Chapter 8 The Tectonic Development of the Limpopo Mobile belt and the Evolution of the Archaean Cratons of Southern Africa1

Chapter 8

THE TECTONIC DEVELOPMENT OF THE LIMPOPO MOBILE BELT AND THE EVOLUTION OF THE ARCHAEAN CRATONS OF SOUTHERN AFRICA' J. M. BARTON JR. and R. M. KEY

ABSTRACT The tectonic development of the Limpopo Mobile Belt is reviewed and a plate-tectonic model is presented to explain this development and the evolution of the southern African cratons. It is postulated that the belt formed 3570 Ma ago with the establishment of a continental rise that collapsed into an aulacogen. The aulacogen and the sedimentary and igneous rocks that filled it were successively deformed as a result of differential movements between crustal plates encompassing what are now commonly termed the Rhodesian and Kaapvaal cratons. The timing of events in the evolution of the Limpopo Mobile Belt correlates well with the chronology of the tectonic development of the granite-greenstone terrains of the Rhodesian and Kaapvaal cratons. The pattern of crustal evolution in southern Africa may be viewed in terms of the accretion of a series of deformed back-arc basins and island arcs onto a nucleus of continental rocks including the Limpopo Mobile Belt.

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INTRODUCTION

Various attempts have been made t o interpret the evolution of Precambrian polymetamorphic terrains in terms of plate-tectonic models (see e.g. Anhaeusser, 1973; Burke and Dewey, 1973; Dewey and Burke, 1973; Hoffman, 1973, Talbot, 1973; Bridgwater et al., 1974; Hunter, 1974a, b; Tarney e t al., 1976; Wynne-Edwards, 1976; Martin and Porada, 1977a, b; Key, 1977; Groves et al., 1978; Barton, 1979a). However, the degree to which any such interpretation can provide a unique solution that is meaningful in detail depends on how well the evolution of the specific terrain is understood. It is rare that both the chronology and the physical conditions during the tectonic evolution of a Precambrian polymetamorphic terrain are understood a t all well, especially if the development began during Archaean times. The evolution of the Limpopo Mobile Belt of southern Africa is a notable exception. In this Chapter, the morphology and tectonic development of the Limpopo Mobile Belt as they are presently understood are described. The

A publication of the South African Contribution to the International Geodynamics Project.

186 constraints that these observations place on plate-tectonic models are discussed and a tentative plate-tectonic model is presented for the evolution of the belt. This model involves the formation of an aulacogen which was periodically deformed over 3500 Ma by differential movements between plates encompassing the Rhodesian and Kaapvaal cratons. While this platetectonic model may not provide the only conceivable explanation for the evolution of the belt it is, nevertheless, plausible and consistent with all the observed characteristics. Some of the implications of this model for the evolution of the continental crust of southern Africa are discussed.

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THE LIMPOPO MOBILE BELT

Morphology and lithology The Limpopo Mobile Belt (Mason, 1973) is a polymetamorphic terrain situated between the Rhodesian and Kaapvaal cratons of southern Africa (Figs. 8-1 and 8-2)(see also MacGregor, 1953;Cox et al., 1965). It may be conveniently divided into' three zones, a Central Zone bordered more or less symmetrically by Marginal Zones (Cox et al., 1965;Mason, 1973).Although these divisions were originally proposed on the basis of analysis of aerial photographs, subsequent work has proved them to be valid in modified form. The Marginal Zones may be shown by similarity of rock compositions and continuity of certain rock units to be primarily comprised of reworked equivalents of the granite-greenstone terrains of the adjacent cratons (see e.g. Bennett, 1971;Mason, 1973;Robertson, 1973, 1974, 1977;Ode11 and Phaup, 1975;Key et al., 1976;du Toit and van Reenen, 1977;van Reenen and du Toit, 1977). They are separated from the compositionally distinct supracrustal rocks of the Central Zone by major fault zones (Figs. 8-2 and 8-3)(see e.g. Mason, 1973;Coward et al., 1973;Key and Hutton, 1976;Key, 1977).The contact of the Southern Marginal Zone with the Kaapvaal Craton is gradational with respect to both metamorphic grade and degree of deformation (Graham, 1974;du Toit and van Reenen, 1977;van Reenen and du Toit, 1977, 1978). The contact of the Northern Marginal Zone with the Rhodesian Craton, originally thought t o be gradational (MacGregor, 1953; Robertson, 1968),is in part a zone of thrust faults dipping shallowly southward (James, 1975). However, in the west the original contact-has been altered by repeated deformation and may have been at one time a zone of thrust faults dipping northward (Key, 1977). The Marginal Zones are primarily comprised of leucocratic gneisses of variable but often tonalitic composition with small amounts of amphibolite, serpentinite, banded iron formation and other metasedimentary rocks preserved as synformal keels within the leucocratic gneisses. Conformable layers of K-feldspar porphyritic gneisses occur throughout the Northern Marginal

