with the limestonesamples generally yielding higher concentrations than the dolomiticsamples (Fig. 10b). The Berg Aukas Formation dolomicrite shows lightREE depletion ((Nd/Yb)SN=0.37) and amarked positive Eu anomaly ((Eu/Eu⁎)SN=1.84). All the Maieberg Formation carbonates, both limestoneand dolostone, independent of stratigraphic position within theformation, display remarkably uniform REE+Y patterns without anysignificant element anomalies. An exception is one dolomicrite thatshows a slight positive Y anomaly. The light REE depletion is lesspronounced than in the other carbonates ((Nd/Yb)SN=0.63−1.09).Similarly, the Elandshoek Formation dolomicrite yielded relativelyuniform REE+Y patterns but with a stronger light REE depletion ((Nd/Yb)SN=0.24−0.74). In contrast, the Hüttenberg Formation dolomicriteshows a distinct positive Yanomaly ((Y/Ho)SN=1.40−1.47) butotherwisea pattern similar to the other samples. Noteworthy is a general absenceof a Ce anomaly in any of the Otavi Group carbonates.

4.3. West Congolian Group carbonates

The limestone samples of the Haut Shiloango Subgroup show veryconsistent, uniformREE+Ypatterns (Fig.11a), independentof total REE+Ycontent. No distinct element anomalies are present. Light REE depletion,omnipresent in almost all other carbonates analysed in this study, is onlyvery weakly developed ((Nd/Yb)SN=0.82±0.07). All but two samples lacka positive Y anomaly. To assess the possible influence of shale con-tamination on the overall REE+Y patterns, three calcpelite samples wereinvestigated as well and they yielded indistinguishable patterns butoverall higher REE+Y concentrations that are close to PAAS values.

Similarly, theREE+Ypatternsobtained for thedolomitic limestoneoftheC1and the limestoneof theC3 lithofacieswithin the Schisto-CalcaireSubgroup are very uniform and lack significant element anomalies,except for one limestone sample that displays aweak positive Yanomaly(Fig. 11b). The C1 dolomitic limestones have overall higher (but still anorder of magnitude less than PAAS) REE+Y contents but show nosignificant difference in the distribution of these trace elements.

Amongst the samples from lithofacies C4, dolomitic limestone hasgenerally lower total REE+Y concentrations than the pure limestone; insome cases reaching the lower limit of detection for several elements.Most of the limestone samples, which still have one to two orders ofmagnitude less total REEcontents thanPAAS, the shale-normalisedREE+Ypatterns are flat and uniform (Fig. 11c). In some samples, especially thepartlydolomitisedones, anenrichment in lightREE (particularly La) canbenoted. Two samples yielded a positive Y anomaly ((Y/Ho)SN=2.81),whereas the remainder lacks such an anomaly.

The shale-normalised REE+Y patterns obtained for most of the C5carbonates are again flat and uniform (Fig. 11d) with no elementanomalies. Overall, the dolomite-bearing carbonates have higher totalREE+Y concentrations than pure limestone but display similarpatterns. Three calcpelite samples, for comparison, expectedly containone to two orders of magnitudemore REE+Y (corresponding to PAAS).They show a trace element distribution that is analogous to that of thelimestone and variably dolomitised carbonate rocks. In few carbonatesamples, a slight positive Eu anomaly is observed and those limestonesamples with the lowest total REE content seemingly have a negativeYb anomaly. As the concentrations of the latter are close to theanalytical capabilities, this Y anomaly might not be real.

5. Interpretation and discussion

5.1. Contamination and alteration

Many of the REE+Y patterns obtained in this study do not conformto typical seawater patterns. Therefore the possibility of contamination

needs to be assessed carefully before deducing chemical peculiaritiesof the precipitating waters and thus depositional environment. Themost critical sources of contamination are continent-derived detritalmaterial, notably clay minerals, Fe–Mn-oxides and sulphides. Noth-durft et al. (2004) showed that as little as 2% shale contamination canalter the REE+Y pattern of marine carbonates to such an extent thatelemental anomalies become effectively eradicated, resulting in flat,uniform shale-normalised patterns. Useful monitors for the extent ofterrestrial particulate matter are the concentrations of elements, suchas Zr, Th and Al. They are concentrated in different detrital minerals,

Fig. 11. Shale-normalised REE+Y patterns of carbonates from the West CongolianGroup: (a) Haut Shiloango Subgroup, (b) lithofacies C1, (c) lithofacies C4, and (d)lithofacies C5 of the Schisto-Calcaire Subgroup.

