Sea-level changes in a back-arc-foreland transition: the late Carboniferous-Permian Karoo Basin of South Africa

Sedimentao' Geolo~', 83 ( 1993 ~ 115-131

1t 5

Elsevier Science Publishers B.V.. Amsterdam

Sea-level changes in a back-arc-foreland transition: the late Carboniferous-Permian Karoo Basin of South Africa J o h a n N.J. Visser Department of Geologn., Unit'e~io" qf the Orange Free State, P.O. Bo~ 339, Btocmfontein 9300, South ,4fiiea Received March 19, 1992: revised version accepted September 111, 1992

ABSTRACT Visser, J.N.J.. 1993. Sea-level changes in a back-arc-foreland transition: the late Carboniferous-Permian Karoo Basin of South Africa. Sediment. Geol.. 83: 115-131. The late Carboniferous to early Permian succession in the Karoo Basin consists in ascending stratigraphic order of diamictites, mudrocks, greywackes and immature sandstones. The diamictites are interpreted as glacigene deposits laid down by a marine ice sheet on a shelf. Post-glacial suspension-settling of predominantly mud occurred on the basin floor during open marine and later brackish conditions. Basin floor fan deposits accumulated in the foredeep, whereas deltaic mud, silt and sand finally infilled the basin. Glacigene sedimentation occurred in a back-arc basin during the late Carboniferous and early Permian. The post-glacial succession accumulated in a foreland basin in which the rising orogenic belt was the predominant sediment source. Relative sea-levels, following the major Namurian regression, apparently remained low during the extensive Gondwana glaciation. Marine ice sheets are highly sensitive to sea-level fluctuations: and minor sea-level rises, probably climatically controlled, were recorded as interglacials when parts of the marine ice sheet collapsed. Final collapse of the marine ice sheet was associated with a major transgression. Post-glacial mudrock deposition was charz~cterised by relatively high sea-levels, except for the period of basin floor fan deposition when there was a sea-level Iowstand. Final infitling of the basin by prograding deltas was accompanied by a major regression. The late Carboniferous to early Permian sea-level fluctuations in the Karoo Basin show correspondence with global eustatic cycles reflecting low tectonic interference with sea-level patterns. However, late Permian sea-levels were predominantly tectonically controlled as a result of foreland basin evolution. Possible global eustatic events complemented apparent sea-level changes within the basin during this period.

Introduction The Karoo Basin in southern Africa represents a structural remnant of a widespread Gondwana succession that was deposited during the late Palaeozoic and early Mesozoic. Within its confines sedimentary rocks form a stratigraphic succession about 10 km thick (Visser, 1991a). The late Carboniferous-Permian part of the succession consists of glacigene, deep and shallow marine, brackish sea, basin floor fan and deltaic facies. During deposition the basin configuration changed from a large open basin in the late Carboniferous to a small enclosed one at the Permian-Triassic boundary. Study of the sedimentary facies and, where possible, sequence

stratigraphy give an opportunity to record relative sea-level change in a tectonically controlled basin during a specific time slot. However, to assess the influence of basin tectonics on relative sea-level, an analysis of the basin evolution during the late Carboniferous and Permian has to be included in the description. Very little is known about sea-level change during late Palaeozoic sedimentation in the Karoo Basin. Van Vuuren 11983) briefly referred to sea-level control on coal deposition in the Vryheld Formation. Cooper (1990), in his discussion of the tectonic cycles in southern Africa, suggested a first-order regression in the Cape-Karoo Basin. Recent research has resulted in better constraints on the ages of the stratigraphic units

0037-0738/93/$06.00 © 1993 - Elsevier Science Publishers B.V. All rights reserved

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J.N,J. V I S S E R

ogy (Anderson, 1977), occurs within the lower half of the succession (Fig. 3). This time-line cuts across the succession with the result that the diamictites become younger in an eastern (upglacier) direction as well as towards the northern basin margin where the diamictites and associated mudrocks are confined to valley fill~ ~.Visser, 1989). The palaeotopography, glaciated surfaces, glacial facies and clast composition suggest transport of material from the palaeo-east, palaeosouth and palaeo-west during the late Carboniferous and early Permian (Visser, 1989). The overlying mudrocks consist of the Prince Albert Formation at the base, followed by the whiteweathering Whitehill Formation, that is overlain by the Collingham and Vischkuil Formations at the top (Fig. 2). A white-weathering, thin, carbonaceous shale which contains marine fossils, is present at the base of the Prince Albert Formation along the southern basin margin (Fig. 4). Very sparsely distributed dropstones are present within the basal part of this unit. The main outcrops of the

(cf. Visser, 1990), but lack of fossils, nevertheless, prohibits the definition of higher-order transgressions and regressions. Late Carboniferous-Permian succession

The succession in the Karoo Basin, which covers an area of over 600,000 km 2 (Fig. 1), consists of diamictites at the base, overlain by a predominantly mudrock succession, and sandstones at the top attaining a cumulative thickness of just over 2000 m (Fig. 2). The present basin axis is about east-west, but during the late Palaeozoic this axis was oriented in a north-south direction (Visser, 1991a). The glacigene Dwyka Group consists of massive and stratified diamictites, boulder beds and mudrocks some of which contain ice-rafted debris. Two persistent mudrock horizons are interbedded in the d:amictites (Fig. 3). The upper mudrock unit is correlated with the Eurydesma transgress,~on which is of early Sakmarian age (Dickins, 1984; Visser, 1990). The inferred Permian-Carboniferous boundary, based on palynol-

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:o',i I Fig. 3. Stratigraphic sections of the Dwyka Group in the southern Karoo showing the inferred position of the Permian-Carboniferous boundary (after Anderson, 1977) and the interglacial mudrocks. See Fig. 1. for the location of the section.