LEGEND POST- JURASSIC ROCKS KAROO SUPERGROUP

UMKONDO GROUP SOUTPANSBERG TROUGH Soutpansberg Group and Karoo Supergroup WATERBERG GROUP W= Woterberg Basin BUSHVEL D COMPLEX P= Pilonesberg Complex TRANSVAAL SUPERGROUP LIMPOPO MOBILE BELT NMZ= Northern Marginal Zone CZ=CentrolZone SMZ- Southern Marginol Zone RHODESIAN CRATON KAAPVAAL CRATON .r

I

Fig. 8-1. Generalized geologic map of southern Africa showing the exposed portion of the Limpopo Mobile Belt and the surrounding area (modified from Barton (1979a, fig. 1). M =Messina area. Coarse diagonal striping pattern shows the areas in the Marginal r-L a Zones containing orthopyroxene-quartz bearing metamorphic mineral assemblages (modified from Coward, 1976, fig. 1). 4

RHODESIAN CRATON

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FALILT AND SHEAR ZONES BOUNDING THE CENTRAL ZONE BOUNDARY OF GRANULITE FAClES ROCKS OF LIMPOPO MOBILE BELT

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Fig. 8-3. An aerial photograph of a portion of the Limpopo Mobile Belt near SelebiPikwe, Botswana, showing part of the Central Zone of multiply deformed rocks (south of the Letlhakane fault) and the northern marginal fault zone of the Central Zone. The marginal fault zone here is called the Tuli-Sabi straightening zone and an important fault in it is called the Letlhakane fault.

Zone and these may be correlated with the units of arkosic sandstone that form the lower parts of the stratigraphic sequences within the greenstone belts in the southern part of the Rhodesian Craton (Key et al., 1976). Fig. 8-2. Simplified geologic map of the Limpopo Mobile Belt and the adjacent cratons. The narrow horizontal striping pattern designates granitic and syenitic intrusions of diverse age. Except for the greenstone belts, the rocks of the cratons are undifferentiated but consist primarily of leucogranitic orthogneisses and granitoid plutons. The Marginal Zones consist of reworked rocks of the adjacent cratons. The Central Zone consists largely of a unique suite of supracrustal rocks, now present as gneisses.

