The 2.7–2.0 Ga volcano-sedimentary record of Africa, India and Australia: evidence for global and...

Precambrian Research 97 (1999) 269–302www.elsevier.com/locate/precamres

The 2.7–2.0 Ga volcano-sedimentary record of Africa, Indiaand Australia: evidence for global and local changes

in sea level and continental freeboard

P.G. Eriksson a,*, R. Mazumder b, S. Sarkar b, P.K. Bose b, W. Altermann c,R. van der Merwe a

a Department of Geology, University of Pretoria, Pretoria 0002, South Africab Department of Geological Sciences, Jadavpur University, Calcutta 700-032, India

c Institut fur Allgemeine und Angewandte Geologie, Ludwig-Maximilians-Universitat, Luisenstrasse 37, 80333 Munchen, Germany

Accepted 21 April 1999

Abstract

The 2.7–2.0 Ga volcano-sedimentary records of the African, Indian and Australian cratons indicate two broadlydefined periods of extensive drowning of the emergent continental areas, concomitant with lowered freeboard.Carbonate-banded iron formation (BIF) platforms characterised the first such event, at ca 2.6–2.4 Ga (Africa andAustralia) to 2.7 Ga (India). These earlier globally enhanced sea levels are ascribed to increased mid-ocean ridgeactivity, possibly related to breakup of a postulated Late Archaean ‘southern’ supercontinent. Alternatively, atransition from global-scale catastrophic mantle overturn events to the onset of plate tectonics may have occurred inthe Late Archaean (Nelson, 1998. Earth Planet. Sci. Lett. 158, 109–119). Both explanations of increased mid-oceanridge activity are compatible with significant Early to Middle Archaean crustal growth (Armstrong, 1981. Phil. Trans.R Soc. London A 301, 443–472), with the emergent high freeboard cratons being subjected to aggressive weatheringand erosion. Enhanced continental crustal growth near the Archaean–Proterozoic boundary (McLennan and Taylor,1982. J. Geol. 90, 347–361), related to the development of significant island arc complexes, would have resulted incommon lowered freeboard–enhanced sea level conditions at the passive margins of the ‘southern’ cratons. Thediachronous nature of these earlier transgressions in the various cratons may reflect the effect of local tectonicmovements and/or the thermal state of the cratons. From ca 2.4–2.2 Ga, cratons that make up the present-daycontinents of India, Africa and Australia had relatively high continental freeboard and lowered sea levels. Glacigenicdeposits are preserved on the Kaapvaal (Africa), Singhbhum (India) and Pilbara (Australia) cratons. The secondbroadly defined drowning event, at ca <2.2 and >2.15 Ga, was probably due to post-glacial climatic amelioration.Freeboard was reduced by the combination of eustatic rise and the reestablishment of aggressive weathering aswarmer palaeoclimates returned. In India, carbonates were more prominent than the siliciclastic sediments (includingprominent black shales) seen in Africa and Australia. © 1999 Elsevier Science B.V. All rights reserved.

Keywords: Archaean; Continental freeboard; Proterozoic; Sea level; Volcano-sedimentary record

* Corresponding author. Tel.:+27-12-4202238; fax:+27-12-3625219.E-mail address: pat@scientia.up.ac.za (P.G. Eriksson)

270 P.G. Eriksson et al. / Precambrian Research 97 (1999) 269–302

1. Introduction the 2.7– 2.0 Ga volcano-sedimentary record ofAfrica, India and Australia, in order to try anddiscriminate different kinds of sea level change, asThe constant freeboard model of Wise (1972,

1974) is directly related to sea level changes con- well as variations in the continental freeboard. Inaddition, we will attempt to assess continentalcomitant with supercontinent amalgamation and

fragmentation [e.g. review by Eriksson (1999)]. crustal growth models, particularly whetherenhanced rates prevailed near the Archaean–Supercontinent attenuation and breakup generally

may coincide with enhanced mid-ocean ridge Proterozoic boundary, or Armstrong’s (1981)alternative suggestion of Early Archaean achieve-spreading activity, which leads to rise of sea level

and transgression of dispersing supercontinental ment of significant continental crustal volumes.fragments ( Windley, 1995). As an alternative idea,Nelson (1998) proposes that formation of conti- 1.1. Criteria used to identify and correlate

freeboard changes across cratonsnental crust in the Late Archaean reflects thesuperimposition of global catastrophic mantleoverturn and plate tectonic processes. He further Elevated freeboard of a craton is indicated by

evidence of large scale volcanism, particularlypostulates that global magmatic events were grad-ually replaced by plate tectonics as the dominant flood-basaltic events of inferred mantle plume

affinity. Equally important is the preservation ofcrust-forming process during a Late Archaean–Early Proterozoic transition period. If Nelson’s relatively widespread sedimentary successions of

continental origin, such as alluvial, aeolian and(1998) model is correct, then enhanced mid-oceanridge growth independent of supercontinent atten- lacustrine deposits. Evidence of high rates of ero-

sion, preferably combined with deep regoliths anduation and fragmentation could have been impor-tant during the approximate 2.7–2.0 Ga period. significant unconformities also support high free-

board conditions. While recognising that, ideally,A constant freeboard from ca 2.5 Ga, as pro-posed by Wise (1972, 1974) is compatible with these indicators should be developed on a craton-

wide scale for meaningful correlation of significantcrustal growth models proposing rapid attainmentof most of the continental crustal volume in the freeboard changes within and between cratons, the

preservation of the 2.7–2.0 Ga record makes thisearly Precambrian (e.g. Armstrong, 1981;McLennan and Taylor, 1982; Arndt, 1999; impractical. In addition, high freeboard-related

erosional and non-depositional episodes will gen-Eriksson, 1999). However, there is disagreementwhether enhanced growth rates occurred near the erally be much more difficult to recognise and

date, than widespread marine sediments of lowArchaean–Proterozoic boundary (e.g. Eriksson,1995), or whether much of the present volume of freeboard and high sea level affinity. Although

freeboard and eustatic changes are directly relatedcontinental crust formed prior to 3.8–3.6 Ga(Armstrong, 1981; Arndt, 1999). The geological (e.g. Eriksson, 1999), freeboard may also change

independently of sea level, particularly due torecord of the different Precambrian cratons sug-gests that this accelerated growth of continental thermal changes in a craton.

It is thus important to try and distinguish localcrust was diachronous, on a global scale (Eriksson,1995). As a result, continental freeboard condi- variations in both freeboard and relative sea level

from those of truly global character, and this papertions varied for these cratons, as also seen todayfrom the variable hypsometric curves of the reflects a first attempt, bearing in mind the con-

straints provided by limited accurate geochronol-different continents. Significant variation in free-board elevation (concomitant with variations of ogy, particularly pertinent for India and Africa.

However, the scale and duration of sea level10–40% in post-Archaean continental crustalgrowth rates) falls well within the bounds of several changes and freeboard variation during the chosen

2.7–2.0 Ga period was possibly greater than formechanisms causing sea level changes (Eriksson,1999). the Phanerozoic-modern era, due to significant

formation of continental crust (e.g. Eriksson, 1995;For these reasons, this paper will examine briefly

271P.G. Eriksson et al. / Precambrian Research 97 (1999) 269–302

Windley, 1995). The present investigation is thus exposure of part of a craton. Archaean rocks inthe East Saharan craton are restricted to thenecessarily limited to freeboard and sea level

changes of large geographic and temporal scales. Uweinat inlier, and the southeastern portion, bor-dering on the NE Congo craton (Fig. 1). Thelatter Archaean basement terranes extend to thenorth-northeast, to encompass also the Bomu and2. The 2.7–2.0 Ga volcano-sedimentary record of

Africa West Nile gneissic complexes of the East Saharancraton. Comparable metamorphic and charnocki-tisation events identified within both cratons indi-2.1. Introductioncate that the East Saharan may in fact be anorthward continuation of the Congo cratonThe crustal architecture of Precambrian Africa

comprises the Kalahari, Tanzania, Congo (also (Petters, 1991). Large scale intrusion of granitesand concomitant cratonic stabilisation in this lattercalled Zaire), West African and East Saharan

(Nile) cratons (Fig. 1). The Kalahari craton is region occurred at ca 2.47–2.41 Ga (Cahen et al.,1984).made up largely of the older (>2.7 Ga) Kaapvaal

and Zimbabwe cratons (Fig. 1), and of younger, Palaeomagnetic data (Onstott and Hargraves,1981) allow the possible reconstruction of theLate Archaean and Proterozoic mobile belts. These

two cratons exhibit the association of greenstones, relative positions of the African cratons for theEarly Proterozoic (e.g. Ledru et al., 1989). Thetrondhjemite–tonalite–granodiorite (TTG) and

granitic plutons, typical of all Archaean cratons relative positions of the Tanzania and Congo cra-tons were similar to today, but with the Tanzania(e.g. De Wit et al., 1992; Kramers, 1988). Outward

growth of the Kaapvaal craton (McCourt, 1995) long axis orientated SE–NW (Fig. 2). The smallcratonic and probably Archaean (Petters, 1991)from a nucleus terminated with the collision of the

exotic terrane of the Central Zone, Limpopo Bengweula Block was separated from the Tanzaniacraton by the Early Proterozoic Ubendian mobilemobile belt, at ca 2.68 Ga, and thereafter, at ca

2.58 Ga, the Zimbabwe craton collided with the belt. The Kalahari craton was situated approxi-mately south of the Bengweula and Tanzania cra-Kaapvaal–Central Zone plate (Treloar and

Blenkinsop, 1995). The latter suture zone was tons (Fig. 2). These relative positions indicated bythe palaeomagnetic data possibly resulted fromtranspressionally reactivated at ca 2.0 Ga, during

the widespread Eburnean orogeny of Africa early attenuation and breakup of the LateArchaean ‘southern’ supercontinent postulated by(Petters, 1991; Treloar and Blenkinsop, 1995). In

contrast to these views, Holzer et al. (1998) argue Aspler and Chiarenzelli (1998). A better con-strained African supercontinent had assembled byfor an ca 2.0 Ga age for the main Limpopo

orogeny. the end of the Eburnean orogeny at ca 2.0–1.85 Ga, encompassing the Kaapvaal, Bengweula,Stabilisation of the Tanzania cratonic basement

was achieved by ca 2.57 Ga, with widespread gran- Tanzania, Congo and East Saharan cratons ofAfrica.ite intrusions (Cahen et al., 1984); 2.7–2.0 Ga

volcano-sedimentary successions in this region are The Eburnean orogeny was thus a widespreadevent, affecting essentially the margins of most ofmostly highly altered. The Congo craton is gen-

erally poorly exposed; stabilisation was achieved the cratons of Africa (Fig. 3). In West Africa,reactivation of Archaean basement rocks (Guineaby large scale granitic intrusions at ca 2.7 Ga in

the NW (northern Gabon massif; term as used and Reguibat Rises; Fig. 1) was followed by theca 2.6–2.13 Ga Birimian Supergroup (Ledru et al.,here includes the de Chaillu massif in the south

thereof ), with Kasai and the Angola shield (Fig. 1) 1989; Milesi et al., 1989; 1991; Leube et al., 1990;Mortimer, 1992; Eisenlohr and Hirdes, 1992) andfollowing at ca 2560 Ma, and the northeastern

parts at ca 2430 Ma (Cahen et al., 1984; Petters, by subsequent large scale syntectonic intrusion ofgranites at 2102±1–2092±2 Ma (Hirdes and1991). The term shield is used here to denote a

cratonic nucleus or a geographically defined large Davis, 1998). The main Eburnean orogeny here

Fig. 1. Late Archaean and Early Proterozoic cratons, Eburnean (±2.0 Ga) mobile belts and 2.7–2.0 Ga volcano-sedimentary succes-sions of Africa. Note also the Late Archaean (?) Limpopo mobile belt between the Kaapvaal and Zimbabwe crustal blocks.