Prince Albert Formation consist of micaceous, silty and dark-grey to black carbonaceous shale. The formation shows a facies change towards the northeast and northwest where mere siltstones and sandstones are present (Fig. 4). Concretions, lenses and irregular bodies of carbonate, chert and phosphatic rock are present within the mudrocks. Viljoen (1990) recognised widely distributed, interbedded, volcanigenic material and tuffaceous layers. The upper contact with the Whitehill Formation is sharp, except in the north where a thin upward-coarsening sequence from carbonaceous mudrock to silty shale occurs at the top of the Prince Albert Formation (Fig. 4). At Douglas marine fossils, as well as fossil wood and palaeoniscoid fish remains and coprolites, were recorded (McLachlan and Anderson, 1973). The Prince Albert Formation has an Artinskian to Kungurian age (Visser, 1990).

The Whitehill Formation is developed only in the western and central parts of the basin, because towards the east and northeast it loses its character (cf. Cole and McLachlan, 1991) so that the lower mudrocks become indistinguishable from the upper mudrocks. Therefore beyond the lateral cut-off (Fig. 1), the Pietermaritzburg Formation overlies the diamictite beds (Fig. 4). Black, very thinly laminated, carbonaceous, pyritebearing shale constitutes either the entire or about 75% of the Whitehill Formation. Thin, yellowishweathering, rhyodacitic tuffaceous beds (Rowsell and De Swardt, 1976; McLachlan and Anderson, 1977a; Viljoen, 1990), carbonates, chert and siltstones constitute the remainder of the formation. Interbedded silty shale is more common in the northwest reflecting proximity to the basin margin (Fig. 5). The upper contact is commonly trans;tional in the north where interbedded silty beds

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SEA-LEVEL CHANGES IN A BACK-ARC-FORELAND TRANSITION

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are present. Ferruginous, carbonate concretions with a dolomitic composition are dispersed throughout the formation. The orientation of fossils (Oelofsen, 1987) and facies changes (Fig. 5) suggest major sediment transport from the highlands m the palaeo-east, whereas a subordinate volcaniclastic source was towards the palaeo-west (Viljoem 1990). The biostratigraphy of the Whitehill Formation suggests synchronous deposition across the Karoo Basin as the fossil range zones are confined to the upper part of the succession (Oelofsen, 1987) (Fig. 5). The fossils include a swimming reptile (Mesosaurus tenuidens) (Oelofsen and Araujo, 1987), insect wings (McLachlan and Anderson, 1977b), plant remains (Glossopteris leaves and fossil wood), sponge spicules, palaeoniscoid fish (Palaeoniscus capensis) and arthropods (Notocaris tapscotti). Visser (1990)

suggested a possible late Kungurian to early Ufimian age for the Whitehill Formation. A thin, black shale with marine fossils is p~esent at the base of the Pietermaritzburg Formation. The presence of phosphorite in the black shale (Biihmann et al., 1989) and the absence of chert may suggest correlation with a higher stratigraphic level in the Prince Albert Formation (Fig. 4). The overlying mudrocks also show the same characteristics as the upper part of the Prince Albert Formation. Along the northeastern basin margin three upward-coarsening coal-bearing cycles belonging to the Vryheid Formation overlie a basal Pietermaritzburg shale (Stavrakis, 1986). Glauconite is present within the sandstone associated with the upper coal seam. Dark-grey shales of the Volksrust Formation overlie the coal measurcs.

The Collingham and Vischkuit Formations

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Fi~. ft. Strafi#raphic ~cciion~- of lhe upper part (Vbx-hkuit/Coliingham, Skoorsteenberg/Laing.~burg/Ripon. Fort B r o w n / T i e r b e r g and W a t e r f o r d / K o e d o e s b e r g Forma[flms) of the Eeca Group, The Koonap Formation of lhe Beaufort Group largely represents

delta plain sedimentation.

122

overlying the Whitehi[l Formation consist of rhythmically bedded shale, siltstone, very finegrained sandstone, chert and thin tuffaceous layers. Dolomiti,; concretionary layers and lenses and fossil wood are present. The Tanqua, Laingsburg and Ripon sandstones within the thick mudrock succession (Figs. 2 and 6), form a discontinuous outcrop, althoagh they occur on the same stratigraphic level. The Ripon Formation consists of two well-defined cycles (Kingsley, 1981) of which the lowermost may be stratigraphically older than thc main sand depositional event. The formalions consist of alternating fine-grained greywacke and shale forming upward-coarsening sequences. In the greywackes, Bouma sequences are commonly incomplete (Kingsley, 1981). There i~; an abundance of trace fossils, slumping and sole marks. The palaeoslope for the deposition of the turbidite sands was towards the palaeo-east and palaeosouth (Kingsley, 1981). The overlying Fort Brown and Tierberg Formations consist of shale, silty shale and subordinate fine-grained sandstone. The succession becomes arenaceous towards the top where it forms a transitional contact with the Waterford and Koedoesberg Formations (Fig. 2). These formalions consist of alternating thickly bedded sandstone, siltstone, mudstone and shale arranged in tip to four upward-coarsening sequences. Plant and reptilian remains (Rubidge and Oelofsen, 1981), molluscs (Cooper and Kensley, 1984) and trace fossils are present. The bivalves and tctrapods have a mid to late Kazanian age (Cooper and Kensley, 1984). The sand infilling the basin was derived predominantly from tile palaeo-north and palaeo-west (Visscr, 1991a). Basin tectonics, lithofacies interpretation and sequence stratigraphy