190 Batholiths of syntectonic granitic rocks and thick sequences of volcanic rocks are largely lacking in the Marginal Zones. However, a line of plutons including the Matok and possibly the Schiel Complexes (Fig. 8-2)strikes obliquely into the Southern Marginal Zone from the Kaapvaal Craton (Barton e t al., 1981a). Those in the Southern Marginal Zone are deformed. The Central Zone near Messina (Fig. 8-2) is composed of a metamorphosed quartz-dioritic to granodioritic basement complex, including the Sand River gneisses (Barton e t al., 1977, 1978, 1 9 8 1 ~Barton ; and Ryan, 1977; Fripp, 1980), unconformably overlain by a sequence of supracrustal rocks.’ The basement complex contains a large component of probably metamorphosed greywacke (Barton, 1979b) intruded by gabbroic dykes (Fig. 8-4). The supracrustal rocks consist of metamorphosed sandstone, shale and carbonate rock with small amounts of banded iron formstion, metamorphosed basalt and possibly rhyolite (see e.g. Sohnge, 1945; Sohnge e t al., 1948; Jacobsen, 1967; Bahnemann, 1972; Mason, 1973; Light e t al., 1977; Light and Watkeys, 1977; Fripp e t al., 1979). The carbonate rock occurs primarily near the margins of the Central Zone, while the banded iron formation occurs most commonly near the centre, suggestive of sedimentary facies changes. Major differentiated sills of gabbroic rocks, emplaced a t high crustal levels and collectively termed the Messina Layered Intrusion, are found within the supracrustal rocks and several generations of gabbroic dykes intrude the supracrustal rocks and those of the Messina Layered Intrusion (Barton e t al., 1977, 1979a, 1981b). Deformed granitic bodies such as the Bulai pluton are unusual in this part of the Central Zone and are confined t o the area west of Messina (Fig. 8-2)(Sohnge, 1945; Sohnge e t al., 1948; Light e t al., 1977; Light and Watkeys, 1977). A major body of undeformed granite, the Mahalapye pluton, occurs within the Central Zone in eastcentral Botswana (Fig. 8-2). How far this succession of basement and cover rocks can be traced eastnortheastward of Messina is unknown. T o the west-southwest, however, compositionally similar rocks and rock successions occur as far as Baines T h e basement complex, originally recognized by Bahnemann (1971 ), has been variously termed and includes t h e Artonvilla Formation (Jacobsen, 1 9 6 7 ) and t h e Macuville Group in S o u t h Africa and Zimbabwe (see e.g. Light e t al., 1 9 7 7 ) . T h e supracrustal rocks were originally termed t h e Messina Formation (Sohnge e t al., 1 9 4 8 ) and later in South Africa and Zimbabwe t h e Beit Bridge G r o u p made u p of several sub-groups and formations (see e.g. Light e t al., 1977). Application of these divisions has proven to be nearly impossible owing to t h e difficulty of recognizing basement and cover rocks in outcrop and t h e large a m o u n t of deformation experienced by t h e rocks. In addition, the “basement rocks” exposed o n Farm Macuville are actually in part supracrustal rocks a n d in part a phase of t h e Bulai pluton. Some of t h e units grouped by Sohnge e t al. ( 1 9 4 8 ) i n t o t h e Messina Formation have been recognized as belonging t o t h e basem e n t complex. F o r these reasons, we have decided n o t to use formal stratigraphic names in this paper.

191

Fig. 8-4. A sketch map showing the relationship in the basement complex of the Limpopo Mobile Belt between the leucocratic and grey gneiss facies of the more than 3790 Ma old Sand River Gneisses and the approximately 3570Ma old mafic dykes (modified from Barton e t al., 1977, fig. 2 ) .

Drift and the confluence of the Seoka and Limpopo rivers (Fig. 8-2; Key, 1977, 1979). Rocks compositionally similar t o those of the supracrustal sequence also occur in the area around Selebi-Pikwe (Fig. 8-2; Wakefield; 1974, 1976), but large concentrations of metabasalt also occur here and the rocks that might be correlated with those of the Messina Layered Intrusion are anomalously rich in quartz (Key, 1977), possibly as a result of hydrothermal alteration. Near Mahalapye (Fig. 8-2) the stratigraphic succession appears to be quite different (Ermanovics, 1977), suggestive of a new and probably younger succession of rocks being exposed in the Central Zone west of approximately longitude 27"30' E. The fault zones bounding the Central Zone have had long and complex histories involving both strike-slip and dip-slip movements (see e.g. Mason, 1973; Coward e t al., 1976a; Key and Hutton, 1976). The distance separating the fault zones and hence the width of the Central Zone decreases from westsouthwest t o east-northeast (Fig. 8-2), and the apparent amounts of displacement along these fault zones increase going in the same direction. In addition, both fault zones bend slightly t o the north at a point in eastern Botswana (Figs. 8-2 and 8-5) and dip steeply to the south (Cox et al., 1965). The fault zones have been intermittently active over the history of the belt and they have controlled the pattern of more recent sedimentation in both