Fig. 2. Reconstruction of the relative positions of the East Saharan, Congo, Tanzania, Bengweula and Kalahari cratons in the EarlyProterozoic, based on palaeomagnetic data (Onstott and Hargraves, 1981; Ledru et al., 1989). Also shown are the Eburnean-agedmobile belts and the 2.7–2.0 Ga volcano-sedimentary successions. Note the present-day north direction as a frame of reference.

occurred between ca 2130 and 2050 Ma (Cahen mentary record of the West African craton (Fig. 1)is not one suitable for examining evidence for seaet al., 1984) (Fig. 3). The Birimian supracrustals

partially resemble Archaean greenstone belts level and continental freeboard changes. Similarly,the equivalent record in both the Tanzania and(Condie, 1989), thereby implying that cratonisa-

tion in West Africa was continuing during Birimian East Saharan cratons is unsuitable, largely due toalteration. Recourse must thus be made to thedeposition. Therefore, the 2.7–2.0 Ga volcano-sedi-

Fig. 3. Geohistory summary chart for the 2.7–2.0 Ga volcano-sedimentary successions of Africa. Note, at the base of each column,the approximate age of stabilisation of continental crustal basement (high freeboard).Two inferred transgression-epeiric marinerelated successions are shown: a lower, ca 2.65–2.43 Ga succession including BIF–carbonate rocks; an upper ca >2.15 Ga volcano-clastic sedimentary succession with common black shales. The onset of the Eburnean orogeny is shown for each column as well asmajor magmatic and tectonic events preceding this Africa-wide tectono-thermal event.

2.7–2.0 Ga volcano-sedimentary successions on and Walraven, 1993; Walraven et al., 1994), andthe 2714–2709 Ma (Armstrong et al., 1991)the Congo and Kalahari cratons, in order to try

and discriminate local from global sea level Ventersdorp Supergroup, of predominantly maficvolcanics and subordinate sediments, are an indica-changes, and in order to examine variation in

continental freeboard conditions. tion of extensional tectonics and the onset ofrifting (e.g. Van der Westhuizen et al., 1991). TheVentersdorp flood basaltic lavas were followed by2.2. Kaapvaal cratonthe ca >2659–2050 Ma (e.g. Eriksson et al., 1995)Transvaal Supergroup (Eriksson and Reczko,Although the Kaapvaal craton is thought to

have become fully cratonised by ca 2.7 Ga, depos- 1995; Catuneanu and Eriksson, 1999) (Table 1),preserved in three structural basins on theition of the ca 2.97–2.78 Ga (Robb and Meyer,

1995) Witwatersrand and coeval Pongola Kaapvaal craton: Transvaal; Griqualand West;and Kanye basins (Fig. 1).Supergroups already indicated a large measure of

crustal stability. The acidic rocks of the Gaborone Following the ca 2.68 Ga Limpopo mobile beltevent (Treloar and Blenkinsop, 1995), peneplana-Granite Suite and Kanye Volcanic Complex,

formed at ca 2.83–2.77 Ga (Sibiya, 1988; Grobler tion of the Ventersdorp rocks preceded the devel-

Table 1The Transvaal Supergroup, Kaapvaal craton: geological overview (data from: Button, 1973; Beukes, 1980, 1983, 1987; Eriksson,1988, 1997; Henry et al., 1990; Eriksson et al., 1993a, 1993b, 1994, 1995, 1996, 1998a; Altermann and Wotherspoon, 1995; Erikssonand Reczko, 1995, 1998; Reczko et al., 1995; Altermann and Siegfried, 1997; Altermann and Nelson, 1998; Eriksson andAltermann, 1998)

Supergroup Group Formation Lithology Age Sedimentary Sea level Freeboardhistory

Transvaal Pretoria [ca Post- Alternating Eastern and Probably Enhanced,<2350 Ma Magaliesberg mudrock and western remnant intermediate due to commonand formations arkosic to basins: arkosic>2054 Ma; quartzose continental detritus andEriksson and sandstone, sedimentation continentalReczko minor lava, and coastal sedimentation(1995); carbonate rocks wind-tidal flats systems along aArmstrong coastlineand McCourt(1997)]

Magaliesberg Mature Regressive Dropping Increasing duesandstones tidally due to regression

reworked braid- to generaldeltas regressive

settingSilverton Mudrocks, fine >2150 Ma Thick General Reduced due to

sandstones, (unpublished hemipelagic transgressive low sedimenteastern lens of Rb–Sr data, muds, silts, epeiric setting calibreandesitic lavas; first author) fine sands withtotal deposited below enhanced seathickness wave level>2000m base, and storm

and fair weatherdeposits abovewave base

Daspoort Mature Rift-related Rising due to Initially high andsandstones, alluvial fans, marine decreasing asminor mudrocks with marine reworking transgressionand reworking in east proceedsconglomerates of basin

Strubenkop Mudrocks, Lacustrine basin; Low Highminor lateral faciessandstones equivalent of

DwaalheuwelDwaalheuwel Immature Fluvial sheet Low High

sandstones, sandstonesminorconglomerates

Hekpoort Basaltic 2223±13 Ma Continental Low High (plume-(Ongeluk andesites (Cornell flood basalt related)Formation in et al., 1996) (subaerial )GriqualandWest)Boshoek Conglomerate, Alluvial-fluvial Low High

immature systemssandstone

Timeball Hill Up to 1400 m; Epeiric basin Locally high Enhanced due tolower and upper pelagic (epeiric sea) regionalmudrocks, sedimentation, glaciation inmedial contourites, Griqualand Westsandstones; distal low- basin and

Table 1 (continued )

Supergroup Group Formation Lithology Age Sedimentary Sea level Freeboardhistory

Transvaal Pretoria [ca minor periglacial density delta-fed periglacial<2350 Ma detritus; turbidites; for deposits locallyand Makganyene sandstones, tidal in Transvaal>2054 Ma; Diamictite flat reworking of basinEriksson and Formation in the regressiveReczko Griqualand West fluviodeltaic(1995); basin systems; glacialArmstrong depositionand McCourt(1997)]

Rooihoogte Breccia, Karst-fill alluvial Low High — due toconglomerate, and lacustrine long erosivesandstone and sedimentation hiatus at base ofmudrock Pretoria Group

Chuniespoort Duitschland Clastic sediments Regressive Decreasing Elevated, due to(Transvaal ) Formation; in Duitschland; shallow marine erosion ofand Ghaap Koegas mixed siliciclastic coastline underlying(Griqualand Subgroup in and chemical dolomites toWest) (ca Griqualand sediments in supply clastic2642 Ma– West Koegas particles<2.43 Ga)

Penge BIFs ca 2.5– Deeper shelf BIF High Low — due toFormation 2.43 Ga chemical lack of clastic(Transvaal ) (Beukes, sedimentation sedimentationand 1980, 1983;Kuruman Trendalland et al., 1990)GriquatownFms.(GriqualandWest)Malmani Dolomites, ca 2583–2550 Carbonate High Low — lack ofSubgroup minor chert and to 2500 Ma platform silicilastic(Transvaal ) mudrock (Altermann developed under sedimentationand and Nelson, shallow shelfCampbellrand 1998; Martin conditions inSubgroup et al., 1998) extensive epeiric(Griqualand seaWest)Black Reef Black Reef: Schmidtsdrif Black Reef: Generally High — due toFormation conglomerate– lavas: peneplanation low dominance of(Transvaal ) sandstone– 2642 Ma and fluvial sheet siliciclastics inand mudrock; ( Walraven sandstones; bothSchmidtsdrif Schmidtsdrif: and Martini, Schmidtsdrif: Schmidtsdrif andSubgroup siliciclastic and 1995) shallow marine Black Reef(Griqualand chemical coastline atWest) sediments, minor craton margin

lavas‘Protobasinal Various Lavas and ca 2.66 Ga Linear fault- Low High —units’ formations pyroclastic (Eriksson bounded basins extensive

and rocks, immature and Reczko, filled with volcanismgroups — coarse-clastic 1995) volcanic rocks, (plume)restricted and more continentalfault-related reworked fine- deposits andbasins of clastic sediments finer, deeperlocalised basinalextent sediments

opment of a number of restricted volcano- Kanye basins) (Table 1) and of the correlatedPostmasburg Group (Griqualand West basin) fol-sedimentary successions (‘protobasinal units’, a

purely descriptive term; Table 1) within the lowed. This hiatus is compatible with lowered sealevel and elevated freeboard conditions. PeriglacialTransvaal basin. Tectonic control on protobasinal

sedimentation most likely reflects a continuation detritus shed into the Timeball Hill shallow seafrom a glacial centre on the Vryburg Rise palaeo-of the extensional regime associated with the

Ventersdorp volcanism. Peneplanation and depos- high, separating the Transvaal and GriqualandWest basins, correlates with the terrestrial basalition of the Black Reef Formation (Table 1) in the

Transvaal and Kanye basins may logically have Makganyene glacial diamictites in the latter depos-itory (Table 1) (Visser, 1971; Eriksson et al.,followed upon the second Limpopo collisional

event, at ca 2.58 Ga (Treloar and Blenkinsop, 1993a). The latter glacial deposits, resting erosivelyon older rocks in the Griqualand West basin, point1995). The uppermost parts of the Black Reef

Formation reflect shallow marine inundation, as to eustatic lowering of sea level and enhancedfreeboard; the Timeball Hill epeiric sea was thusdoes the Schmidtsdrif Subgroup (Ghaap Group,

Griqualand West; Table 1). probably a local exception to this trend. Thecombination of subaerial lavas and fluvio-Carbonate platform sedimentation began in the

southwestern Prieska sub-basin (see Fig. 7, lacustrine continental sediments (Boshoek toDwaalheuwel Formations), succeeding theSection 4) of the Griqualand West basin, between

2642 and 2588 Ma (Altermann and Nelson, 1998), Timeball Hill Formation, suggests continuation oflow sea level–high freeboard conditions (Table 1).followed by a major transgression on the Kaapvaal

craton at ca 2583 Ma (Martin et al., 1998) to Renewed rifting of the Pretoria Group basin(Daspoort Formation) ushered in a second and2550 Ma. The resultant thick carbonate succession

was followed by a second major transgression after more widespread transgressive epeiric sea envi-ronment (Silverton–Magaliesberg Formations,2516 Ma (Altermann and Nelson, 1998), during

which banded iron formation (BIF ) deposits devel- Table 1), compatible with high sea level–relativelylow freeboard conditions. Such conditions areoped over much of the Kaapvaal craton (Beukes,

1980, 1983; Nelson et al., 1999) (Table 1; Fig. 7). reflected also in the BIF of the Hotazel Formationwhich succeed the Ongeluk Formation volcanicsIn general terms, low freeboard and high sea level

are indicated for the ca 2550–2430 Ma carbonate– of the Postmasburg Group in Griqualand West.Intrusion of the Bushveld Complex, at 2061±BIF epeiric trangression of the Kaapvaal craton.