Glacial sedimentation Tile large ba,~;in (designated the Sowegon basin after SOuthWEstern GONdwana) in which tile glacial deposits accumulated in southwestern Gondwana, probably formed by subduction of a palaeo-Pacific plate underneath the Gondwana

J.N.J. V I S S EI~.

plate during tile mid-Carboniferous (cf. Smellie, 1981; Mahlburg Kay and Mpodozis, 1991). Subduction led to the formation initially of a magrustic arc (proto-Precordillera) and a large backarc basin which, together with the adjoining highlands and mountain ranges, had been largely icecovered since the mid-Carboniferous (Fig. 7A). Analysis of the glacial facies suggests the presence within the basin of a passive margin on the palaeo-eastern craton side and an active margin on the arc side. H~ilbich et al. (1983) reported the first evidence of a compressional paroxysm at about 278 Ma which suggests the onset of a change from extensional to compressional conditions within the back.arc basin. On the shelf (passive margin) a thick glacial sequence was deposited initially by grounded ice and later by rain-out of ice-rafted debris~ suspension settling of mud and sediment gravity flow (Visser, 1989). The glacial sedimentation in the Karoo Basin represents the deglaciation phase of tile extensive Gondwana ice sheet when higher relative sea-levels initiated unstable glacial conditions. It can be assumed that sea-levels were very low during maximum glaciation (mid-Carboniferous) (Fig. 2). During the late Carboniferous large sections of the continental ice sheet became decoupled from its substrate and were transformed into an unstable marine ice sheet with increased glacial sedimentation rates. Such a transformation in the ice sheet is ascribed to a rise in relative sea-level which cannot be entirely attributed to isostatic depression of the basin floor. Partial or total collapse of such an unstable ice sheet supplied large vohimes of meltwater which would have resulted m a cumulative sea-level rise on a global scale. The glacial Dvo'ka Group thus represents a record of low sea-level in the beginning and relatively higher sea-levels for the remainder of the glaciation, except for thc two interglacial periods when relative sea-level was high (Fig. 2). 2. Syn- to post-ghwial mud deposition The transition from a back-arc to a foreland basin during tlle late early Permian resulted in a major change in basin configuration (Fig. 7B).

SEA-LEVEL CHANGES IN A BACK-ARC-FORI'2kANDTRANSITION

123

The post-glacial open sea inundated large parts of the continental interior, subsidence of which can be attributed to crustal flexure. The inundated region formed a gently sloping foreland ramp which merged into a foredeep in the palaeo-west. The palaeo-western margin of the foreland basin was defined by a rising inner arc (cf. Kingston et al., 1983) of which the formation can possibly be attributed to the change from an extensional to compressional regime (cf. Wilson, 1991). H~iibich et al. (1983) reported a major compressional paroxysm in the Cape Fold Belt at about 258 Ma which suggests that ongoing compression along the inner arc eventually produced a rctro-arc fold-thrust belt. Sediment ,_'nput into the incipient foreland basin was predominantly from the highlands in the palaeo-east and to a lesser extent fl-om the palaeo-west as suggested by the facies changes in the Prince Albert and Whitehill Formations. Acidic volcanism along the distant magmatic arc supplied windblown volcaniclastics to the foreland basin (cf. Viljoen, 1990). At the time of the final collapse of the marine section of the ice sheet the passive margin in the basin was depressed beyond a critical limit and no isostatic adjustment occurred. The ice front thus retreated in relatively deep (> 250 m) water (Visser, 1991b). The thin, black, fossi!iferous mudrock horizon overlying the diamictite with an abrupt contact is interpreted as a marine condensed section (MCS) (cf. Mitchum, 1977; Haq, 1991; Figs. 2 and 4), which was deposited during a sea-level highstand. The disseminated pyrite and remains of a shark within this section suggest very low sedimentation rates and scdiment starvatkm (ct. Haq, 1991; Garzanti, 1991). This marine transgression is attributed to sudden collapse of the marine ic'," sheet (cf. Carter et al., 1986;

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+ + Highland (~)(~ Orogenic fold-thrust belt Ice flow/palaeocurrent

Fig, 7. A. Configuration of the back-arc Sowegon basin during the late Carboniferous ( + 300 Ma). The basin had been icecovered since the mid-Carboniferous. B. Basin configuration during the latc early Permian sea-level highstand ( +_26(I Ma). The development of an inner arc and a foredeep substantially changed the shape of the basin. C. Configuration of the foreland basin during the late Permian regression ( +_255 Ma). See Fig. 2 for the time slots of A, B and C.

124

J.N.,I. VISSI£R

Lopez Gamundi, 1989: Anderson and Thomas, 199t) causing sea-level to rise. Although the ice sheet collapsed over the shelf,

mountain ice caps were maintained and glacial outwash was fed in~,~ ti~e basin. The fine mud togeti~.er wi'~h phosphatic material, biogenic silica

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Sequence boundary

. . . . . Par~sequenee boundary

Fig. N. Generalized sccti~ms of |b.c Prince Albcr~ and \¥hi[c'hitt F,~rmatio~s ~md of the c()al-bcaring VryhL.id Form'~tiun (after Stavrakis. 198h). ~d~owmgscaJc~cI fluctu~tion~ during the ear};' and cartydatc Permian. Ages of the forrn~i~ms a~tcr Visscr (It}90). P~tb = Pictcrmarilzburg.

l:m .-: F ' o r m ~ t i o m ),1(.~, : m u ~ i n c c o n d c n , , c d , < c t i ( m , L , = [ r a n : , e , ~ e s s i o n