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such as zircon and clayminerals. Consequently, a positive correlation isto be expected between REE and Zr, Th as well as Al concentrations.Such a relationship is exemplified by the samples from the WestCongolian Belt, which include calcpelites, as depicted in Figs.12a and b.A very good positive correlation exists also between Al and Th(R2=0.99; not shown). Those samples with high REE concentrationsare clearly contaminated by a considerable shale component and arethus not used any further for palaeoenvironmental reconstructions. Itshould be noted, however, that even those carbonate samples thathave very low REE contents and less than 2% of the Zr concentration ofPAAS (e.g. sample HFWC67) yielded the same overall REE+Ypattern asthe others (see Fig. 11).

Terrestrial detrital contamination should follow the trends shown inFigs.12a and b. Plotting all data into a total REE versus Zr space (Fig.12c)reveals, however, two distinctly different trends. The steep trend on thisdiagram is defined essentially by the West Congolian samples andreflects contamination by terrestrialmaterial. The other trend, followingonly a marginal increase in Zr with increasing REE concentration, isdefinedby theRoshPinahFormationcarbonates. For the latter, bothfieldrelationships and geochemical evidence for a strong hydrothermalinfluence have beenpresentedpreviously (Frimmel and Lane, 2005) andthe elevated REE contents in some samples can be explained byhydrothermal alteration. This is supported by REE+Y patterns thatdeviate markedly from PAAS (Fig. 8) and by the fact that those samplesthat were most intensely affected by hydrothermal fluid–rock interac-tion display distinct Eu anomalies. They can be either positive ornegative, depending on the redox state of the hydrothermal fluid and itstotal sulphur activity, with positive Eu anomaly observed in pyrite-bearing carbonate and negative Eu anomaly in Fe-oxide rich domains.These examples illustrate that hydrothermal contamination of carbo-nates can be readily distinguished from terrestrial-detritalcontamination.

In the following interpretation of the REE+Y patterns obtained forthe various Neoproterozoic carbonates, an upper threshold value of

4 ppm Zr will be applied, which corresponds to about 2% shalecontamination. This restriction should ensure that the observed REE+Ytrends can be considered to approximate those of the contemporaneouscarbonate-forming water unless the original REE concentration in thecarbonate was very low — an effect that may be seen, for example, insome of the West Congolian carbonates (Fig. 11). The majority of thestudied samples conform to the above requirement, but for somestratigraphic units, such as the Haut Shiloango Subgroup, only very fewsamples meet this criterion. All shale-normalised diagrams and valuespresented in this study are based on the PAAS values. If other shalecomposites are used as reference, such as the North American ShaleComposite (NASC), this would not change the overall trends and criticalelement anomalies (Alibo and Nozaki, 1999). The latter authors showedthat seawaterhas a positiveHoanomalywhennormalised againstNASC,but this anomaly is not seen when normalisation is carried out againstother reference values, including PAAS. The same was noted for Lu/Lu⁎.Strong positive La, negative Ce and even the weak positive Gd anomalyin seawater seem to be, however, independent of the normalisation (DeBaar et al.,1985; Alibo and Nozaki,1999). For a more detailed discussionof normalisation artefacts in REE+Y patterns see Kamber et al. (2005).

Comparison of the various shale-normalised REE+Y patternsobtained in this study reveals that Eu anomalies are independent ofthe total REE content and thus unlikely to be related to shalecontamination. The Eu3+/Eu2+ redox potential in aqueous solutionsdepends mainly on temperature and to a lesser extent on pressure, pHand REE speciation (e.g., Bau, 1991), which explains the positive Euanomalies typically found in acidic, reducing hydrothermal fluids.Thus the positive Eu anomalies observed in some samples may be theproduct of either admixture of hydrothermal fluids or co-precipitationof hydrothermal Fe-sulphide. The carbonates from the Rosh Pinah andBerg Aukas Formations provide good examples of precipitation fromhydrothermally influenced waters. Both formations are characterisedby syn-sedimentary exhalative base metal sulphide mineralisationand both contain carbonates with distinct positive Eu anomalies.