The largely siliciclastic sediments of the 27 Ma [ Walraven et al. (1990); more recently givenas 2054 Ma by Armstrong and McCourt (1997)]Schmidtsdrif Subgroup and Black Reef Formation

(2642–2588 Ma) point to elevated freeboard com- terminated Transvaal Supergroup sedimentation(Fig. 3).pared to the younger carbonate–BIF platform.

Radiometric age data from the Limpopo belt, To the west of the Kaapvaal craton, theNamaqua–Natal mobile belt includes basementdiscussed previously, indicate the possibility of

compressional orogenic events on the northern rocks and supercrustal lithologies which wererelated to the ca 2.0 Ga Eburnean event. Themargin of the Kaapvaal craton during the 2642–

2430 Ma period. However, there does not seem to supercrustal rocks of the Kheis province (Fig. 1),with a maximum age of 1.93 Ga ( Walraven andhave been any influence on the northern part of

the Transvaal epeiric sea from this orogeny, as Martini, 1995), essentially fall without the timeperiod discussed here. Eburnean basement rocksthere is no evidence for significant clastic influx

into the carbonate–BIF platform succession. include the Achab Gneis in the Richtersveld prov-ince ( Watkeys, 1986), and those in the Okwa RiverFollowing a widespread hiatus, approximately

between 2430 and 2350 Ma, during which consider- area of Botswana (Thomas et al., 1993), northwestof the Kaapvaal craton (Fig. 1). The Haibable weathering and erosion of BIF and the under-

lying carbonate platform rocks occurred, Subgroup metavolcanic and metasedimentaryrocks in the Richtersveld province are thought todeposition of the Pretoria Group (Transvaal and

reflect a continental margin island arc (Colliston Formation in the Transvaal basin, Kaapvaalcraton, also thought to be >2150 Ma old. As anand Schoch, 1998), analogous in genesis to the

Birimian Supergroup of the West African craton alternative view to the depositional models givenin Table 1, the combination of pelagic suspension-(Section 2.1). The relatively high grade metamor-

phism and deformation in these rocks prevents deposited mudrocks, subaqueous volcanics, turbid-ites and subordinate carbonate rocks in theinferences being made on sea level and freeboard

changes for these terranes. Piriwiri, may also occur within epeiric marinesettings (e.g. Eriksson and Reczko, 1998), such asalso suggested for the Silverton Formation2.3. Zimbabwe craton(Eriksson, 1997). High sea level may also beinferred for the Piriwiri Group, although its higherLarge parts of the Zimbabwe craton were

affected by compressional tectonics between 2.7 metamorphic grade (e.g. Munyanyiwa andMaaskant, 1998) signals a different burial andand 2.6 Ga, caused probably by terrane accretion

from the north and a southward subduction of tectonic history for this part of the Zimbabwecraton when compared to the Transvaal basin.oceanic crust beneath the Tokwe–Zimbabwe conti-

nental crust ( Kusky, 1998). Such a tectonic regime In view of palaeomagnetic evidence (Fig. 2)that the Kalahari craton lay to the south of theis not obvious on the Kaapvaal craton at these

times (compare with Section 2.2). Only with the Tanzania and Bengweula crustal blocks during theEarly Proterozoic, Deweras rifting may well haveintrusion of the 480 km long Great Dyke at ca

2.58 Ga (Mukasa et al., 1996), did a regionally been related to the Eburnean collision betweenZimbabwe and the two northern cratonic terranes.widespread extensional tectonic regime prevail on

the Zimbabwe craton, by which time thick conti- Rifting is also proposed to have been related tothe intrusion of the Bushveld Complex into thenental crust and mature cratonic conditions had

presumably been attained. However, an approxi- Kaapvaal craton (e.g. Eriksson et al., 1991; VonGruenewaldt and Harmer, 1993) at 2054 Mamate 2.58 Ga age is also inferred for collision of

the Zimbabwe craton with the Kaapvaal–Central (Armstrong and McCourt, 1997); this was conceiv-ably also related to the ca 2.0 Ga Eburnean colli-Zone plate during the Limpopo orogeny (Treloar

and Blenkinsop, 1995) (Fig. 3). Although it is sional event. The tilting and erosion of the DewerasGroup, which preceded deposition of the subse-possible that this collision may have produced

tectonic conditions conducive to Great Dyke intru- quent Lomagundi Group (Table 2), was possiblyalso associated with this widespread orogenicsion, other workers (e.g. Barton et al., 1994;

Cheney, 1996; Holzer et al., 1998) argue for youn- event. The Lomagundi Group probably reflectsdrowning following thermal subsidence of theger Limpopo orogenic ages of 2.47 to 2.05–

1.95 Ga. Deweras rifted basin (Stagman, 1981; Master,1991). Continuation of the Magondi orogeny untilThe Zimbabwe craton lacks rocks equivalent to

the lower and middle Transvaal Supergroup (ca ca 1.8 Ga resulted in S-SE directed thrusting ofthe Piriwiri over the Lomagundi Group, and of2600 to >2222 Ma carbonate–BIF and mudrock–

tilloid succession). Such sediments were either not the latter onto the Deweras Group (Stagman,1981; Master, 1991). Master (1993) suggests thatdeposited on the Zimbabwe craton, or were eroded

before the Magondi Supergroup was laid down, the Magondi, Okwa and Kheis terranes (Fig. 1;Section 2.2) formed as a result of an Eburnean-thus reflecting a high freeboard condition at some

time between 2600 and 2200 Ma. Thermal uplift aged destructive plate margin developed to thenorthwest of the Kalahari craton.related to Great Dyke intrusion and concomitant

rifting presumably contributed to this high free-board. The Magondi Supergroup overlies this late 2.4. Southwestern Congo cratonArchaean basement, and comprises three groups(Table 2). The Piriwiri lithologies are analogous The only 2.7–2.0 Ga succession here, the

Oendolongo Supergroup (Fig. 1), rests uncon-to the epeiric marine deposits of the Silverton

Table 2Magondi Supergroup, Zimbabwe Craton — geological overview (data from: Stagman, 1981; Cahen et al., 1984; Leyshon and Tennick,1988; Master, 1991, 1995)

Supergroup Group Formation/facies Lithology Sedimentary Sea level Freeboardhistory

Magondi Lomagundi Mature Marine High Low(<2.0 Ga) quartzose coastline —

sandstones, carbonatecarbonate rocks, platform andmudrocks and stable shelfsubordinatevolcanic rocks(3475 m)

Deweras Southern facies Arkose Braided streams Low High — plume(∏2060±100 Ma) (mudrock) and volcanism and

(800 m) volcanism — riftingstrike–slip faultsystems

Tholeiitic lavas(agglomerate)(1000 m)Immature clasticsediments (thin)

Northern facies Siliciclastic Volcanism Low High —sediments, mafic volcanismlavas (plume) and rift(pyroclastics) setting(600 m)Calcareous — Playa lake andevaporitic flat, locallymudrocks aeolian(sandstones)(500 m)Arkosic red beds Distal fanandconglomerate(200 m)

Piriwiri Chitena Graphitic Deep euxinic — High Low(>2.15 Ga) phyllite, chert, slope abyssal fan

black shale and plain oreugeosynclinalflysch

Kanyaga Phyllite, wacke,micaceous,quartzite, minorcarbonate rocks

formably on a granitic and migmatised basement be Late Archaean in age (Carvalho, 1972; Silva,1977; Bassot et al., 1981). The succeedingof at least 2600 Ma age (Cahen et al., 1984). The

Oendolongo Supergroup comprises a basal pre- Chivanda Group follows unconformably, andlocally exhibits a succession of schists (possiblydominantly volcano-sedimentary Jamba Group,

which includes BIF, localised carbonate rocks and originally tuffs), medial sandstones and conglomer-ates, and upper black schists (Torquato and Tomasclastic metasediments, and which is considered to

Oliviera, 1977). The Chivanda Group underwent 2078±36 Ma being obtained on one sample(Cahen et al., 1984) (Fig. 3).a regional metamorphic event at 2161±99 Ma and

a thermal event at 2063±56 Ma (Bassot et al.,1981; Cahen et al., 1984). 2.6. Northwestern Congo craton (North Gabon

Massif)The Jamba Group, including BIF and localisedcarbonate rocks, and which appears to be LateArchaean in age, is possibly analogous with the There are two major 2.7–2.0 Ga volcano-sedi-

mentary successions within this terrane: thelower transgressive carbonate–BIF succession(Campbellrand–Malmani Subgroups) of the Kimezian Supergroup in the south of the NW

Congo craton (Fig. 1), and the lithologies associ-Kaapvaal craton (Fig. 3). The black schists inthe upper part of the Chivanda Group, ated with the Gabon orogenic belt in the north of

the NW Congo cratonic area (Petters, 1991). The>2161±99 Ma in age, are comparable in generallithology and lower age limit with the >2150 Ma ca 2600–2088±60 Ma Kimezian Supergroup origi-

nally comprised a shelf association of mudrocks,Silverton Formation (Pretoria Group, TransvaalSupergroup, Kaapvaal craton) transgression and quartzose sandstones and limestones which were

metamorphosed to amphibolite facies; furtherthe probably epeiric marine >2150 Ma PiriwiriGroup (Zimbabwe craton) (Fig. 3). alteration occurred during the Late Proterozoic

Western Congolian orogeny, which produced mig-matites and gneisses (Cahen et al., 1984). The2.5. Kasai Shield, southeastern Congo cratonpoorly dated Kimezian is possibly analogous tothe general depositional setting envisaged for theVolcano–sedimentary stratigraphic units within

the 2.7–2.0 Ga interval in this region (Fig. 1) are Late Archaean–early Palaeoproterozoic LowerTransvaal ( Kaapvaal craton)–Jamba Group (SWlimited to the Luiza Supergroup, and, possibly,

also to the Lulua Group (Petters, 1991). The Congo craton)–Luiza Supergroup ( Kasai Shield)epeiric marine assemblages, which also included<2560 Ma and >2432±48 Ma Luiza Supergroup

unconformably overlies basement tonalites and BIF (Fig. 3).A more intense tectono-thermal collisionalgranodioritic gneisses (Cahen et al., 1984), and

consists of quartzites, mica–schists and BIF event in the north of the NW Congo cratonproduced the Gabon orogenic belt, which com-(Petters, 1991). The presence of BIF and the

apparent age of the Luiza indicate the possibility prises, from west to east, Archaean basementrocks, the metamorphic rocks of the Ogooueof a widespread occurrence of raised sea level and

decreased continental freeboard in parts of Africa Supergroup, in turn thrust eastwards over theforeland of the Francevillian Supergroup (Ledruat the Archaean–Proterozoic boundary, encom-

passing also the carbonate–BIF succession of the et al., 1989). The Francevillian Supergroup(Table 3) unconformably overlies 2.7–3.3 Ga gra-lower Transvaal Supergroup on the Kaapvaal

craton, and the Jamba Group in the SW Congo nitic basement rocks. The inferred drowning eventconcomitant with pyritic–asphaltic sandstonecraton (Fig. 3).