S E A - L E V E L C H A N G E S IN A B A C K - A R C - F O R E L A N D T R A N S I T I O N

and volcanic ash accumulated in the deeper part of the basin (Fig. 7B). High sea-levels enhance the formation of phosphorite and biogenic chert (cf. Cook, 1984) so that the lower part of the Prince Albert Formation represents a highstand wedge (Figs. 2 and 8). In the Kroonstad area a thin Pietermaritzburg shale, which formed during the Pietermaritzburg transgression (Fig. 8), overlies the glacial outwash (Stavrakis, 1986). The upper part of the Prince Albert Formation is correlated with the lower deltaic cycle of the coal measures along the northeastern basin margin (Fig. 8). In the coal measures a major regression wh'ch followed on the Pietermaritzburg transgression, coincides in the Prince Albert Formatioa with a silty horizon or a sharp break between dark-grey shale with chert and micaceous shale and mudstone representing distal turbidite deposits (Fig. 4). This break represents a sequence boundary (cf. Mitchum and Van Wagoner, 1991), whereas the contact between the Prince Albert and Whitehill Formations is a parasequence boundary (Fig. 8). The coal-forming deltas formed during a regional highstand (cf. Haq, 1991), whereas the suspended mud was deposited together with terrestrial organic material in a distal prodclta setting. Thus during deposition of the upper part of the Prince Albert

| 25

Formation relative sea-level remained high (Fig. 2). The overlying Whitehill Formation consists of highly carbonaceous, laminated shale which distinctly differs from the underlying mudrocks (Fig. 9). The shales formed by suspension settling of fine mud in a highly anoxic environment. Towards the shallower basin margins more oxygenated bottom conditions prevailed and the deposition of coarser grained sediments took place (Figs. 1 and 5). The Whitehill Formation is correlated with the middle and upper coal-bearing deltaic cycles of the northeastern basin margin (cf. Cole and McLachlan, 1991)(Fig. 8). Although marine fossils are absent in the lower part of the Whitehill Formation, the R b / K ratio of the shales suggests marine conditions (Visser and Young, 1990). The lack of medium-grained sediment input into the basin and the strongly anoxic conditions can be attributed to a sea-level highstand. This is substantiated by the presence of coal-forming deltas along the basin margin. The glauconite in the siltstones and sandstones overlying the upper coal seam suggests a flooding surface (cf. Mitchum and Van Wagoner, 1991). The or~n.ic material in the Whitehill Formation consists partly of finely comminuted plant remains which were produced by shoreface erosion as the

Fig. 9. The white-weathering Whitehill Formation (arrow) which possibly represents a marine condensed section. A dolerite sill caps the mountain.

~.2~1

~ransgressive front moved over the coal-forming coastal deposits (cf. Garzanti, 1991). During temporary shallower conditions carbonatcs and silty shales were deposited towards ~,he basin interior. The reptilian, arthropod and fish remains in the upper Whitehill shales suggest brackish conditions (cf. Oelofsen, 1987; Visser and Young, t990). This abrupt change from marine to brackish conditions is attributed to a change in basin configuration whereby marine conditions became more restricted due to impeded oceanic circulation (cf. Fig. 7B). The lower part o f the Whitehi?i Formation is thus interpreted as a possible condensed section (Fig. 8), whereas the .upper part may also represent deposithm during a sea4evel highstand. Superimposed on the highstand was a regression terming a parasequence boundary at about the middle of the Whitehil[ Formation (Fig. 8).

Basin floor mud and fan deposition Late Permian basinward migration of the rising fold-thrust belt caused loading of the cratonic lithosphere resulting in further subsidence of the foreland ramp to form a wide flexural basin (Fig. 7C). A nmv source area developed in the palacenorth, whereas the rising fold-thrust belt i~1 the palaeo-west formed the major sediment source. The dark-coloured Ecca mudrocks, overlying the Whitehill Formation and the coabhearJng strata along the basin margin, lknmed by suspensio~ settling of mud as well as distal mt~d turbiditcs (Lock, 1973) in a possible brackish environment. The presence of fossil wood in the shales suggests rapid drowning of the coastal lowlands accompanied by transportation of the dead wood basinwards. The Vofksrust transgression terminated Whitehill deposition and sealevels remained thus relat~ely high during the mud deposition (Figs, 2 and 8). Basin tloor fans intcrbedded in lhc mudrock succession were deposited in the foredeep of the R)reland basin (Fig. 6). The sandstones which dominated the fans, were deposited by turbidity currents forming an upward-coarsening sequence in the Laingsburg section. The fans are radial in pattern and were probably canyon-fed from local

J.N.J. \ ' | SSE~',I

rivers (cf. Stow et al., 1984). Direct funnclling of sediment through canyons and fan valleys to deeper basins occurs when there are narrow shelves associated with sea-level towstands (Stow et al. 1984). The upward-coarsening sequence is interpreted as a lowstand wedge (cf. Shanmugam and Moiola, 1982) (Fig. 2) of which the deposition was terminated with a transgression when a blanket deposit of suspension mud was laid down on the basin plain. Some of the muds may represent a distal prodelta facies (cf. Kingsley, 1981). However, the possibility that these fans could also have formed as a result of a marked increase in submarine slope caused by shortening in the orogenic wedge (cf. Allen et a!., 198~), cannot be ignored. The separate but simultaneous formation of the fans rather suggests a lowering of sea-level which, possibly together with tectonically induced unstable slopes, were the triggering mechanisms for fan development. The abrupt termination of each fan cycle indicates a sudden cut-off in sediment supply which could have been caused by switching of the input system or a sea-level rise (Wickens, 1990).