192 the Tuli and Soutpansberg Troughs (Jansen, 1975; Key and Hutton, 1976; Ermanovics e t al., 1978; Barton, 1979a). Fault zones conjugate t o those bordering the Central Zone and striking approximately from northwest t o southeast exist throughout the Central Zone (Fig. 8-2). These have also been intermittently active during the history of the belt. Gravity and magnetic character

PROTEROZOIC AND PHANEROZOIC ROCKS

4 BOUGUER GRAVITY .Oe.

/

/

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ISOGRAD ( 1 0milligal)

FAULT ZONES BOUNDING THE CENTRAL ZONE OF THE LIMPOPO MOBILE BELT

0

ORTHOPY ROXENE ISOGRAD

Fig. 8-5. Gravity pattern over the exposed portion of the Limpopo Mobile Belt (modified from Fairhead and Scovell, 1977, fig. 2).

193 broad lineaments of relatively high value, one along the northern margin of the Central Zone and the other along the northern margin of the Southern Marginal Zone (see e.g. Fairhead and Scovell, 1977; Reeves, 1977). In each instance the gravity high may be interpreted t o indicate the existence of a thinner than normal, low-density upper crust in these areas and the presence nearer t o the surface of higher-density rocks such as characterize the lower crust and upper mantle. The gravity highs become more subdued towards the west-southwest through the belt and disappear into a rather uniform but low pattern in eastern Botswana (Reeves and Hutchins, 1975). The gravity highs are continuous northeastwards as far as the gravity high associated with the Lebombo monocline which marks the eastern edge of the Kaapvad Craton. Although some of the gravity pattern in the eastern end of the belt may be attributable to the Lebombo monocline, it nonetheless suggests that a portion of the Central Zone has been thrust over the Northern Marginal

NMZ \

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Fig. 8-6. A generalized crustal cross section of the eastern portion of the Limpopo Mobile Belt showing an average gravity profile for comparison. Assumed ayerage rock densities are shown. N M Z = North Marginal Zone; CZ = Central Zone; S M Z = Southern Marginal Zone.

194 Zone and that a portion of the Southern Marginal Zone has been thrust over the Central Zone (Fig. 8-6). Field evidence for this thrusting may be found in Botswana (Key, 1977). This thrust faulting has resulted in the turning of the continental crust partially on edge in each case, exposing lower levels of crust along the northern margins respectively of the Central and Southern Marginal Zones. The amount of thrust faulting decreases t o nothing towards the west-southwest and the convergence and bending of the fault zones bordering the Central Zone reflect the increasing amount of thrusting towards the east-northeast. N o gravity high marks the northern margin of the Northern Marginal Zone, suggesting that any thrust faulting along this margin must have either been along shallowly dipping zones or of a minor amount. The magnetic and gravity patterns in Botswana (Reeves and Hutchins, 1975; Reeves, 1977, 1978a) show that the fault zone separating the Southern Marginal Zone from the Central Zone continues westward beneath the rocks of the Waterberg Group, Karoo Supergroup and Kalahari Group and is truncated by the Kalahari line (Reeves, 1977)(Fig. 8-7). The Kalahari line is a major crustal discontinuity that marks the western edge of the Kaapvaal and Rhodesian cratons. The fault zone separating the Central Zone from the Northern Marginal Zone becomes progressively less distinct westward and is unrecognizable west of about longitude 27'30' E (Key and Hutton, 1976; Reeves, 1978a). This may reflect a termination of this fault zone, in which case the rocks of the Central Zone and those of the Rhodesian Craton are continuous between about longitude 27'30' E and the edge of the Rhodesian Craton marked by the Kalahari line. Alternatively, and perhaps more reasonably, the gradual disappearance of the fault zone may reflect progressively deeper burial beneath younger rocks and this zone may actually continue at depth westward t o the Kalahari line. The Kalahari line itself becomes less distinct t o the north (Fig. 8-7) and is obscured by a feature termed the Makgadikgadi line (Reeves, 1977) which probably represents a fault zone along which younger rocks from the northwest are thrust southeastward over the rocks of the Rhodesian Craton. This thrust faulting did not distort the rocks containing the Kalahari line (Reeves, 1977). If this interpretation is correct, then the Central Zone of the Limpopo Mobile Belt completely separates the Rhodesian Craton from the Kaapvaal Craton and some of the rocks in the western portion of the central Zone, as suggested in the previous section, are younger than and overlie those in the east. Perhaps these younger supracrustal rocks are correlative with or constitute part of those in the basin postulated by Coward e t al. (197613) t o have occupied the region immediately to the north of the Central Zone. The destruction of this basin gave rise t o the upper greenstone belts in the southwestern portion of the Rhodesian Craton.