In addition, selective Eu mobilisation by diagenetic fluids maycause Eu anomalies. It is also possible that Eu is preferentially releasedduring weathering if present as Eu2+ in the weathered source (Nozakiet al., 2000), especially if it is rich in feldspar (Kamber et al., 2005). Thismay be the case in some samples from the Schisto-Calcaire Subgroup,notably the C5 lithofacies, which display positive Eu anomalies. The Euanomalies do not seem to be an artefact of insufficient correction forBaO+ because no significant correlation exists between Ba and Euconcentrations (Table A6). Consequently, Eu seems to be of little usefor assessing the original depositional setting.

Many of the analysed carbonates are dolomitic and thus thequestion arises as to the effect of dolomitisation on the REE+Ypatterns. Comparison of dolostone and limestone data from thevarious units, such as the Dabie River Formation (Fig. 9a), the proximalsection through the Bloeddrif Member (Fig. 9b), the MaiebergFormation (Fig. 10b), and the C4 and C5 lithofacies of the Schisto-Calcaire Subgroup (Fig. 11c,d), reveals that there are no systematicdifferences in the REE+Y patterns between dolomitised and non-dolomitised samples. No relation can be observed between the degreeof dolomitisation and REE abundance, which confirms the observationby Banner et al. (1988), who found that dolomitisation of Mississip-pian limestones did not significantly affect their REE signatures.

5.2. Marine versus non-marine origin

A conspicuous feature common to the majority of the analysedcarbonate rocks is the deviation of their REE+Y distribution from anormal seawater composition. Only the carbonates from the distalsection through the Bloeddrif Member, the Kombuis Member, andfrom the Hüttenberg Formation have positive Y anomalies with Y/Horatios (47±3, 45±3, and 39±1, respectively) that are markedly higherthan that of the upper continental crust. Estimates of the latter are 27.5

Fig. 12. Positive correlation between the total REE concentration and respectively Al2O3

(a) and Zr (b) in the carbonate rocks and calcpelites of theWest Congolian Group reflectsvariable shale contamination. (c) Total REE versus Zr concentration plot for all analysedsamples shows two trends, one depicting shale contamination (see Fig. 11b), the otherhydrothermal contamination (dominated by Rosh Pinah Formation samples).

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(Taylor and McLennan, 1985) and 26.24 (based on more recent workon sediments from Queensland by Kamber et al., 2005). Forcomparison, the Y/Ho ratio in open seawater is typically between 60and 90 but is strongly dependent on salinity (Lawrence et al., 2006).

The relatively unfractionated shale-normalised REE+Y patternsdisplayed by the majority of Neoproterozoic carbonates could beexplained by formation in a meteoric environment. A lacustrinesetting for most of the analysed carbonates is, however, very unlikelyon the basis of field relationships and geochemical/isotopic data. Onlyin the case of the Rosh Pinah Formation, a lacustrine setting cannot beexcluded and would be in agreement with the field evidence of thincarbonate beds in cycles that are dominated by siliciclastic depositswith graded bedding (Alchin et al., 2005). Except for the lack of aredox-controlled Ce anomaly, the Rosh Pinah patterns resemble thoseobtained for Lake Tanganyika, which is not only one of the largest anddeepest lakes in the world but is also affected by sub-lacustrinehydrothermal activity (Barrat et al., 2000). A similar hydrothermalinfluence into a restricted basin (near-shore) is also suggested for theBerg Aukas Formation, which hosts syn-sedimentary exhalativemineralisation.

Although Y and Ho have similar ionic radii, identical charge, andthus similar geochemical behaviour, Ho is removed from seawatertwice as fast as Y because of differences in the surface complexationbehaviour (Nozaki et al., 1997). This makes the Y/Ho ratio aparticularly useful monitor for the differentiation between marineand non-marine deposits (Bau,1996; Nothdurft et al., 2004). Amongstall those studied stratigraphic units that have Y/Ho close to the uppercontinental crust value, samples exist that lack a shale-normalised Yanomaly in spite of no shale contamination being recognisable. Thislack of a significant Y anomaly is therefore likely to be a primaryfeature of the precipitating water. Yttrium/Ho ratios in riverwaters are equal to, or slightly higher than, the PAAS value butwell below seawater values, i.e., below 60 (e.g., Lawrence et al., 2006).The slight variability indicates limited fractionation of Y/Ho in thefluvial environment, depending on weathering reactions of indivi-dual REE+Y-bearing minerals and the pH of the water (Elderfieldet al., 1990; Lawrence et al., 2006).