Deltaic sedimentation Towards the cud of the deposition of thc Ecca Group an almost hind-locked basin had dcvcloped. Rapid uplift and basinward migration of the fold-thrust belts in the palace-west and to a lesseE e~tcnt in the palace-north formed prominen|-mountain ranges which were the predominant sediment sources for the thick late Permian to carl}, Triassic basin fill. Thinning and possibly thermal weakening of the lithospkcre led 1o extensive crustal flexure immediateiy adjaccnt to the growing orogenic belts (cf. Allcn et al., 198~) to accommodate the scdimem pile m thc shrinking basin. The final phase in the infilling of the basin was by prograding deltas from the palace-west, palace-north and palace-southeast (Visscr, 1991a). Ripple-laminated silty mud and sand with plant and vertebrate remains were deposited as prode[ta and dclta front layers (Figs. 2 and 6). There is stratigraphic and palaeontological evidence that the late Permian deltaic infiiling of

SI/A-I.I:\,'LI. ('ttAN(;I:S IN ?., I~;A('K-AI~.('-FOR}~I,ANDTRANSITION

the Karoo Basin represents a diachronous event (cf. Jordaan, 1981) with the deltas along the active palace-northern and palaeo~western margins of the basin being older than tho~e in the palaeo-east. Uplift of the orogenic belt in the palaeo-north and palaeo-west may have been the first of a series of tectonic pulses in the source areas during the Permo-Tciassic (cf. Visser and Dukas, 1979). Thus the major regression in the basin was controlled by the sheer volume of sediment input exceeding the rate of basin subsidence. This is illustrated by the highly increased sedimentation rate in the basin during the late Permian (Fig. 10).

127

i SEDIMENTATION AGE RATE(m/MB) (Ma) iO 100 200 248 z LU--< ~

W

Upliit with compressional deformation. Subsidence I of foreland ramp

o~

2 Sub.~tant~al change in

Z

basin configuration.

Formation of inner arc. Volcaniclastics

0.

E E

g~ I

"r"

uJ

286 4 ~

~ Decrease in plale .~ ' movement. Thermal subsidence phase. Passive b a c k - a r c basin in e~tensional setting?

,,=

i

i

0

TM

Fluvial --z-Ji~"i "~ ~v .ca

258

Relative sea-level changes vs global eustatic fluctuations

The measure of relative sea-level change in a basin is the interaction of regional tectonism, eustasy and the rate of sediment supply (Hag, 1991). The start of a foreland basin is comm.only characterised by an underfillcd stage due to a delay in the orogenic cycle awaiting the emergence of the fold-thrust belt above sea-level (Allen et al., 1986). Starved sedimentation (sedimentation rate < 20 m / M a ) and the deposition of dark shales in the Karoo Basin during the period 265 to 258 Ma (Fig. 10) suggest relatively orogenic quiescence (cf. Tankard, 1986). The rapid increase in sedimentation rates (> 200 m/Ma), the deposition of turbidite fans and delta progradalton during the late Permian indicate major uplift in the fold-thrust belt. Thus, the lk~reland fold-thrust belt dominatcd sedimentation since the start of the tat, Permian. Cooper (1990) prcscqted a very simplistic view of eustatic events in the Karoo Basin by inferring a first-order regression from the Devonian up to the Triassic. There is evidencc that the regression peaked during the mid-Carboniferous (Namurian) with the onset of glaciation (Veevers and Powel[, 1987; Visser, 199(}) and was followed by a major highstand most probably associated with the collapse of the large Permo-Carboniferous ice sheet over Gondwana. Sea-level highstands tend to show a correlation with periods of global warming (Haq, 1991)and in the Karoo Basin this is aptly

TECTONIC EVOLUTION

315

I

,

t* High rate of plate movement. Subduction of PalaeoPacrlic oceanic crust beneath Gondwana. Partial destruction of shell

Fig. 1(). Tectonic evolulion (~f die Karoo Basin during |he late ('arbimiferot, s and Permian. The subdanlial increase in sedinlcillLtl[Oll Fate illdic:.lles thc starl of lI!e {/lOgelliC phase in tile evolution ~1 Ihc l~rcland basin.

illustrated by the dark-coloured mudrock succession overlying the glacial deposits and onlapping onto the basement. The queslion arises whether the relative sealevel changes recorded in the Karoo Basin during the late Carboniferous and Permian (Fig. l l A ) could be correlated with global sea-level trends. A third-order sea-level curve after Ross and Ross (1987, 1988) has been used as a norm for global comparison (Fig. liB). The Ross and Ross curve shows high sea-levels during the late Carboniferous and part of the early Permian (Sakmarian and Artinskian) and fluctuating sea-levels during the Asselian and Kungurian. Global sea-levels were low during the mid-Carboniferous and early late Carboniferous (Namurian ~and Weslphalian) and then again during the late Permian.

128

J.N..I. VISSEP,

RELATIVE ;EA-LEVEL AGE I LOW ~

~H!GH

=~........ "&"

.~

HIGH

LOW ~ ' -

13

253 "1 258.-- 263 MCS ._,,= t 268

1#,~ J1 C ~

o.-('2 E.=

f

i

?