Metamorphism Large areas of the Marginal Zones are composed of rocks bearing the min-

195

I8"

24"

30"

12"

18"

24"

30"

36"

Fig. 8-7. A generalized tectonic map of southern Africa showing the inferred extent of the Rhodesian Craton (RC), the Central Zone of the Limpopo Mobile Belt ( L ) and the Kaapvaal Craton ( K C ) . K = Kalahari line. M = Makgadikgadi line. The stippled area shows that portion of the western Central Zone in which presumed younger supracrustal rocks are exposed.

era1assemblage orthopyroxene-quartz (Fig. 8-1)and have been metamorphosed under pressure and temperature conditions characteristic of the granulite facies (Robertson, 1968; van Reenen and du Toit, 1977, 1978). The grade of metamorphism declines, in some cases over a fairly short distance (van Reenen and du Toit, 1977), passing into the adjacent cratons where mineral assemblages characteristic of lower amphibolite facies or greenschist facies prevail. The grade of metamorphism also decreases towards the westsouthwest as does the width of the zone of granulite facies rocks (Figs. 8-1 and 8-2). Granulite facies rocks also occur in the zone of deformation in Botswana between the Matsitama greenstone belt and Pikwe (Fig. 8-2). In the Southern Marginal Zone the area containing the orthopyroxene-quartz assemblage may indicate the deepest level of the crust exposed. In the Central Zone it is unusual to find orthopyroxene in quartz-bearing rocks, the rocks characteristically containing amphibole and biotite. However, petrologic studies suggest that these rocks also have been subjected to pressures and temperatures characteristic of granulite metamorphism

196 (Bahnemann, 1972; Clifford, 1974; Schreyer and Abraham, 1976; Horrocks, 1980). Rocks in large areas of the Central Zone may have been metamorphosed under conditions of higher water proportion in the fluid phase than have been those of the Marginal Zones. In the Central Zone, as with the Marginal Zones, the grade of metamorphism appears to become lower in the west, perhaps reflecting in that direction rocks from higher crustal levels at the surface.

De forma tional his tory Locally the deformational histories of the rocks of the Marginal Zones of the Limpopo Mobile Belt are reasonably well understood (see e.g. Coward et al., 1973, 1976a, b; Graham, 1974; Hickman and Wakefield, 1975; James, 1975; Coward, 1976; Key et al., 1976; du Toit and van Reenen, 1977). These histories may be correlated with those of the adjacent cratons, although the regional intensities of the deformational events are variable, accounting for the sometimes different tectonic styles of the Marginal Zones. In the Northern Marginal Zone an early phase of recumbent folding preceded the formation of tight, upright north-south trending folds that are locally co-axially overfolded. This early folding formed the major nappes defined by the greenstone belts in the southern part of the Rhodesian Craton (Litherland, 1973) and was followed by a major period of shearing that formed the pronounced east-northeast-trending grain of the Northern Marginal Zone. A simple shear model with P,,, parallel t o the regional grain has been proposed for this second event (Key and Hutton, 1976). In addition, a rifting event preceded the intrusion of the Satellite Dykes of the Great Dyke Complex (Robertson and van Breemen, 1969). In the west, a young zone of deformation, spatially confined to a belt trending northwest to southeast between the Matsitama greenstone belt and Pikwe, intersects both the Rhodesian Craton and the Northern Marginal Zone and has affected both areas equally (Fig. 8-2). In the Southern Marginal Zone a relatively simple tectonic history has been preserved. Here tight, upright east-west-trending folds have been refolded and cut by shears of thrust faults which also trend from east to west or from northeast to southwest (du Toit and van Reenen, 1977). These deformational events gave a pronounced east-west grain to the Southern Marginal Zone. In the Central Zone the basement complex was penetratively deformed a t least once prior t o the deposition of the supracrustal rocks and was intruded subsequently by gabbroic dykes (Barton et al., 1977, 1979a). A basin then formed into which the sequence of supracrustal rocks was deposited, and this basin was subjected to tensional stresses during the emplacement of the Messina Layered Intrusion (Barton et al., 1979a). Following