The Y/Ho ratios for almost all analysed carbonate samples, exceptfor the distal Bloeddrif Member, the Kombuis Member, and theHüttenberg Formation, are within the range given for river waters. Aconsiderable freshwater influence may be, therefore, inferred for thedepositional environments of these carbonates.

General light REE depletion, marked positive La and negative Ceanomalies as well as a weak positive Gd anomaly are further featurestypically ascribed to a seawater origin. The Gd anomaly, however, doesnot seem to be unique to seawater. It has been described also frommodern river waters in Australia (Lawrence et al., 2006). Conse-quently, the weak positive Gd anomalies that are present in almost allof the analysed samples are not considered diagnostic of a specificdepositional environment.

Those samples with little light REE depletion or even enrichmentare characterised by Y/Ho ratios close to PAAS (Fig. 13a). The latterrelationship is in agreement with a strong freshwater componentbecause dissolved river load typically displays uniform, largelyunfractionated REE+Y patterns with only mild light and heavy REEdepletions (Goldstein and Jacobsen, 1988; Lawrence et al., 2006).

A positive La anomaly is present in almost all samples, but it isgenerally much smaller than expected for a normal marine carbonate.Expectedly, those sampleswith a positive La anomaly also tend to displaya stronger Y anomaly, in agreement with a marine origin (Fig. 13b).Contamination, be it hydrothermal or detrital (shale), results in thedisappearance of the La anomaly and this is clearly seen in the (La/La⁎)SNversus Zr diagram (Fig. 13c). The same diagram also illustrates, however,that not all of those samples that show very little or no signs ofcontamination (Zr b1 ppm) display a distinct positive La anomaly.Consequently, the lackof a La anomaly in these samples is taken as anear-

primary signal of the depositional environment, which is unlikely to beopen marine but may be strongly influenced by freshwater.

As oxidised Ce4+ is less soluble and more readily adsorbed ontoparticles than Ce3+, the extent of Ce depletion reflects the oxygenationstate of the water. Not surprisingly, Ce oxidation takes placepreferentially at shallow water depths (e.g., Alibo and Nozaki, 1999),and evidence of Ce oxidation in seawater exists for the time since thePalaeoproterozoic, but not from the Archaean (Kamber and Webb,2001; Kamber et al., 2004; Bolhar and Van Kranendonk, 2007). Ceriumoxidation is more sensitive to pH than to oxygen fugacity and favouredby alkaline solutions (Elderfield and Sholkovitz, 1987). For this reason,and for a shift of the Ce4+/Ce3+ redox equilibrium towards higheroxygen fugacity with increasing temperature, negative Ce anomalies

Fig. 13. (a) Relationship between positive Y anomaly and light REE depletion, expressedas (Nd/Yb)SN; (b) positive correlation between Y and La anomalies for all samplesanalysed; (c) Relationship between La anomaly and contamination, expressed in termsof Zr concentration.

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are typically absent in low-pH and/or hydrothermal precipitates.Effectively all of the analysed carbonate samples lack a negative Ceanomaly. This phenomenon is independent of total REE concentrationand Zr or Al content and thus seems to be a primary feature that couldimply only poorly oxygenated and/or relatively acidic waters. Themost extreme samples, a dolostone from the Dabie River Formation(PF10) and a limestone sample from the Nooitgedagt Member (PF201)both of which are characterised by very low total REE contents, lightREE depletion and a positive Y anomaly, are distinguished by even amarked positive Ce anomaly ((Ce/Ce⁎)SN=1.42 and 1.48, respectively).An explanation for this behaviour of Ce may be obtained from thecomparison of data from proximal and distal sections through thesame stratigraphic unit.