/

Fig. I1. Relative sea-level fluctuations during the late Carbor, iferous and Permian: A. Karoo Basin; B. inferred global fluctuations (after Ross and Ross, 1985, 1987, 1988). Timescale after ttarland et al. (1982). aa', bb', co' and Jd ~ may represent synchronous events. MC'S= marine condensed section

Despite the lack of good age constraints on the late Palaeozoic Karoo strata certain correlations between their relative sea-level curve and global eustatic events are evident. The sharp rise in sea-level during the Westphalian (aa'), the early Sakmarian Eurydesma transgression (bb')which Dickins (1984) showed to be a well-established global event, a regression at the time of the Artinskian-Kungurian boundary (cc') and a possible late Kungurian transgression (rid') show remarkable correlation (Fig. 11). That some of these events are slightly out of phase must be ascribed to the lack of precise dating of the Karoo strata. A major difference between the Karoo sealevel curve and the Ross and Ross global curve is the inferred lowstand during the P e r m o CarboniSzrous glaciation. A global regression which is mainly attributed to the growth of the Gondwana ice sheet, occurred during the midCarboniferous (Veevers and Powell, 1987). This large ice sheet was at least maintained up to the early Permian, although evidence from the central South American basins suggests a collapse of its southwestern section already at the beginning of the late Carboniferous (Lopez Gamundi, 1989). It is theoretically unlikely that high sea-levels, which would have influenced global climate, occurred during this period (cf. Haq, 1991). Vast amounts of water were locked up in the continental ice sheet over Gondwana as well as in its marine sections located over the shelf regions, all of which would probably enhanced a sea-level lowstand. Evidence from the Karoo Basin suggests a change during the late Carboniferous from grounded ice on the shelves to floating ice which Visser (1989) attrEmted to the formation of a marine ice sheet in the Sowegon basin (Fig. 11A). For the decoupling of ice on a shelf a relative rise in sea-level which could have been the result of isostatic depression of the shelf by the ice or a global eustatic event, was necessary (Visser, 1991b). Also, the formation of a marine ice sheet which is inherently unstable, would have resulted in a thinner overall ice sheet, thereby decreasing the water volume locked up in ice over land. It can thus be argued that relatively higher global sea-levels could have been maintained since the

S E A - L E V E L C H A N G E S IN A B A C K - A R C - F O R E L A N D T R A N S I T I O N

formation of the Permo-Carboniferous marine ice sheets (cf. the existence of the present West Antarctic marine ice sheet despite higher postPleistocene sea-levels). The transformation of the shelf section of the Gondwana ice sheet to a marine ice sheet probably took place during the sea-level rise in the Westphalian (a" in Fig. llB) and subsequent sporadic lowering in sea-level (e.g. during the Asselian) may then reflect the re-occurrence of grounded ice conditions on the shelf. This may imply a much greater prevalence of floating ice during the Permo-Carboniferous glaciation. The major post-glacial transgression may, however, be an oversimplification because effects of regional variation are not taken into account. Marine ice sheets are very sensitive to sea-level fluctuations (cf. Anderson and Thomas, 1991) so that the collapse of the marine ice sheet in the Sowegon basin was most probably triggered by a pre-collapse sea-level rise. Crustal adjustment by glacio-isostacy, hydro-isostacy and sediment loading also could have caused local transgressions and regressions (cf. Carter et al., 1986). Because marine ice sheets are partly afloat, glacio-isostatic adjustment after collapse of the ice sheet is insignificant. However, hydro-isostacy and sediment loading may have have contributed to the regional inundation of the basin. Inundation of the palaeotopography in the Karoo Basin suggests a sea-level rise of more than 200 m during the early Artinskian (Fig. 8), but this figure may be unrealistic because of the cumulative effect of isostasy and basin subsidence on the sea-level rise. Lopez Gamundi (1989) recorded a sim~!ar sequence of events for Carboniferous basins in South America, thereby !ending, further support for a global rise in sea-level related to late Carboniferous deglaciation (cf. higher eustatic levels at the time of the Namurian-Westphalian boundary--Ross and Ross, 1987). Heckel (1986) and Ross and Ross (1987) also emphasised the control of the Gondwana glaciation on the late Carboniferous cyclothems in North America. These cyclothems display rapid marine transgressions when melting of the ice took place and slow interrupted regressions when there was a build-up of ice. Thus, even if the sea-level rise in the Sowegon basin during the early Artinskian was

129

much less than predicted (Stavrakis, 1986, suggested a rise of 100 to 150 m), one would have expected a reflection of this event in tt;e global sea-level curve. The lack thereof in the curve cannot be explained satisfactorily. However, melting of the West Antarctic marine ice sheet which is about of the same dimensions as the Sowegon ice sheet, would upon collapse increase global sea-level by only 6 m (Oerlemans and Van der Veen, 1984). The question, therefore, remains whether the net effect of the melting of the marine ice sheet on global sea-level had been so small that it is not reflected in a third-order curve.

The cyclic eustatic changes towards the end of the early Permian (Fig. 11) can be attributed to the final melting phase of the remaining Gondwana ice caps. On the other hand, cyclic sea-level changes during the late Permian could have also been caused by episodic thrusting and lithospheric relaxation in the forelapd basins in southwestern Gondwana. Possible global eustatic fluctuations thus merely complemented the influence of tectonics on depositional patterns. All interpreted eustatic fluctuations for the late Permian in the Karoo Basin must be cons;dered as highly speculative because sedimentation was controlled by tectonic instability in the source areas and within the basin itself, For the late Carboniferous and largest part of the early Permian a possible global eustatic signal can be discerned in the Karoo Basin above the "noise" generated by isostatic and tectonic effects. Global changes of sea-level for the Carboniferous and from the late Triassic to the present are reasonably welt documented with respect to their ages, durations and relative amplitudes. Data for the Permian are less abundant (cf. Vail et al., t977; Ross and Ross, 1985) and, therefore, interpretation of eustatic changes in the Karoo Basin contributes to our understanding of possible global events during one of the major continental accretion phases in earth's history.