197 this, the rocks of the Central Zone around Messina were deformed under compressional stresses during four major events (Barton e t al., 1979a). The first deformational event after deposition of the supracrustal rocks involved isoclinal folding and thrusting about approximately northsouth-trending axial planes that dipped gently t o the west. The second period of deformation which was the major fabric-forming event affecting both the supracrustal rocks and those of the Messina Layered Intrusion involved upright folding around steeply dipping northwesterly trending axial planes. Gabbroic dykes were intruded into these rocks after the second period of deformation. The Bulai pluton was emplaced after the second deformational event and was affected by and possibly intruded syntectonically during the third deformational event. This third event involved upright folding around axial planes nearly coplanar with those of the second deformational event. The fourth deformational event involved upright folding around east-west-trending axes and thrusting along east-west-striking axial planes. It was during this last major deformational event that the thrust faulting took place by which the rocks of the Central Zone over-rode those of the Northern Marginal Zone and the rocks of the Southern Marginal Zone over-rode those of the Central Zone. For the rocks northwest of Messina in Zimbabwe a similar tectonic history has been proposed by Light and Watkeys (1977) except that they postulate the emplacement of the Messina Layered Intrusion synchronous with the late stages of the first deformational event affecting the supracrustal rocks. Farther t o the west again a roughly similar tectonic history has been proposed for the Central Zone but different workers have disagreed on points of detail. A possible basement complex, which has been recognized near the confluence of the Seoka and Limpopo Rivers (Fig. 8-2), may have been penetratively deformed and then intruded by gabbroic dykes prior t o the deposition of the supracrustal rocks (Key, 1977). A basin then formed into which the supracrustal rocks were deposited, and these rocks in turn were intruded by the rocks of the Messina Layered Intrusion. The supracrustal sequence was tightly folded into recumbent structures with axial planes trending from north-northeast t o south-southwest (Wakefield, 1977). A suite of gabbroic dykes was emplaced prior t o a second deformational event that involved upright isoclinal folding about north-south-trending axes (Key, 1974; Wakefield, 1977). A third deformational event subsequently involved refolding of the upright folds about the same axial planes. All of the resulting interference structures plunge t o the west and the basement complex and the supracrustal rocks appear t o pass beneath a younger sequence of supracrustal rocks a t about longitude 27'30' E. Thrust and strike-slip faulting accompanied and followed the generation of the fold structures. Structures associated with the last deformational event near Messina have not been recognized in the rocks near Pikwe.