The Lulua Group is generally fault-bounded deposition in the upper part of FA, at ca>2143 Ma, may be analogous to that interpretedand in faulted contact with the Luiza rocks, and

comprises ca 6000 m of clastic metasedimentary for the shelf-related black mudrocks of the>2150 Ma Silverton Formation (upper Transvaalrocks and greenstones, the latter including spillitic

basalts and pillow lavas (Petters, 1991). The Lulua Supergroup, Kaapvaal craton), the graphitic phyl-lites of the >2150 Ma Piriwiri Group (Zimbabwehas been interpreted as either unconformable upon

the Luiza Supergroup, or as a foreland equivalent craton) and the black schists of the >2161 MaChivanda Group (SW Congo craton) (Fig. 3).onto which the Luiza was thrust (Cahen et al.,

1984). Available geochronology mostly gives Crustal convergence and crustal shortening aregenerally proposed for the Gabon orogenic belt atMiddle to Late Proterozoic whole rock Rb–Sr

ages for the Lulua, with a possible age of ca 2.0 Ga (e.g. Shackleton, 1986; Ledru et al.,

Table 3The Francevillian Supergroup, NW Congo craton: geological overview (data from: Weber and Bonhomme, 1975; Bonhomme et al.,1982; Cahen et al., 1984; Gauthier-Lafaye, 1986)

Supergroup Sub-units Lithology Sedimentary history Sea level Freeboard

Francevillian FE Sandstones Possibly littoral Decreasing? Increasing?(>±2.14 Ga– (interbedded<2050 Ma) mudrocks)

FD Ignimbritic tuffs Marine shelf High LowBlack mudrocks

FC Banded chert and Marine shelf High Low (chemical shelfmassive dolomite sediments)(±50 m)

FB (±600 m) Ma-carbonate and BIF Marine shelf High Low (chemical shelf( lower part sediments)synchronous withN’Goutou)

Calcareous blackmudrocks (sandstones)Submarine maficvolcanics and sandypelites

N’Goutou Complex Subvolcanic rocks Subvolcanic ? ?(2143±143 Ma) (pyroclastic) ComplexFA Coarse feldspathic Continental Initially low, Changing, relatively,

sandstones; lower sedimentation in followed by from higher toevaporitic red beds active tectonic transgression lowerunconformably setting; upperoverlain by pyritic- drowning.asphaltic sandstoneswith U deposits(∏1000 m)

1989), and palaeomagnetic data support an > 2.05 Ga (Cahen et al., 1984) (Fig. 3). Postulateson sea level and freeboard for the UsugaranEburnean collision between the Sao Francisco and

Congo cratons at this time (Onstott and Supergroup in the mobile belt of the same name(Figs. 1 and 2) are similarly precluded by a lackHargraves, 1981; D’Agrella-Filho et al., 1996).of chronological data (Cahen et al., 1984) (Fig. 3)and by widespread high grade metamorphism.2.7. NE Congo craton and the Tanzania cratonPetters (1991) and Lenoir et al. (1994) discuss thegeological history of the Eburnean collisionalThe 2.7–2.0 Ga volcano-sedimentary record in

this region encompasses metamorphosed rocks events in this general region.within the Ubendian, Usugaran and Ruwenzorimobile belts (Figs. 1 and 2). High grade metavol-canic and metasedimentary rocks which character- 3. The 2.7–2.0 Ga volcano-sedimentary record of

Indiaise the Ubendian belt preclude estimation of sealevel and freeboard conditions. The Ruwenzoribelt (Fig. 1), encompassing the predominantly vol- 3.1. Introductioncanic Buganda-Toro Supergroup (Tanner, 1970,1973), and the analogous, east–west trending Extensive Deccan lavas and younger sediment

cover, as well as the paucity of accurate chrono-Luhule-Mobisio Group in the NE Congo craton(Fig. 1) are poorly constrained in age: <2.5 and logical data are problems inherent in studying the

Precambrian evolution of India. The South Indian up early Precambrian India (e.g. Paliwal, 1998).As ‘shield’ was defined as a component or nucleus(SIB) and North Indian (NIB) Blocks were amal-

gamated along a linear orogenic belt (the Central of a craton in Section 2.1 of this paper, consistencyrequires that this Indian usage be excluded here.Indian tectonic zone, or CITZ) (Fig. 4) and com-

prised the united SIB–NIB terranes at 2.0 Ga; The SIB and NIB differ from each other in manysignificant aspects (Table 4) and there is generalother crustal fragments probably accreted later. It

should be noted here that many Indian authors agreement about their different cratonic evolution.Controversy, however, exists about the timinguse the term ‘Peninsular Indian Shield’ for the

combined cratons and mobile belts which make of their amalgamation (Radhakrishna, 1983;

Fig. 4. Crustal structure of Pensinsular India: three early Precambrian crustal provinces separated by linear rifts in the SIB, and theAravalli–Bundelkhand province of the NIB; SIB and NIB separated by the CITZ, also a rift zone. See Table 5 and Fig. 5 for ages.

Table 4Significant differences between the SIB and NIB [some data from Sivaraman and Odom (1982) and Sinha Roy (1988)]

Aspects SIB NIB

Dharwar Singhbhum Bastar Aravalli–Bundelkhand

Gravity contour Open ClosedInitial rifting ca 3.0 Ga ca 3.3 Ga ca 3.2 Ga ca 2.5 GaClosing of initial rift basin ca 2.5 Ga ca 2.1 Ga ca 2.5 GaBanded magnetite quartzite Substantial Moderate Substantial InsignificantBanded hematite quartzite Substantial Substantial Substantial AbsentInferred Wilson cycle operation No No Yes YesGranitoid intrusion events 3.0, 2.5 Ga 3.3, 3.1, 2.1 Ga 2.5, 2.2 Ga 2.9, 2.6, 2.0 GaArchaean K-rich granite intrusion None 2.6 GaInferred mantle plume upwelling Middle–late Archaean Middle Archaean– — Early Proterozoic

Early ProterozoicBasement thickness Moderate (38 km) Moderate (48 km) Moderate (42 km) Very high (82 km)Carbonate deposits Moderate Low Moderate HighDominant structural trend North–South Northeast–Southwest

Radhakrishna and Naqvi, 1986; Radhakrishna and to cratonisation having been completed by ca 3.1–3.0 Ga in Singhbhum. Immediately overlying theRamkrishnan, 1988).

Prior to 2.0 Ga, there were four crustal prov- Singhbhum basement is the ca <3.3 and ca>3.1 Ga (Singh, 1998) Older Iron Ore Groupinces or cratons, three of them in the SIB and the

other in the NIB (Fig. 4). All of them are typi- (Fig. 5).cal Archaean tonalite–trondhjemite–granodiorite The 2.7–2.0 Ga volcano-sedimentary record in(TTG)–greenstone terranes. TTG basement Singhbhum (Table 5) begins with the Lategrowth continued in the SIB up to ca 3.0 Ga, while Archaean Younger Iron Ore Group (Fig. 5), char-in the NIB crustal stabilisation was achieved at ca acterised by haematitic iron formations (Table 5).2.5 Ga. All four crustal provinces are characterised Shallow marine to shelf deposition under generalby greenschist to lower amphibolite facies meta- highstand conditions is inferred for these rocksmorphism, with local granulites occurring in the and a combination of raised sea level and loweredNIB (Sharma, 1995; Dasgupta et al., 1997). In the continental freeboard would logically have beenSIB, early Precambrian volcano-sedimentary suc- pertinent. The presence of inferred glacigenic rockscessions remain confined to the Archaean in the (field work in progress, second–fourth authors),Bastar and Dharwar provinces; only in Singhbhum evidence for higher freeboard-type continental sed-do they transgress the Archaean–Proterozoic imentation and plume-related subaerial volcanismboundary. In the Aravalli–Bundelkhand province in the Dhanjori Formation (Table 5) suggest pos-(NIB), on the other hand, the supracrustal record sible correlation with the global ca 2.4–2.2 Gaappears to have been initiated only at ca 2.0 Ga glaciation event, which Aspler and Chiarenzelli(Table 5; Fig. 5). (1998) also relate to plume uplift of continental

crust.3.2. Singhbhum crustal province A punctuated Early Proterozoic transgression

(Fig. 5) is possible for the Chaibasa Formation–Dhalbhum Formation–Kolhan Group successionThe oldest rocks in India [ca 3.6 Ga (Pb isotope)

basement rocks; Saha and Roy (1984); (Table 5). Common slumps and slides within theDhalbhum Formation point to active tectonism,Bhattacharya (1998)] and the most extensive Late

Archaean–Early Proterozoic volcano-sedimentary probably related to thermally (mantle plume)enhanced freeboard. The Kolhan and Chaibasarecord are found in the Singhbhum province.