,Acknowledgements Financial support for the study of the tectonic evolution of the Karoo Basin by the Foundation

]30

for Research Development and the University of the Orange Free State is gratefufiy acknowledged. My sincerest thanks to John Veevers and Ken Eriksson for constructive comments on an earlier draft of the manuscript. References Allen, P.A., Homewood, P. and Williams G.D., 1986. Foreland basins: an introduction. In: P.A. Allen and P. Homewood (Editors), Foreland Basins. Spec. Publ. Int. Assoc. Sedimentol., 8: 3-12. Anderson, J.B. and Thomas, M.A., 199l. Marine ice-sheet decoupling as a mechanism for rapid, episodic sea-level change, the record ef such events and their influence on sedimentation. Sediment. Geol., 70: 87-104. Anderson, J.M., 1977. The biostratigraphy of the Permian and the Triassic Part 3. A review of Gondwana Permian palynology with particular reference to the northern Karoo Basin, South Africa. Mem. Bot. Surv. S. Afr., 41, 188 PP. Bi.ihmann, D., BiJhmann, C. and Von Brunn, V., 1989. Glaciogenic banded phosphorites from Permian sedimentary rocks. Econ. Geol., 84: 741-750. Carter, R.M., Carter, L. and Johnson, D.P., I986. Submergent shorelines in the SW Pacific: evidence for an episodic post-glacial transgression. Sedimentology, 33: 629-649. Cole, D.I. and McLachlan, I.R., 1991. Oil potential of the Permian Whitehill Shale Formation in the main Karoo Basin, South Africa. In: H. Ulbrich and A.C. Rocha Campos (Editors), Gondwana Seven, Proc. Inst. Geoci., Univ. Sao Paulo, pp. 379-390. Cook, P.J., 1984. Spatial and temporal controls on the formation of phosphate deposits--a review. In: J.O. Nriagu and P.B. Moore (Editors), Phosphate Minerals. SpringerVerlag, New York, N.Y., pp. 242-274. Cooper, M.R., 1990. Tectonic cycles in southern Africa. Earth Sci. Rev., 28: 321-364. Cooper, M.R. and Kens!ey, B.. 1984 Endemic South American Permian bivalve molluscs from the Ecca of South Africa. J. Paleontol., 58: 136(/-1363. Dickins, J.M., 1984. Late Palaeozoic glaciation. BMR J. Aust. Geol. Geophys., 9: 163-169. Garzant;, E., 1991. Non-carbonate intrabasinal grains in arcnites: their recognition, significance, and relationship to eustatic cycles and tectonic setting. J. Sediment. Petrol., 61 : 959-975. Hfilbich, I.W., Fitch, F.,I. and Miller, J.A., 1983. Dating the Cape orogeny. In: A.P.G. S6hnge and 1.W. Hfi[bich (Editors), Geodynamics of the Cape Fold Belt. Spec. PuN. Geol. Soc. S. Air., 12: 149-164. Haq, B.U., 1991. Sequence stratigraphy, sea-level change, and significance for the deep sea. In: D.IM. Macdonald (Editor), Sedimentation, Tectonics and Eustasy: Sea-level Changes at active Margins. Spec. Publ. int. Assoc. Sedimentoi., 12: 3-39.

J.ri.J. VISSER

Harland, W.B., Cox, A.V., Llewellyn, P.G., Pickton, C.A.G., Smith, A.G. and Waiters, R., 1982. A Geologic Time Scale. Cambridge University Press, 131 pp. Heckel, P.H., 1986. Sea-level curve for Pennsylvanian eustatic marine transgressive-regressive depositional cycles along midcontinent outcrop belt, North America. Geology, 14: 330-334. Jordaan, M.J., 1'48l. The Ecca-Beaufort transition in the western parts of the Karoo Basin. Trans. Geol. Soc. S. Afr., 84: 19-25. Kingsley, C.S., 1981. A composite submarine fan-delta-fluvial model for the Ecca and lower Beaufort Groups of Permian age in the eastern Cape Province, South Africa. Trans. Geol. Soc. S. Aft., 84: 27-411. Kingston, DR., Dishroon, C.P. and Williams, P.A., 1983. Global basin classification system. Am. Assoc. Pet. Geol. Bull., 67: 2175-2193. Lock, B.E., 1973. Distal turbidites from the Middle Ecca. Trans. Geol. Soc. S. Aft., 76: 169-171. Lopez Gamundi, O.R., 1989. Postglacial transgressions in late Paleozoic basins of western Argentina: a record of glacioeustatic sea level rise. Palaeogeogr., Palaeoclimatol., Palaeoecol., 71: 257-270. Mahlburg Kay, S. and Mpodozis, C., 1991. Late Paleozoic to Triassic Gondwana magmatism and tectonism akmg the South American margin (28 ° to 33°S). Abstr. XII ICC-P, Buenos Aires, p. 59. McLachlan, I.R. and Anderson, A.M., 1973. A review of the evidence for marine conditions in southern Africa during Dwyka times. Palaeontol. Aft., 15: 37-64. McLachlan. I.R. and Anderson, A.M., 1977a. Carbonates, "stromatolites'" and tuffs in the lower Permian White Band Formation. S. Afr. J. Sci., 73: 92-94. McLachlan, I.R. and Anderson, A.M., I977b.. Fossil insect wings from the early Permian White-'Band Formation, South Africa. Palaeontol. Aft., 2(1: 83-86. Mitchum, R.M., 1977. Seismic stratigraphy and global changes of sea level, Part I. Glossary of terms used in seismic stratigraphy. In: C.E. Payton (Editor), Seismic Stratigraphy --aaplications to Hydrocarbon Exploration. Mere. Am. Assoc. Pet. Geol., 26: 205-212. Mitchum, R . M and Van Wagoner, J.C., 1991. ttigh-frequency sequences and their stacking patterns: sequence-stratigraphic evidence of high-frequency eustatic cycles. Sediment. Geol., 70: 131-160. Oelofsen, B.W., 1987. "Fhe biostratigraphy and fossils of the Whitehill and lrati Shale Formations of the Karoo and Paranfi Basins. In: G.D. McKenzie (Editor), Gondwana Six: Stratigraphy, Sedimentology, and Paleontology. Geophys. Monogr. Am. Geophys. Union, 41: 13t-138. Oelofsen, B.W. and Araujo, D.C,, 1987. Mesosaunts tenuidens and Stereosternum tumidton from the Permian Gondwana of both southern Africa and South America. S. Afr. J. Sci., 83: 370-371. Oerlemans, J. and Van der Veen, C.J., 1984. Ice Sheets and Climate. Reidel, Dordrecht, 217 pp. Ross, C.A. and Ross, J.R.P., 1985. Late Paleozoic deposi-