198 Chronology’ The rocks in the Marginal Zones of the Limpopo Mobile Belt, being metamorphosed equivalents of the granite-greenstone terrains of the adjacent cratons, are older than 3200 Ma (see e.g. Allsopp, 1961; Oosthuyzen, 1970; Hickman, 1974, 1976, 1978; Hawkesworth e t al., 1975; Davies and Allsopp, 1976; Key e t al., 1976; Moorbath et al., 1976; Barton, 1980). However, ages 2870Ma have been reported for rocks from the Northern only as old as Marginal Zone (Hickman, 1976, 1978) and ages only as old as 2650Ma have been measured for rocks from the Southern Marginal Zone (Barton and Ryan, 1977; Barton e t al., 1981a). In the Southern Marginal Zone nearly 2600 Ma and the postall of the deformed rock units have yielded ages of tectonic units yield slightly younger ages. I t appears probable, therefore, that the metamorphism that affected the Marginal Zones was sufficiently severe t o erase a t least the Rb-Sr whole-rock isotopic evidence for these older ages. The rocks of the Marginal Zones were eroded to approximately their present 1950Ma ago (Barton and Ryan, 1977; Jansen, 1977), and postlevel by kinematic gabbroic dykes were emplaced in the Southern Marginal Zone 1900Ma ago (Barton, 1979a). In marked contrast t o the Marginal Zones where metamorphism has obliterated the radiometric record of early tectonic events, a long and complex isotopic history covering c. 3800Ma is preserved in the rocks of the Central Zone near Messina and probably extending as far west as SelebiPikwe (Barton and Ryan, 1977). The basement complex is a t least 3790 Ma old and was intruded by gabbroic dykes 3570Ma ago after suffering a t least one deformational event (Barton and Ryan, 1977; Barton e t al., 1977, . supracrustal rocks were deposited sometime during the 1978, 1 9 8 1 ~ )The interval between 3570 Ma ago and 3270 Ma ago (Barton and Ryan, 1977; Barton e t al., 1977, 1979a). The Messina Layered Intrusion and other related rocks were intruded 3270 Ma ago and, near Messina, the first deformation of the supracrustal rocks occurred sometime between that time and 3150Ma ago (Barton e t al., 1979a, 1981b). The second deformational event occurred 3150 Ma ago and the Bulai pluton was probably emplaced 2700 Ma ago (van Breemen and Dodson, 1972; Barton e t al., 197913). The third de2700 Ma ago and 2600 Ma ago and formational event occurred between 2600Ma ago (Barton and Ryan, 1977; Barton the fourth event occurred e t al., 1979a, 1981a). A metamorphic event for which n o period of penetrative deformation has been assigned occurred 3000 k 50Ma ago (Barton and Ryan, 1977; Barton e t al., 1981b). This event was, however, accompanied by the emplacement of gabbroic dykes. High level granitic intrusions such as the Mahalapye pluton and some gabbroic dykes were

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All radiometric ages mentioned in this section have been calculated with the decay constants recommended by Steiger and Jager (1977).

199

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emplaced 2200Ma ago (van Breemen and Dodson, 1972; Barton, 1979a) and uplift and erosion t o nearly the present surface level was achieved 1950Ma ago (Barton and Ryan, throughout most of the Central Zone by 1977). Local thermal metamorphism and uplift preceded the formation of 1770Ma ago (Barton, 1979a) and some movethe Soutpansberg Trough ment along the fault zones bounding the Central Zone accompanied the deposition of the rocks of the Karoo Supergroup in the Soutpansberg and Tuli Troughs 180Ma ago (Manton, 1968; Barton and Ryan, 1977). In general, every tectonic event that has occurred on the Rhodesian and Kaapvaal cratons has been accompanied by some contemporaneous tectonic activity in the Central Zone of the Limpopo Mobile Belt (see the review in Barton and Ryan, 1977).