Potassium-rich (3.22%) granodiorites, with a high deposits both reflect higher sea level conditions,while the Dhalbhum points to a truncation of thisinitial 87Sr/86Sr ratio of 0.711 (Saha, 1994), point

Table 5Late Archaean–Early Proterozoic stratigraphy of the Indian craton (data from: Heron, 1953; Roy and Paliwal, 1981; Sarkar, 1982;Gupta et al., 1985; Srinivasan and Ojakangas, 1986; Bhattacharyya et al., 1988; Chadwick et al., 1988; Deb et al., 1989; Deb andSarkar, 1990; Ghosh and Chatterjee, 1990, 1994; Roy, 1990; Ramakrishnan, 1990; Ahmad and Rajamani, 1991; Alvi and Raza, 1992;Bhaskar Rao et al., 1992; Dasgupta et al., 1992; Ghosh et al., 1992; Roy et al., 1993; Sarkar et al., 1993; Sinha-Roy et al., 1993;Ahmad and Tarney, 1994, Saha, 1994; Bandyopadhaya et al., 1995; Mazumder, 1996; Ray et al., 1996; Bose et al., 1997; Senguptaet al., 1997; Singh, 1997)

Cratonic Stratigraphy Lithology Sedimentary Historyprovince

Singhbhum Kolhan Group Lower arkose– Fluvial ( lower) tocrustal (>2.1 Ga) quartzite – mudrock shallow marineprovince (conglomerate) (upper), in fault-

succession and upper bounded basincarbonate succession

Dhalbhum Formation Predominant Thermal (plume)(unconf.) bimodal volcanics elevation, volcanism,

and immature fluvial/aeolianquartzites (finer sedimentationsiliciclastics)

Chaibasa Formation Mature quartzites Tectonic quiescence(unconf.) and interbedded and transgressive

mudrocks, minor marine settingvolcanics

Dhanjori Formation Basal glacigenic Glacial–(unconf.) boulder bed, fluviolacustrine;

conglomerates and plume elevationimmature arkoses,common volcanicrocks (komatiite–basalt)

Younger Iron Ore Hematitic BIF, Shallow marine toGroup (>2.55 Ga mudrocks, tuffaceous shelf depositionand <3.0 Ga) shale, chert,

sandstone (dolomiteand conglomerate);komatiitic–basalticvolcanics

Dharwar Dharwar Supergroup Chitradurga Group Hiriyur Fm Mudrocks, chert Relatively deepcrustal (ca 3.0–2.5 Ga) (conglomerate) marineprovince

Ingaldhal Fm Chert, Mn and Fe Deeper shelf–marinedeposits, bimodal basinvolcanics(pyroclastics)

Vanivilas Fm Carbonate rocks Marine shelf(mudrocks)

Bababudan Group Mundre Fm Wackes, mudrocks Transgressive marine(conglomerates) shelf conditions

Jagar FmMulaingiri Fm BIF (volcanics,

mudrocks)Santavery Fm Bimodal volcanic Thermal elevation,

rocks continentalvolcanism and fluvialsedimentation

Table 5 (contined )

Cratonic Stratigraphy Lithology Sedimentary Historyprovince

Allampur Fm Mafic–ultramafic sillsand quartzites

Kalaspura Fm Oligomictconglomerates,quartzites, maficvolcanics andpyroclastics(subaerial )

Bastar (ca 2.95–2.6 Ga) Nandagao Group Bimodal volcanics Volcanism andcrustal (pyroclastics) crustal thickeningprovince

Bailadila Group BIF, chert, mudrocks Marine(coarse siliciclastics) transgression–shelf

sedimentationBengpal Group Conglomerate– Regressive(unconf.) quartzite–schists and continental

intercalated depositionmetabasites

Aravalli– (Aravalli Supergroup,Bundelkhand ca 2.0–1.75 Ga;crustal Bijawar and Gwaliorprovince basins, Bundelkhand

craton, ca 2.0–1.4 Ga)

transgressive trend in the Early Proterozoic of the formations of the Bababudan Group (Fig. 5). TheChitradurga Group of the Dharwar SupergroupSinghbhum crustal province (Fig. 5; Table 5).

Anorogenic granitic magmatism at ca 2.1 Ga ter- begins with shelf deposits (Table 5), thus indicat-ing a further continuation of high sea level, loweredminated Proterozoic deposition in Singhbhum,

ushering in an enormous stratigraphic hiatus there freeboard conditions (Chadwick et al., 1988), fol-lowed by deeper shelf to possibly marine basinuntil the Tertiary (Saha et al., 1988).conditions (Bhattacharyya et al., 1988; Naqviet al., 1988). Stromatolitic carbonates of the3.3. Dharwar crustal provinceVanivilas Formation at the base of the ChitradurgaGroup (Fig. 5) appear to be deep shelf depositsThe Middle Archean TTG–greenstone base-

ment of Dharwar has an extensively developed (Naqvi et al., 1988), thus indicating greater waterdepths than normally inferred for shelf carbonates.regolith (Fig. 5), indicating exposure and most

likely elevated freeboard at ca 3.0 Ga, upon which Naqvi et al. (1988) suggest an active continentalmargin setting for the Chitradurga argillites. Thethe Late Archaean Dharwar Supergroup was

deposited unconformably (Chardon et al., 1998) Late Archaean volcano-sedimentary record in theDharwar crustal province was terminated by gra-(Table 5). Low sea level and high freeboard condi-

tions appear to have continued into the basal part nitic intrusions at ca 2.5 Ga (Friend and Nutman,1991) (Fig. 5), followed by a large stratigraphicof the lower Bababudan Group of this unit

(Table 5; Fig. 5). Reduction in freeboard and gap until the Middle Proterozoic. Peucat et al.(1989) suggest that this 2.5 Ga event in Dharwarrising sea levels are inferred for the upper three

Fig. 5. Geohistory summary chart for the 2.7–2.0 Ga volcano-sedimentary successions of the Singhbhum, Dharwar, Bastar andAravalli–Bundelkhand crustal provinces, Peninsular India. See Table 5 for sources of data and text for details. Note, Late Archaeancarbonate–BIF-(siliciclastic) association probably related to transgressive drowning of cratonic terranes, and alternation of high sealevel and high freeboard deposits in the Early Proterozoic period.

included accretion of ca 2.5 Ga crust in a subduc- ered sea level and enhanced freeboard. Subsequenttransgressive marine sediments of the Bailadilation setting.Group may be the approximate correlates of theLate Archaean Younger Iron Ore Group in

3.4. Bastar crustal province Singhbhum and of the uppermost Bababudan andsucceeding Chitradurga BIF in Dharwar (Fig. 5).

The TTG–greenstone basement in Bastar is Marine sedimentation in Bastar was terminated byunconformably overlain by a succession of three the Nandagao Group, an enormous thickness ofvolcano-sedimentary groups of Late Archaean age volcanic rocks possibly related to crustal thicken-(Table 5), terminated by anorogenic granite intru- ing of the Bastar craton (Mahadeven, 1998). Twosion at ca 2.6–2.5 Ga (Sarkar et al., 1993) (Fig. 5). further supracrustal volcano-sedimentary unitsThe Bengpal Group is ascribed to regressive depos- occur in the Bastar crustal province, but their

stratigraphic constraints are uncertain (Fig. 5).ition (Ramakrishnan, 1990), thus suggesting low-

3.5. Aravalli–Bundelkhand crustal province possibility of high freeboard-lowered sea level con-ditions for this region.

This comprises the Aravalli craton, lying to thenorthwest and separated from the Bundelkhandcraton to the southeast by younger rocks (Fig. 4). 4. The 2.7–2.0 Ga volcano-sedimentary record ofIt is thought that the Bundelkhand craton may Australiarepresent the central part of a much more wide-spread Archaean Bundelkhand province, encom- 4.1. Introductionpassing the Aravalli region and extending also

The Australian continent comprises differentfurther northwards to form the basement to theArchaean crustal blocks including the PilbaraLesser Himalayan belt (Sharma, 1998). Thecraton, the Yilgarn craton and the Gawler cratonArchaean basement of the Bundelkhand craton,(Fig. 6), and Proterozoic trough sediments andcomprising quartzites, BIF, schists and amphibo-volcanics. The Pilbara craton stabilised at >3.0 Galite, engulfed in TTG intrusives, is thought to be(Nelson et al., 1999), the Yilgarn crust at ca 2.7–comparable to the Banded Gneiss Complex of the2.6 Ga (Nelson, 1998) and the Gawler crust at caAravalli craton (Sharma, 1998). The greater2.5 Ga (Daly and Fanning, 1990). The YilgarnBundelkhand crustal province thus appears to becraton, however, yields the oldest known detritala typical Archaean cratonic terrane. Granite intru-zircons, of 4.28 Ga (Myers, 1995). Late Archaeansions into Aravalli craton basement rocks at casedimentary cover rocks (2.7–2.5 Ga) are wide-2.5 Ga (Fig. 5) were followed by the deposition ofspread only on the Pilbara craton. Earlythe Aravalli Supergroup (Table 5), a successionProterozoic cover rocks are more common andcommonly held to be Palaeoproterozoic in agecan be traced on all three cratons and in some(e.g. Bhattacharya, 1998). However, Pb isotopelinear belts or inliers of central and northerndata (Deb et al., 1989; Deb and Sarkar, 1990)Australia, such as the 2.1–1.9 Ga Pine Creek inlier,indicate clearly that the Aravalli Supergroup is caor the Tennant Creek inlier (Fig. 6), both with2.0–1.75 Ga (Sinha Roy, 1988; Banerjee andturbiditic and BIF, deep water, isoclinally foldedBhattacharya, 1994; Verma, 1998), and it thus fallssediments. The Gawler craton experienced threeoutside the compass of this paper.deformation episodes with a peak deformation andIn the Bundelkhand craton to the southeast,metamorphism during the ca 2450 Ma Sleafordianthe 2.7–2.0 Ga record is more substantial butorogeny, followed by granitic intrusions at 2440–includes no significant volcano-sedimentary suc-2300 Ma (Daly and Fanning, 1990). Sediments ofcessions. Cratonisation here became complete with2.5–2.0 Ga are not known from the Gawler cratonlarge scale granitoid intrusion and associated rhyo-and were either never deposited, or eroded awaylitic eruptives, possibly related to micro-platebefore the N–NE trending, ca 1.95–1.85 Gaaccretion (Rahman and Zainuddin, 1993); geo-(Fanning et al., 1988) Hutchison Group was laidchronology indicates that this felsic magmatismdown.was Palaeoproterozoic in age, and data suggest

2.5–2.1 Ga or a narrower time frame of 2.5–2.4 Ga(Sarkar et al., 1984; Sharma, 1998). Hydrothermal 4.2. Yilgarn cratonquartz veining on a large scale, dated at ca 2.0 Ga(Sharma, 1998), preceded deposition in the The Yilgarn craton comprises high grade

gneisses and associated lithologies of the WesternBijawar and Gwalior basins, and their volcano-sedimentary successions developed marginally to Gneiss terrain, and an eastern low grade granitoid–

greenstone terrain (Myers, 1997). The Westernthe Bundelkhand craton at some time between 2.0and 1.4 Ga (Sharma, 1998). Sharma (1998) pro- Gneiss terrain contains an old epicontinental to

continental metasedimentary sequence with detri-poses a compressional tectonic regime for theBundelkhand craton during the 2.7–2.0 Ga period tal zircon ages between 3.3 and 4.28 Ga (Narryer

Gneiss Complex), and has a conspicuous lack ofconsidered in this paper, thereby suggesting the

volcanic intercalations (Myers, 1995). It thus pro- the Great Dyke of Zimbabwe, but significantlyyounger at 2.42 Ga than their African counterpart.vides evidence for the earliest formation of conti-

nental crust and, concomitantly, of high freeboardconditions. The eastern terrain generally has youn- 4.3. Pilbara cratonger igneous ages, <3.0 Ga, and the metasedimentsare associated with basic volcanics and are of The Pilbara craton exhibits a widespread and

thick late Archaean to Middle Proterozoic sedi-typical greenstone affinity. The major tectono-thermal event througout the Yilgarn craton, mentary cover (Table 6; Fig. 8). Lithological sim-

ilarities to the Ventersdorp and Transvaalincluding the Narryer Gneiss Complex, is dated atca 2650 Ma (Myers, 1995, 1997). Late Archaean Supergroups of the Kaapvaal craton (Fig. 7) led

in the past to several comparisons of the develop-cover sediments have yet to be identified on theYilgarn craton. ment of 2.7–2.0 Ga basins on the two cratons

(Trendall, 1968, Button, 1976), culminating in theThe northern Yilgarn craton exhibits 2.2–1.6 Ga lithologies in two sedimentary cycles within suggestion of a joint ‘Vaalbara’ continent (Cheney,

1996). However, as discussed by Nelson et al.the Nabberu basin. This basin contains thick,Superior-type granular and pelletal BIF in pre- (1999), with the accumulation of more precise

isotopic age data, such correlations and palaeogeo-dominantly shallow water facies. Low to moderatecontinental freeboard can be interpreted for the graphic reconstructions become untenable and the

differences of the basinal development and theNabberu basin sediments, which include the shelfsediments of the Glengarry (2.2–1.8 Ga) and over- sedimentary record between the Kaapvaal and

Pilbara cratons become striking and geneticallylying Earaheedy Groups (ca1.8–1.7 Ga). TheGlengarry Group sediments may be a lateral, deep important. Microtectite layers preserved in both

the Pilbara and Kaapvaal successions ( Woodheadwater facies equivalent of the Wyloo Group (seebelow, Pilbara craton) of the central Capricorn et al., 1998; Simonson et al., 1999) may provide

tie-lines (Fig. 7) in attempted detailed correlationsorogen, where the Yilgarn and Pilbara cratonsamalgamated at ca 2.0–1.8 Ga. The central Yilgarn of the early Precambrian supracrustal rocks of

these two cratons. However, palaeomagnetismcraton exhibits Early Proterozoic intrusive rocksthat are genetically and morphologically similar to does not support contiguity.