5.3. Near-shore versus open marine environments

One of the main findings of this study is the observation that, ingeneral, almost all of the analysed carbonates are richer in traceelements (including REE) than Archaean or Proterozoic carbonates. Tofind an explanation for this difference, it is advantageous to assess thesignificance of depositional environment, i.e., the proximity to palaeo-shore lines, on the carbonate geochemistry. The data obtained on thetwo sections through the Bloeddrif Member at the bottom of theHolgat Formation neatly illustrate the differences in trace elementbehaviour between near-shore and distal environments. At the sametime they might explain the non-marine signature observed in themajority of the analysed carbonates. The two studied sections throughthe same stratigraphic unit, a cap carbonate succession above theglaciogenic Numees Formation diamictite, differ in terms of lithology,chemistry and isotopic composition (Frimmel and Fölling, 2004). Onesection consists entirely of limestone with high Sr concentrations(1078–2483 ppm) and low Mn/Sr, Fe/Sr and Ca/Sr ratios indicative ofonly very limited post-depositional alteration. Low SiO2 (0.4–3.1 wt.%),Y (≤2 ppm), Zr (≤16 ppm) and low Rb (≤5 ppm) contents reflect onlyminimal continental detrital input, which is in agreement with a distaldepositional environment. In contrast, the other section is largelydolomitic, has thin sandstone beds intercalated, and the carbonatestherein contain an order of magnitude more Y, Zr, and Rb, reflecting alarger continental detrital component in a proximal, near-shoreenvironment.

Of particular importance is the observation that the C isotopiccomposition of the least altered carbonates follows different trends inthe two sections (Frimmel and Fölling, 2004; Fig. 4). The bottom ofboth sections contains a negative δ13C excursion (b−3‰) but furtherupwards, recovery of the δ13C ratios to around 0‰ is noted in thedistal section, whereas a marked positive δ13C excursion (+5‰) iscontained in the proximal section. This enrichment in 13C within theproximal section has been interpreted as evidence of a strongersalinity stratification in a poorly circulated sea (Frimmel and Fölling,2004). If such a suggested strong palaeo-environmental control onδ13C exists, it would cast doubt on a series of chemostratigraphiccorrelations of Neoproterozoic strata many of which are typicallybased on C isotopic anomalies. Since it has been shown that salinityexerts a strong control on the distribution of REE+Y (Lawrence et al.,2006), the data obtained in this study might help in betterinterpreting the noted difference in δ13C in distal and proximalprofiles.

The REE+Y patterns obtained for the various limestone samplesfrom the distal section are all essentially identical, irrespective of totalREE concentration. Contamination is negligible. The patterns bearfeatures typical of seawater, such as a positive Y anomaly but the Laand Ce anomalies are not as pronounced as expected for open marinechemical sediments. The continental detrital influx in the proximalsection is also reflected by overall higher REE, Zr, Th, and Al contents aswell as a tendency towards uniform, flat shale-normalised REE+Ypatterns. Most representative of the original water composition are

those samples with the lowest total REE (about 1% of PAAS values) andthese differ from the distal limestone samples by having a lesspronounced Yanomaly, lack of a La anomaly, and overall less light REEdepletion. This, together with the observed lack of a negative Ceanomaly in all of the proximal carbonates, points to a considerableriverine influx in the near-shore depositional section of the BloeddrifMember.

As the mean oceanic residence time of C is even shorter than thatof most REE (exceptions are Ce and possibly Pr and Eu; Alibo andNozaki, 1999), it is doubtful whether near-coastal carbonate rocks areuseful proxies for the C isotopic composition of ancient ocean waters.Similarly as the REE+Y patterns in these rocks do not reflect openmarine conditions, their C isotopic composition is unlikely to berepresentative when and where local changes in the ocean waterchemistry occurred, be it due to enhanced evaporation or dilution byriver water. This explains the discrepancy in the C isotope profilesthrough the distal and proximal sections of the Bloeddrif Member. Thedifference in δ13C, with a positive excursion in the proximal sectionand the lack thereof in the distal section, is not necessarily just afunction of diachronous sediment deposition from the slope up to theshelf, as suggested by Hoffman et al. (2007), but is likely to becontrolled by local palaeoenvironmental circumstances. Conse-quently, the C isotope ratios are not representative of ambientseawater composition and caution must be taken when using Cisotope profiles from proximal sections for chemostratigraphiccorrelation.