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tkmal sequences are synchronous and worldwide. Geology, 13: 194-197. Ross. C.A. and Ross. J.R.P.. 1987. Late Paleozoic sea levels and depositional sequences. Spec. Publ. Cushman Found. Foram. Res., 24: 137-168. Ross. C.A. and Ross, J.R.P., 1988. Late Paleozoic transgressive-regressive deposition. In: C.K. Wilgus. B.S. Hastings, H. Posamentier, J. van Wagoner, C.A. Rcw~ and C.G.St.C. Kendall (Editors), Sea-Level Changes: An Integrated Approach. SEPM Spec. PUN., 42: 227-247. Rowsell, D.M. and De Swardt, A.M.J., 1976. Diagenesis in Cape and Karroo sediments, South Africa, and its bearing on their hydrocarbon potential. Trans. Geol. Soc. S. Aft., 79: 81-145. Rubidge. B.S. and Oelofsen, B.W., 1981. Reptilian fauna from Ecca rocks near Prince Albert. South Africa. S. Aft. J. Sci., 77: 425-426. Shanmugam, G. and Moiola, R.J., 1982. Eustatic control of turbidites and winnowed turbidites. Geology, 10: 231-235. Smellie, J.L., 1981. A complete arc-trench system recognized in Gondwana sequences of the Antarctic Peninsula region. Geol. Meg., 118: 139-159. Stavrakis, N., 1986. Sedimentary. environments and facie~ of the Orange Free State coalfield. In: C.R. Anhaeusser m d S. Maske (Editors), Mineral Deposits of Southern Africa. Geol. Soc. S. Aft., Johannesburg, pp. 1939-1952. Stow. D.A.V., Howell, D.G, and Nelson. C.H., 1984. Sedimentary, tectonic, and sea-level controls on submarine fan and slope-apron turbidite systems. Geo-Mar. Left., 3: 5764. Tankard, A.J., 1986. On the depositional response to thrusting and lithospheric flexure: examples from the Appalachian and Rocky Mountain basins. In: P.A. Allen and P. Homewood (Editors), Foreland Basins. Spec. Publ. Inl. Assoc. Scdimentol., 8: 369-392. Will. P.R., Mitchum. R.M. and Thompson, S., 1977. Seismic stratigraphy and global changes of sea level, Part 4. Global cycles of relative changes of sea level. In: C.E. Payton (Editork Seismic Stratigraphy--Applications to Hydrocarbon Exploration. Mere. Am. Assoc. Pet, Geol., 26: 83-97,

Van Vuuren, C.J., 1983. A Basin Analysis of the Northern Facies of the Ecca Group. Unpubl. Ph. D. Thesis, Univ. Orange Free State, Bloemfontein, 249 pp. Veevers, J.J. and Powell, C.M., 1987. Late Paleozoic glacial episodes in Gondwanaland reflected in transgressive-regressive depositional sequences in Eurameriea. Geol. Soc. Am. BulL, 98: 475-487. Viljoee. J.H.A., 199{). K-bentonites in the Ecca Group of the so'~lth and central Karoo Basin. Abstr. 23rd Congr. Geol, Soc. S. Afr., Cape Town, pp. 576-579. Vissm. J.N.J., 1989. The Permo-Carboniferous Dwyka Formation of southern Africa: deposition by a predominantly subpolar marine ice sheet. Palaeogeogr., Palaeoclimatot., Palaeoecol., 70: 377-391. Visser, J.N.J., 1990. The age of the late Palaeozoic glacigene deposits in southern Africa. S. Afr. J. Geol., 93: 366-375. Visser. J.N.L 1991a. Geography and climatology of the late Carbonifert,us to Jurassic Karoo Basin in south-western Gondwana. Ann. S. Afr. Mus.. 99: 415-431. Visser, J.N.J., 1991b. Self-destructive collapse of the PermoCarboniferous marine ice sheet in the Karoo Basin: evidence from the southern Karoo. S. Aft. J. Geol., 94: 255-262. Visser, J.NJ. and Dukas, B.A., 1979. Upward-fining fluviatile megacycles in the Beaufort Group, north of Graaff-Reinet. Trans. Geol. Soc. S. Afr., 82: 149-154. Visser, J.N.J. and Young, G.M., 1990. Major element geochemistry and paleoclimatology of the Permo-Carboniferous glacigene Dwyka Formation and post-glacial mudrocks in southern Africa, Palaeogeogr., Palaeoclimatot,, Palaeoecol., 81: 49-57, Wickens, H. de V.. 1990. The sedimentology of the Skoorsteenberg Formation and its applicabiiity :is a model for submarine fan turbidite dep,osits in the Bredasdorp Basin. Unpubl. Rep. Soekor. Parow, 47 pp. Wilson, T.J., 19~1. Transilion from back-arc to foreland basin development in the southernmost Andes: stratigraphic record from the Ultima Esperanza District, Chile. Geol. Soc. Am. Bull.. 1113: 98-111.