Local uplift and erosion were active during thedeposition of the Fortescue Group (Table 6;Fig. 8), and a high to moderate continental free-board must be assumed. The Fortescue Groupformed during a period of extensional, continentaltectonics that culminated in deep shelf-type basinson cratonic basement, in which continuous chemi-cal sedimentation with very little clastic distur-bance took place (overlying Hamersley Group).The sediments of the Hamersley Group (Table 6;Fig. 7) show no significant continental influx apartfrom suspension-deposited mudrocks, and there-fore indicate low continental freeboard. The>2400 m thickness and 180 Ma duration ofHamersley Group sedimentation thus points to along period of very low continental freeboardconditions for large parts of the Pilbara craton.The lowering of continental freeboard was proba-Fig. 6. Distribution of the Archaean Pilbara, Yilgarn andbly initiated during the Jeerinah FormationGawler cratons of Australia and of some Early Proterozoic

folded sedimentary inliers, as discussed in the text. (Fig. 8) volcanism of the Fortescue Group,

Table 62.7–2.0 Ga volcano-sedimentary record of the Pilbara Craton, Australia. (Data from: Trendall, 1976, 1981, 1983, unpublished; Thorne1983, 1986; Walter, 1983; Arndt et al., 1991; Simonson et al., 1993; Cheney, 1996; Barley et al., 1997; Trendall et al., 1998; Martinet al., 1999; Nelson et al., 1999)

Group Formation Lithology Geological history Sea level Freeboard

Wyloo Clastic sediments Detailed Probably (relatively) Increasing from low(2209±15– and basalts investigations not low (?) to moderate (except1900 Ma) available; upper for uppermost part)(unconformable) shallowing-upward

sequence includingstromatoliticcarbonates

Turee 5000 m; Sandstones, Uncertain; Lowered sea level (?) Cessation ofCreek wackes and subsidence of S. Hamersley chemical(<2449±3 Ma siltstones (carbonate margin of preceding sedimentation andand rocks); diamictite of Hamersley basin onset of Turee Creek>2209± Meteorite Bore and/or uplift of clastics points to15 Ma) Member in basal Pilbara block; raised freeboard(conformable) formation diamictite and

carbonates probablyshallow-subaqueous

Hamersley Boolgeda BIF Deep shelf chemical High Low–very low(>2400 m sedimentationthick)(conformable) Woongarra Felsic volcanics Subaqueous High

(2449±3 Ma) volcanismWeeli Wolli BIF Deep shelf chemical High

sedimentationBrockman Iron Yandicoogina Mbr.- Shallower shelf Relatively lower-Formation shales sedimentation shallower shelf

Joffre Mbr.-BIF Deep shelf chemical Highsedimentation

Whaleback Mbr.- Shallower shelf Shallow shelf-shales sedimentation relatively lowerDales Gorge Mbr.- Deep shelf chemical HighBIF sedimentation

Mount McRae Mudrocks Shallower shelf Relatively lower–sedimentation shallower shelf

Mount Sylvia BIF Deep shelf chemical Highsedimentation

Wittenoom Dolomites Gravity flows below Platform or shelffair weather wave setting, that is,base - clastic relatively shallowercarbonates fromArchaean Carawinecarbonate platform

Marra Mamba BIF Deep shelf chemical High(2597±5 Ma) sedimentation

Fortescue Basalt, lesser Extensional tectonics Low Generally high;(2775±10– komatiitic basalts, and probably plume- substantial lowering2629±5 Ma) felsic lavas and clastic related continental in upper Fortescue

sediments; 7000 m in volcanism; due to gradationaltotal; several sediments= upper contact withunconformities and lacustrine and Hamersley Grouplocal disconformities alluvialdo not encompassmajor time gaps (seeFig .8)

Fig. 7. Schematic stratigraphic profile through the Hamersley Group, Pilbara craton, Australia ( left-hand column). Also shown isthe upper, Jeerinah Formation of the underlying Fortescue Group and the top contact of the Hamersley with the overlying TureeCreek Group. On the right-hand side, schematic profiles through the carbonate–BIF platform successions of the Transvaal Supergroup,Kaapvaal craton are given for the Prieska and Ghaap Plateau sub-basins (Ghaap Group) of the Griqualand West basin and for theTransvaal basin (Malmani Subgroup dolomites and overlying BIF, all of the Chuniespoort Group; underlying Black Reef Formation).On the right-hand side of each Transvaal profile, inferred second order transgression–regression cycles of the Transvaal Supergroupare compared to cycles interpreted according to the same principles (sensu Altermann and Nelson, 1998) for the Hamersley Group(column A). Column B depicts cycles and megasequence subdivision of the Hamersley succession after Blake and Barley (1992). Thevertical scale used throughout is the age of the formations in Ma. Note that the boundary between the Chichester Range Megasequenceand the Hamersley Range Megasequence falls within the continuous deep shelf sedimentation of the Mount McRae Shale and isthus not definable. The transgression–regression cycles depicted are obviously of different orders (150 and <10 Ma duration).Transgressive–low freeboard–deep shelf conditions can be deduced from the curves for the Hamersley and Prieska columns between2550 and <2450 Ma, and for the Ghaap Plateau and Transvaal columns between ca 2500 and 2450 Ma. The transgressive depositionof the Jeerinah Formation over older Fortescue Group rocks (Hamersley basin) preceded the transgressive deposition of the Vryburgrocks on the Kaapvaal craton by around 50 Ma (Prieska and Ghaap Plateau sub-basins) and deposition of upper Black Reeftransgressive rocks (Transvaal basin) by probably significantly >50 my. The deposition of BIF and carbonate sediments on theKaapvaal and Pilbara cratons is thus only broadly simultaneous (Nelson et al., 1999). The listed radiometric ages are on zircons,except for the Pb–Pb age that is directly on carbonates: A, Arndt et al. (1991); Al, Altermann and Nelson (1998); B, Barley et al.(1997); M, Martin et al. (1998, 1999); N, Nelson et al. (1999); S, Woodhead et al. (1998); T, Trendall et al. (1990); T2, Trendallet al. (1995); T3, Trendall et al. (1998); W, Walraven and Martini (1995).

ble with the Makganyene Formation glacial depos-its of the Transvaal Supergroup, Kaapvaal craton(Section 2.2) (Trendall, 1976). Trendall (1981,1983) emphasises that the conformably basedTuree Creek Group must be seen as the terminaldeposits of the Hamersley basin, characterised bythe cessation of chemical deposition and thesudden supply of terrigenous material, causedprobably by subsidence of the southern margin ofthe basin. This may conceivably also be viewed asreflecting uplift of the Pilbara block relative to sealevel, to expose the source area of the Turee Creekclastic sediments to erosion. Such an interpretationwould support elevated freeboard conditions forat least the central part of the Pilbara cratonduring Turee Creek time (<2449±3 Ma). Thepoorly studied Wyloo Group (Fig. 8; Table 6),which unconformably overlies the Turee CreekGroup, suggests progressively increasing freeboardconditions (Table 6).

5. Discussion

Early Precambrian supercontinents are difficultto recognise, and their reconstruction relies onpalaeomagnetic data (often problematic), struc-tural, lithological and mineralisation trends (e.g.Aspler and Chiarenzelli, 1998). Hoffman(1989,1992) postulated a ca 2.8–2.6 Ga NorthFig. 8. Schematic stratigraphic column for the Pilbara craton,

Australia, showing the Fortescue, Hamersley, Turee Creek and American supercontinent, Kenorland, which maylower Wyloo Groups, as discussed in the text; not to vertical also have included the Baltic and Siberian shieldsscale. Note the continuity of the sedimentary section from (Aspler and Chiarenzelli, 1998). The latter authors2690 Ma at the base of the Jeerinah Formation to the top of

also proposed the ‘southern’ supercontinent dis-the Turee Creek Group, for at least 250 Ma. Unconformitiescussed here, but supporting evidence remains tenu-within the Fortescue Group are of minor importance, covering

relatively short time spans. The disconformity above the Kylena ous. Arndt (1999) discusses widespread continentsBasalt is questionable and developed probably only on a very as early as 3.0 Ga, and extensive Archaean con-local scale. The radiometric ages are on zircons except for the glomerates with tonalitic basement clasts on thePb–Pb ages, directly on carbonates. A, Arndt et al. (1991); B,

Zimbabwe craton also suggest early continentalBarley et al. (1997); M, Martin et al. (1998, 1999); S, Woodheadcrustal growth (e.g. Petters, 1991). Auriferous con-et al. (1998); T, Trendall et al. (1990); T2, Trendall et al. (1995);

T3, Trendall et al. (1998). glomerates are a common Archaean facies (e.g.Srinivasan and Ojakangas, 1986), with gold proba-bly derived from reworked greenstones. The sizebecause the transition to the Hamersley Group is

conformable (Table 6). of the Witwatersrand gold-bearing basin and thestability of the Kaapvaal craton during its depos-The ca 2.4–2.2 Ga Turee Creek Group (Fig. 8)

includes the conspicuous diamictite of the ition (ca 2.97–2.78 Ga; Robb and Meyer, 1995)are compatible with a central supercontinentalMeteorite Bore Member (Table 6) of the basal

Kungarra Formation. The latter may be correlati- location.