5.4. Freshwater versus near-shore colloids

Many of the relatively flat shale-normalised REE+Y patternscannot simply be explained as being contaminated and thus notrepresentative of ambient water chemistry. This is best seen from thewide range in La anomalies at low Zr concentrations (Fig. 13c). Basedon the REE+Y distribution alone, carbonate formation in eitherfreshwater or brackish water in an estuary would be possibleexplanations for the relatively flat, unfractionated REE+Y patterns.As discussed above, a lacustrine origin of the carbonates can beexcluded except for maybe the Rosh Pinah Formation carbonates.Strontium isotope ratios may help to distinguish between a coastalfringing and a lacustrine environment. In contrast to C and REE, whoseoceanic residence time is on the order of hundreds to a few thousandyears (Alibo and Nozaki, 1999), Sr has a very long oceanic residencetime of about 2.4myr (e.g. Jones and Jenkyns, 2001) and is, therefore, amore reliable proxy. Most of the analysed carbonates yielded leastaltered 87Sr/86Sr ratios that are in the range expected for contempora-neous Cryogenian or Ediacaran seawater (Fölling and Frimmel, 2002;Frimmel et al., 2006). Although this argument may be circumstantialbecause the chronostratigraphic position has been indirectly inferredin some cases from Sr isotope ratios, the least altered 87Sr/86Sr ratiosare in all cases well below values that are expected for river waters, i.e.0.711 to 0.712 (Palmer and Edmond,1989). A freshwater source for thecarbonates can thus be excluded.

As shown by Lawrence and Kamber (2006), marine REE+Y signalsform already at very low salinities when mixing river water withseawater. Moreover, mixing with freshwater should lead to a dilutionof several other elements whose concentrations are generally elevatedin the studied carbonates. Consequently, precipitation from estuarinewaters can be also excluded.

An alternative and preferred explanation is the incorporation ofnear-shore colloids, possibly related to Fe-oxihydroxide scavenging.This is maybe best illustrated by the observed lack of a negative Ceanomaly in all of the proximal Bloeddrif Member carbonates. Themarine origin of these samples would require the presence of such ananomaly. A highly acidic and reducing depositional environment isunlikely and thus the absence of that anomaly would imply theintroduction of some material with a positive shale-normalised Ce

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anomaly. Sholkovitz (1992) showed that Fe-colloids in river watermake an excellent candidate for such a material and such a derivationseems indicated by a positive correlation between Fe and total REEcontents as illustrated in Fig. 14. The analysed samples follow twotrends on that figure. One trend passes almost through the origin,which shows that essentially all REE in these carbonate rocks wereintroduced via Fe-bearing phases. The only limestone sample, whichrepresents the protolith prior to dolomitisation, contains no detect-able Fe and only 1.55 ppm total REE. A second trend intersects the x-axis at about 7 ppm REE, which indicates a small variation in Fe-independent total REE contents in the dolomitised limestone. Thegood correlation between Fe and total REE in these samples highlightsthat riverine Fe-colloids have a strong control over the REEdistribution in these carbonates.

The same kind of trace element distribution as found in theproximal Bloeddrif Member carbonates seems to be a common featurein most of the analysed carbonates, especially the various micro-bialaminites and stromatolites (Rosh Pinah Formation, Dabie RiverFormation, Maieberg Formation, most of the West Congoliancarbonates). Many of them represent former microbial reefs. Basedon the elevated incorporation of near-shore colloids as inferred fromthe REE+Y patterns obtained in this study, a coastal fringing reefenvironment is therefore suggested.

6. Conclusions

The majority of a variety of Cryogenian and Ediacaran marinecarbonate rocks from the Pan-African Saldania, Gariep, Damara andWest Congo Belts do not displaya normalmarineREE+Y signature. Theyshow relatively flat, unfractionated shale-normalised REE+Y patternswithout the positive La, negative Ce and positive Y anomalies that aretypical of marine precipitates. Comparison of samples with variableshale contamination revealed the same patterns also for effectivelyuncontaminated carbonates. Moreover, comparison of the same calcar-eous stratigraphic unit, an Ediacaran post-glacial cap carbonate(Bloeddrif Member), in an open marine, distal position and in a near-shore, proximal position illustrated a strong dependence of the REE+Ydistribution on riverine particle influx. This reflects the overall relativelyshort residence time of these elements in seawater. River-born Fe-colloids are recognised as the particles that have the strongest effect ontheREEdistribution in themixingzonebetweenpure seawaterand riverwater. This effect is seen most prominently in the proximal Bloeddrifsection in the external Gariep Belt, the Maieberg and ElandshoekFormations of the Otavi platform in the northern Damara Belt as well asin most of the West Congolian carbonates. A coastal fringing reefenvironment is envisaged for these largely microbialaminitic tostromatolitic carbonates. Early diagenetic dolomitisation, which ismore widespread in the proximal environments, seems to have littleinfluence on the overall REE+Y pattern.