The Congo and East Saharan cratons appear suggests a similar regime. This evidence for riftingis compatible with attenuation and dispersal of ato have formed a large unified crustal plate in the

Late Archaean (Figs. 1 and 2). Amalgamation of ‘southern’ Late Archaean supercontinent. In con-trast to this idea, large scale granite intrusion andthe SIB and NIB (Fig. 4) may predate 2.5 Ga

( Kale, 1998), supporting a Late Archaean super- final stabilisation of many of the cratons of Africapersisted between 2.7 and 2.0 Ga (Section 2), andcontinent possibly centred on India (Rogers,

1993). Comparison of Kaapvaal’s Transvaal a major tectono-thermal event is recorded in theYilgarn craton at ca 2.65 Ga (Section 4). TheSupergroup basins with the Hamersley basin suc-

cession in Pilbara have given rise to a suggested possibly 2.68–2.58 Ga (Treloar and Blenkinsop,1995) Limpopo collisional event between the‘Vaalbara’ supercontinent [Button (1976);

Stanistreet (1993); Cheney (1996); see, however, Zimbabwe and Kaapvaal cratons may also haveoccurred during the early phase of this possibleNelson et al. (1999)]. The Yilgarn and Superior

(North America) cratons underwent a strikingly fragmentation of the postulated Late Archaeansupercontinent. The alternative ca 2.0 Ga Limpoposimilar magmatic and sedimentary evolution from

ca 2.7–2.4 Ga, suggesting that they belonged to a age postulated by Holzer et al. (1998) implies thatthis collision was part of the Eburnean reassemblycommon Late Archaean craton; however, recent

precise age data make such a scenario unlikely event discussed above.There is thus evidence for ca 2.7–2.0 Ga rifting(Nelson, 1998). Comparisons such as those above

need not necessarily reflect adjoining positions for and extension in the >2.0 Ga terranes that makeup the present-day continents of Africa, India andthe ‘southern’ cratons, but similarities could also

merely be ascribed to similar conditions of Australia. This, along with well established colli-sional events (Yilgarn at 2.65 Ga; Limpopo possi-supracratonic sedimentation, related to global pro-

cesses (e.g. Nelson et al., 1999). For the 2.0 Ga bly at 2.68–2.58 Ga; Bastar in Late Archaean;Aravalli–Bundelkhand in Late Archaean–EarlyEburnean–Amazonian collisional event, the evi-

dence is clearer that at least much of Africa and Proterozoic) would be compatible with the ideasexpressed by Nelson (1998) rather than with aparts of South America were assembled into a

supercontinent (Sections 1 and 2.6). In India, more simplistic attenuation and dispersal of a‘southern’ supercontinent. Nelson (1998) discussessupracratonic sedimentation ceased in Singhbhum

by about the same time (±2.0 Ga) suggesting the global scale catastrophic magmatic events whichappear to have had a significant influence on thepossibility that Peninsular India, or at least parts

of it, may also have been subject to such an formation of Early Archaean crust, and suggestsfurther that by Late Archaean–Early Proterozoicassembly or collisional event. In Australia, the

Yilgarn and Pilbara cratons amalgamated along time, a transition to plate tectonic processesoccurred. The increase in mid-ocean ridge volumesthe Capricorn orogen at ca 2.0–1.8 Ga (Section 4).

In India, the Late Archaean Singhbhum and due to plate tectonics would have raised global sealevels (e.g. Eriksson, 1999), thereby flooding pas-Dharwar crustal provinces underwent extension

and rifting for much of the 2.7–2.0 Ga period; in sive margin settings, while concomitantly increas-ing subduction rates promoted island arc-conti-Bastar, Late Archaean collision is suggested by

large scale crustal thickening and widespread calc- nent and small continent–continent collisions.Freeboard would have been reduced for floodedalkaline volcanism (Section 3), while in the

Aravalli–Bundelkhand crustal province, an ana- passive margins and enhanced for collisional ter-ranes. Local tectonic conditions and inheritedlogous compressional regime is inferred for this

time period (Sharma, 1998). On the Kaapvaal Archaean structural crustal features may explainthe diachronous nature of sea level changes seencraton the ca 2.7 Ga Ventersdorp flood basalts

support extensional tectonism, as do the 2775– on the Indian, African and Australian cratons anddiscussed below.2629 Ma Fortescue Group volcanics in Pilbara

(Section 4). The emplacement of the 2.58 Ga Great Logically, the extensional tectonic regimes inthe Late Archaean cratons discussed here wereDyke in Zimbabwe (Mukasa et al., 1996) also

preceded by mantle plume activity. Concomitant raised sea level and lowered freeboard. Otherpossible causes of raised sea level-lowered free-enhanced mid-ocean ridge growth in the oceans

and cratonic extension would reasonably have led board include changes in geoid relief and glacio-isostacy. Although the former is impossible toto trangressive marine depositional environments

on the passive margins of the cratonic plates. resolve in rocks of this age, there is no evidenceof glaciation, either in Africa or elsewhere in theMantle plumes are indicated for the ca 2.7 Ga

Ventersdorp lavas of Kaapvaal (Van der rock record, at ca 2.6–2.4 Ga (e.g. Erikssonet al., 1998b).Westhuizen et al., 1991), for the ca 2775–2629 Ma

Fortescue lavas of Pilbara (Nelson et al., 1999), All three crustal provinces in the SIB, India,also exhibit a Late Archaean BIF–lesser carbon-for the 2.58 Ga Great Dyke in Zimbabwe and for

2.42 Ga intrusions in the Yilgarn craton, Australia. ate–fine-clastic assemblage, indicative of loweredfreeboard and raised sea levels (Section 3) (Fig. 5).In India, Early Proterozoic volcanics in

Singhbhum and Aravalli have a MORB affinity In Dharwar, local tectonism may also have playeda role as the Bababudan and Chitradurga Groups’(Bose et al., 1989; Alvi and Raza, 1992).

Ophiolites, indicating closure of oceans during the BIF are separated by intervening comparativelyhigher freeboard sediments (nearshore carbonates)Eburnean 2.0 Ga collisional event, are inferred for

the North Gabon massif and Ruwenzori belts of (Fig. 5). In Australia, high freeboard in the LateArchaean is inferred from the 2650 Ma orgenicAfrica (Petters, 1991; Fig. 1).

The 2.7–2.0 Ga African rock record reveals two event in Yilgarn, and from the subaerial volcanismand continental sedimentation in the Fortescuepossibly correlatible transgressive epeiric volcano-

sedimentary successions on many of the cratonic Group, Pilbara craton. The Hamersley GroupBIF–carbonate platform on the latter craton,plates (Fig. 3). It is important to note here

that such correlations, as discussed throughout dated at 2597±5–2449±3 Ma (Section 4) is sig-nificantly younger than the Bababudan BIF fromthis paper, may change with the collection of

more precise age data. The first possible wide- the Dharwar craton (thought to be ca 2.7 Ga). Asstated above, local tectonism and the thermal statespread transgression is exemplified by the

Ghaap–Chuniespoort Groups ( lower Transvaal of the various cratons would have been variableand can perhaps explain these apparently diachro-Supergroup, Kaapvaal craton, 2642–2432 Ma), the

Jamba Group (SW Congo craton, estimated to be nous relationships, within a long Late Archaean–Early Proterozoic time period characterised by2.6–2.5 Ga), the Luiza Supergroup ( Kasai shield,

Congo craton, 2560–2432 Ma) and, possibly, common drowned passive margin conditions onmany ‘southern’ cratons. Postulated transgression–by the poorly constrained Kimezian Supergroup

( W Congo craton, <2.6 Ga and >2143 Ma) regression cycles for the Hamersley Group (Pilbaracraton, Australia) and the Transvaal Supergroupand Konse Group ( Usugaran Supergroup, SE

Tanzania craton, <2573 Ma and >2.1 Ga) ( Kaapvaal craton, South Africa) (Fig. 7) providefurther examples of such diachroneity within(Fig. 3). All these successions include either car-

bonate rocks or BIF, or both (in addition to broadly correlatible global episodes of generallyraised sea level conditions.siliciclastic and volcanic lithologies). Carbonate–

BIF platforms would be a logical consequence of Within the pre-2.0 Ga terranes that make uppresent-day Africa and Australia, these Latedrowned passive margins of stable cratonic plates

and raised sea levels/lowered continental freeboard Archaean low freeboard–high sea level conditionsappear to have persisted only into the earliest partconditions; such conditions are compatible with

the idea of enhanced mid-ocean ridge activity or of the Proterozoic (Figs. 3, 7 and 8), and for theIndian terranes, only until ca 2.6–2.5 Ga (Fig. 5).the less plausible concept of Late Archaean super-

continent fragmentation. The widespread occur- There is evidence for a global glaciation at ca 2.4–2.2 Ga (e.g. Young, 1991, 1995). Within therence of this ca 2.6–2.4 Ga African epeiric

succession would argue against an alternative Kaapvaal craton, the Timeball Hill Formation(Pretoria Group, Transvaal basin), and its corre-interpretation of local tectonism as a cause of the

late in the Griqualand West basin, the Makganyene the central Pilbara at ca 2.4–2.2 Ga. The MeteoriteBore diamictites near the base of the Turee CreekDiamictite Formation, contain substantive evi-

dence for glaciation (e.g. Visser, 1971) and are Group (Section 4) support subaqueous glacigenicdeposition, possibly equivalent to the Makganyenedemonstrably within the 2.4–2.2 Ga age limits

inferred for the first global glaciation. Despite the diamictites of Kaapvaal. In contrast, on theYilgarn craton, 2.42 Ga intrusions may be relatedexpected lower sea levels and raised freeboard

which logically accompanied this event, most of to failed N–S extension (D.R. Nelson, 1998, per-sonal communication) and are thus unrelated tothe rocks of the Timeball Hill Formation reflect

epeiric sea deposition (Eriksson and Reczko, 1998) freeboard variation.Aggressive early Precambrian weathering of aand thus marine transgression. This would there-

fore appear to be another example of the influence possible emergent (high freeboard; mantle plumeuplift) Late Archaean ‘northern’ supercontinentof local tectonics, with subsidence of the Pretoria

Group basin floor during deposition of the (Aspler and Chiarenzelli, 1998) combined with thedevelopment of extensive passive margin carbonateTimeball Hill sediments.

Within the Transvaal basin, the combination of platforms in the ‘southern’ cratons, such as dis-cussed in Section 2.2 for example, would have ledperiglacial rocks in the upper Timeball Hill

Formation, 2223±13 Ma Hekpoort subaerial to high rates of CO2 draw down approximately atthe Archaean–Proterozoic boundary. This proba-continental flood basalts (extending also to

Griqualand West and there partly pillowed), and bly helped promote this first global glaciation eventca 200 Ma later, between 2.4 and 2.2 Ga (Erikssonimmature Pretoria Group alluvial formations both

below and above these lavas (Section 2.2), together et al., 1998b). Due to evidence for thermal uplift(mantle plumes?) of cratons which form part ofare compatible with mantle plume uplift, local

glaciation and generally high freeboard–lowered present-day Africa, Australia and India during this2.4–2.2 Ga period, a mantle-igneous event (possi-sea level conditions. Glacio-eustatic sea level fall

and allied high freeboard conditions are also indi- bly global in scale?) may also have played a rolein this glacigenic event. The cold conditions wouldcated for the Early Proterozoic Dhanjori

Formation, Singhbhum province in India (Fig. 5). generally have led to lowered sea levels, highercontinental freeboard and to large scale massFor the Dhanjori, the glacigenic–continental sedi-

ments-plume volcanism association (Table 5) wasting of the continental regions, with reducedpotential to preserve supracrustal successions. Inresembles that seen in the Pretoria Group. It is

conspicuous that the Gawler craton of Australia the Gawler and Yilgarn cratons of Australia, thelarge gaps in the supracrustal record, respectively,experienced a significant compressional, orogenic

episode at ca 2.45 Ga, and thus presumably had from 2.5–2.0 Ga and >2.2 Ga (Section 4), andthat of the Aravalli–Bundelkhand crustal provincehigh freeboard then as well, which condition prob-

ably persisted until ca 2.3 Ga as a result of the of India (ca 2.5–2.1 Ga; Section 3) may partlyreflect such conditions. Examining Fig. 3, andgranitic intrusions following this orogeny.

Analogous high freeboard conditions appear to bearing in mind the lack of accurate chronologicaldata for Africa, there is some further support forhave prevailed on the Zimbabwe craton from ca

2.6–2.2 Ga (Section 2.3) and freeboard may have such a possible gap.Global glaciation would logically have beenbeen further enhanced due to glacio-eustatic

lowering of sea level during the ca 2.4–2.2 Ga followed by glacio-eustatic sea level rise, but sub-ject to local glacio-isostatic development ofglobal glacigenic event (Young, 1995; Aspler and

Chiarenzelli, 1998). The Aravalli–Bundelkhand enhanced freeboard conditions (Eriksson, 1999).In Africa, a poorly dated ca >2.15 Ga transgres-crustal province in India also appears to have

experienced high freeboard conditions during sive epeiric succession comprising siliciclastic andvolcanic rocks with prominent black shales, ismuch of the Early Proterozoic (Section 3). The

siliciclastic sediments of the Turee Creek Group, relatively widepsread (Fig. 3). Elevated sea levelsare also indicated for the post-glacial formationsPilbara craton, point to enhanced freeboard for

in Singhbhum, India; the punctuated transgression whereas sea level during Wyloo Group deposition(Pilbara craton) appears to have been relativelyin Singhbhum (Dhalbhum Formation; Fig. 5) is

most likely a local event, reflecting either tectonic low. A collision between the Pilbara and Yilgarncratons may thus be indicated. In the Pine Creekor glacio-isostatic effects. Alternatively, thermal

subsidence after the Dhalbhum volcanism, com- Inlier (Fig. 6) of Australia, deep water BIF andturbidite sediments predominate from 2.1–1.9 Gabined with rifting, may explain the Kolhan Group

transgression (Fig. 5). Post 2.2 Ga drowning (Section 4). The possible assembly event of the‘southern’ cratons thus appears to have includedevents are unlikely for the Wyloo Group (Pilbara

craton) in Australia (Section 4), and a localised most of the African cratons, and some from India,South America (Section 2.6) and possibly alsoglacio-isostatic influence possibly played a role.

Once again, this second drowning event in the 2.7– from Australia. Whether a single ‘southern’ super-continent formed at ca 2.0–1.8 Ga, or whether2.0 Ga record is characterised by diachronous pat-

terns from craton to craton, within a broadly various larger cratonic terranes amalgamated thenis uncertain.developed period of elevated ocean levels.

The raised sea levels and epeiric deposition ofblack shales and other fine-grained shelf-type sedi-ments during this second drowning event were 6. Conclusionsmost likely terminated by different mechanisms onthe different cratons of Africa. Thermal uplift One thesis discussed here is that of a postulated

Late Archaean supercontinent, comprising cra-related to intrusion of the Bushveld Complex at2054 Ma (Armstrong and McCourt, 1997) termi- tonic blocks from present-day Australia, India,

Africa, South America and Antarctica.nated deposition within the Transvaal basin onthe Kaapvaal craton. Similarly, emplacement of Subsequently, and probably related to mantle

plume upwelling, attenuation and fragmentationthe Cunene Complex in the SW Congo craton atca 2151–2098 Ma (Petters, 1991) ended epeiric possibly ensued with extensional tectonic-related

Late Archaean flood basalts being preserved onmarine deposition there. The Ubendian–Usugaranorogeny at ca 2.1 Ga (Lenoir et al., 1994), during certain cratons (e.g. Kaapvaal, Pilbara, SIB prov-

inces of India). Alternatively, and perhaps morewhich the Zimbabwe craton probably collided withthe Ubendian terrane and Tanzanian craton lying plausibly, a Late Archaean transition from global-

scale, mantle-dominated crustal growth processesto its north (Fig. 1), terminated Piriwiri passivemargin black shale sedimentation; in addition, this to an increasingly predominant plate tectonic

regime (Nelson, 1998) promoted increased mid-collision led to rifting and deposition of the strike–slip related Deweras Group, followed by thermal ocean ridge activity within the oceans surrounding

the various ‘southern’ cratons. This enhanced mid-subsidence and Lomagundi Group deposition. TheEburnean orogeny and assembly at ca 2.05– ocean spreading led to raised sea levels and lowered

freeboard, and during this Late Archaean–Early1.85 Ga of the African continent affected theWestern Congo craton later, at ca 2.0 Ga (Petters, Proterozoic drowning event, epeiric marine sedi-

ments were laid down on the passive margins of1991) and Francevillian Supergroup depositionthus continued until ca >2050 Ma (Fig. 3). In many of the preserved cratonic terranes. BIF and

carbonate as well as transgressive fine to- coarse-India, the Early Proterozoic succession in theSinghbhum province of the SIB was terminated grained clastic sediments characterised this first

inferred global transgressive event. Rapid mid-by a granitic intrusion event (Fig. 5). These sug-gested conditions are compatible with superconti- ocean ridge growth led to extensional tectonic

regimes in many of the ‘southern’ cratons, but alsonental assembly of a ‘southern’ land mass at ca2.0 Ga. In Australia, the shallow marine sediments led to collisions, between some of the cratonic

blocks or with developing island arc complexes.of the Glengarry Group (Yilgarn craton) werepossibly related to rifting along a passive margin The 2.4–2.2 Ga global glaciation was centred

on a ‘northern’ supercontinent, Kenorland, as yet(D.R. Nelson, 1998, personal communication),

still largely assembled and subject to uplift by Jansen, 1979; McLennan and Taylor, 1982; Nelsonand DePaolo, 1985). Nelson et al. (1999) describestrong mantle plume activity (Aspler andan episodic growth of granite–greenstone crust inChiarezelli, 1998). Glacigenic deposits from theseven major events of <10–70 Ma duration,‘southern’ cratons are best preserved in Kaapvaal,between 3470 and 2760 Ma, compatible with majorSinghbhum and Pilbara, and in each case appearcontinental crust formation during the Middleto be related to mantle plumes and extensiveArchaean (Arndt, this volume). The period ofvolcanism. It is thus possible that a global plumeapproximately 3.0 to 2.76 Ga was probably charac-event was partly responsible for this first refrigera-terised by erosion of many cratons (e.g. Pilbara,tion of the earth, in conjunction with enhancedKaapvaal ) and thus continental freeboard wasCO2 drawdown from the greenhouse atmosphereprobably high. The onset of plate tectonics in thedue to carbonate platform development (in theLate Archaean would have promoted collisions offirst drowning event) and aggressive weathering ofemerging island arc complexes with existing olderemergent continental land masses due to atmo-cratons, thereby enhancing continental crustalspheric composition (e.g. Eriksson et al., 1998b).growth rates at the Archaean–Proterozoic bound-Consequent to this proposed global glaciation, seaary. Such rapid Late Archaean continental crustallevels fell eustatically and due to the high freeboardgrowth would also have had the effect of raisingassociated with plume upliftment.global average sea levels and aggressive weatheringA second epeiric marine transgressive eventmay have lowered continental freeboard (Eriksson,affected the ‘southern’ cratons after glaciation1999) at the Archaean–Proterozoic boundary.ameliorated and sea levels consequently rose, along

with reduction in freeboard due to both eustaticchange and warmer climate weathering. This

Acknowledgementssecond, ca 2.2–2.0 Ga drowning event in Africa,India and Australia was characterised by clastic

PGE acknowledges generous research fundingshallow marine sediments in Africa, includingfrom the University of Pretoria and theprominent black shales, with carbonates beingFoundation for Research Development, Southmore common in India. With assembly of theAfrica. Many discussions with numerous col-‘southern’ cratons into one or more superconti-leagues over the years were very beneficial fornents during the ca 2.15–1.85 Ga Eburnean–PGE. PKB is grateful to the Department ofAmazonian orogeny, the 2.7–2.0 Ga volcano-sedi-Science and Technology, India (Grant No.mentary record was terminated, both tectonicallyESS/CA/A9-13/92) and RM to the Universityand by large scale magmatism.Grants Commission, Government of India, forChanges in sea level and continental freeboardproviding financial assistance. The Indian teamfor the cratonic terranes of present-day India,expresses its gratitude to their department forAfrica and Australia thus appear to reflect theproviding infrastructural facilities. It is also

primary influences of enhanced mid-ocean ridgeindebted to Mr K.K. Roy, ex-Director of the

activity, widespread (global?) mantle plume events Geological Survey of India, for inspiration andand global-scale glacio-eustacy, with lesser cases stimulating discussions at the initial stage of theof localised relative sea level changes due to either work. All the authors owe a debt of gratitude totectonics, glacio-isostacy and small plumes. A Late their friends and colleagues in the GlobalArchaean ‘southern’ supercontinent, or a Late Precambrian Sedimentation Syndicate.Archaean transition from catastrophic mantleoverturn to the onset of a plate tectonic regimeaffecting existing cratonic blocks, would support Referencesthe concept of early Archaean continental crustalgrowth, proposed by Armstrong (1981), and the Ahmad, T., Rajamani, V., 1991. Geochemistry and Petrogenesismodel of enhanced crustal growth near the of basal Aravalli volcanics near Nathdwara, Rajasthan,

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ate platform transition, Kaapvaal Craton, as deduced from Beukes, N.J., 1983. Palaeoenvironmental setting of iron forma-tions in the depositional basin of the Transvaal Supergroup,a deep borehole core at Kathu, South Africa. J. Afr. Earth

Sci. 24, 391–410. South Africa. In: Trendall, A.F., Morris, A.C. (Eds.), IronFormation, Facts and Problems. Elsevier, Amsterdam,Altermann, W., Wotherspoon, J.McD., 1995. The carbonates

of the Transvaal and Griqualand West Sequences of the pp. 131–209.Beukes, N.J., 1987. Facies relations, depositional environmentsKaapvaal craton, with special reference to the Lime Acres

limestone deposit. Mineral. Deposita 30, 124–134. and diagenesis in a major Early Proterozoic stromatoliticcarbonate platform to basinal sequence, Campbellrand Sub-Alvi, S.H., Raza, M., 1992. Geochemical evidence for the vol-

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