Apart from the distal Bloeddrif Member carbonates, those of theKombuis Member and the Hüttenberg Formation yielded typicalmarine REE+Y signatures as well. The identical results obtained forthe limestone of both the Bloeddrif and the Kombuis Membersconfirm the previously suggested correlation of these two units(Fölling and Frimmel, 2002; Gaucher and Germs, 2006). A similarmarine signature is observed for the Nooitgedagt Member limestone,but the Y anomaly is less pronounced and Ce displays a positiveanomaly. The latter is interpreted, analogous to the above conclusions,as indicative of a near-coastal depositional environment. This findingsupports the contention that the Nooitgedagt and Kombuis Members,though traditionally grouped into the same formation, were notdeposited in the same environment (Frimmel et al., 2001a; Fölling andFrimmel, 2002).

Finally, those carbonates that were affected by syn-sedimentaryhydrothermal activity, evident from the presence of stratiform, syn-sedimentary Fe–Mn-oxide and/or base metal sulphide mineralisation,are distinguished in their REE+Y patterns, particularly by showing adistinct positive or negative Eu anomaly in the cases of hydrothermalsulphide or Fe-oxide precipitation, respectively. This is evidentespecially in the Rosh Pinah Formation of the external Gariep Beltand the Berg Aukas Formation in the lower Otavi Group in thenortheastern Damara Belt.

Overall, this study has confirmed that REE+Y distributions incarbonates can be useful monitors of the depositional palaeoenviron-ment. After using REE+Y to better interpret Archaean (Kamber andWebb, 2001; Bolhar et al., 2004) as well as Phanerozoic carbonates(e.g. Nothdurft et al., 2004), this study is a first attempt to obtain abetter understanding of Neoproterozoic carbonates from the distribu-tion of these trace elements. Although the Neoproterozoic seawatercomposition remains relatively poorly constrained, a distinctionseems to be possible with the aid of REE+Y between open marine(distal), near-coastal and possibly lacustrine deposits, as well ashydrothermally influenced deposits.

The results obtained emphasise the presence of concentrationgradients in REE+Y away from shore in agreement with previouslynoted concentration gradients in other trace elements. Shen and Boyle(1988), for example, described elevated Ba concentrations in near-shore corals compared to distal, oceanic corals — an enrichment thathas been explained by terrestrial run-off (Shen and Sanford, 1990).More recently, Sinclair (2005) pointed out, however, that a series ofenvironmental and/or biological factors can influence the Ba distribu-tion in modern corals. Similarly, it can be expected that futuregeochemical studies of ancient carbonates will reveal that a number ofhitherto poorly understood factors control the trace elementdistribution in these rocks.

Considering that even Ba, with a residence time of approximately10,000 years, follows a steep concentration gradient in near-shoreenvironments, an even stronger gradient is to be expected for traceelements, such as the REE and Y, whose residence time is, on average,an order of magnitude shorter. While these elements may serve asuseful monitors of the local depositional environmental, their use forregional or global chemostratigraphic correlation is very limited. Thislimitation is even more strongly applicable to C whose residence timeis even shorter than that of most REE. Overall the conclusion can bedrawn that many Neoproterozoic carbonate successions developed innear-shore environments in which C isotopes cannot be used asreliable chemostratigraphic indicator. Great caution should beapplied, therefore, when using δ13C ratios, especially positive excur-sions, from such succession for inter-basin correlation.

Acknowledgements

A. Späth and F. Rawoort are thanked for assistance with the ICPMSanalyses, D. Reid for XRF analyses. Thanks also go to L. Tack forproviding samples from the West Congo Belt. K.-P. Kelber helped with

Fig. 14. Total REE concentration versus Fe diagram for the proximal section through theBloeddrif Member carbonates (square — limestone, diamond — dolostone).

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the drafting of maps. Very constructive comments by an anonymousreviewer are much appreciated. Financial support from the SouthAfrican National Research Foundation (grant 2069090) is gratefullyacknowledged. This is a contribution to IGCP478 (“Neoproterozoic-Early Palaeozoic Events in southwestern Gondwana”) and IGCP512(“Neoproterozoic Ice Ages”).

Appendix A. Supplementary data

Supplementary data associatedwith this article can be found in theonline version, at doi:10.1016/j.chemgeo.2008.10.033.

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