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A PLATE-TECTONIC MODEL

In order t o properly apply plate-tectonic models t o a polymetamorphic terrain such as the Limpopo Mobile Belt, it is necessary t o dissect the evolution of the terrain into distinct tectonic events. These events must be evaluated separately and the final model must be the sum of the separate solutions. It must be remembered that besides the crustal plates themselves, plate-tectonic models involve spreading zones of crustal creation, shprtening zones of crustal destruction and zones of predominantly strike-slip motion with minor crustal distortion. These last two zones form major crustal discontinuities that tend t o act as permanent zones of weakness that may be preferentially reactivated in response t o later crustal stresses. There is no conclusive evidence that oceanic crust ever existed in what is now the Central Zone of the Limpopo Mobile Belt. Ophiolites have not been recognized and large volumes of volcanic and intrusive rocks of the types that result from destruction of oceanic crust are conspicuously lacking except possibly for the local concentration of metavolcanic rocks west of Pikwe. Therefore, classic (i.e. post-Mesozoic type) subduction of oceanic crust in the evolution of the Central Zone of the Limpopo Mobile Belt seems unlikely unless it was achieved without subsequent magmatism. Furthermore, the lack of large volumes of volcanic and intrusive rocks precludes the possibility that a mantle plume is incorporated into the Central Zone. The earliest tectonic event that may be considered unequivocally t o be a unique part of the evolution of the Limpopo Mobile Belt is the development of a basin, floored b y continental crust, into which the supracmstal rocks 3570 Ma and of the Central Zone were deposited sometime between 3270 Ma ago. The basement complex, albeit older, is compositionally similar to the Archaean gneiss terrains on both the Rhodesian and Kaapvaal Cratons, and it may represent a lower level within the oldest parts of these cratons (Key e t al., 1976). The intrusion of gabbroic dykes 3570Ma ago into the basement complex is apparently coeval with mafic magmatism of the Barberton Mountain Land and the emplacement of the Sebakwian

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200 volcanic rocks of the Rhodesian Craton (see e.g. Jahn and Shih, 1974; Barton e t al., 1977; Wilson e t al., 1978; Hamilton et al., 1979; also Anhaeusser, this volume, Chapter 6 ; Nisbet e t al., this volume, Chapter 7, ed.). The sequence of supracrustal rocks, being composed primarily of clastic rocks of a continental origin and containing few rocks of volcanogenic or chemical provenance, is unique in composition and time of deposition to the belt. The only other sedimentary rocks in southern Africa recognized to possibly be coeval with it are those of the Fig Tree and Moodies Groups of the Barberton Mountain Land (Barton et al., 1977) and the Tutume Group of the western margin of the Rhodesian Craton (Litherland, 1973). However, these sedimentary rocks are derived primarily from volcanic source areas. The linear occurrence of carbonate rocks near the present margins of the Central Zone and the apparent lack of supracrustal rocks related to those of the Central Zone in the Marginal Zones (Key, 1977) suggest that the margins of the basin were not much different than those of the Central Zone today. The obvious plate tectonic analogue for such a basin is an aulacogen or a continental rift valley formed by the collapse of a continental rise emanating from a triple junction (Burke and Dewey, 1973). The fact that the fault zones bounding the Central Zone continue completely across the Rhodesian and Kaapvaal Cratons suggests that the triple junction existed outside of the belt at one end or the other. The second tectonic event of widespread significance in the development of the Limpopo Mobile Belt was the emplacement of the Messina Layered Intrusion into the supracrustal rocks 3270 Ma ago. This event may or may not have closely followed the actual deposition of the supracrustal rocks. However, the tensional regime during which it was intruded and the high initial 87Sr/86Srratio (0.7029, corrected for 3270Ma; see Barton e t al., 1979a) of the magma giving rise t o the Messina Layered Intrusion when compared to its age, are consistent with it having been emplaced during continental rifting or the formation of an aulacogen (Barton, 1979a; Barton et al., 1979a). The first three deformational events in the Central Zone affecting the rocks of the supracrustal sequence and the Messina Layered Intrusion are considered t o be the result of Riedel shear across the belt resulting from differential movements of the adjacent cratons or plates (see e.g. Coward, 1976; Coward et al., 1976a, b; Key and Hutton, 1976). This means that there have been large components of roughly parallel but opposite, motions between the plates encompassing the Rhodesian and Kaapvaal Cratons and that the rocks of the Central Zone have acted as a sort of shock absorber between the plates, deforming in response t o externally applied stresses. The first deformational event, occurring sometime between 3270 Ma and 3150 Ma ago, involved refolding about and thrust faulting along approximately north-south-trending axial planes and resulted primarily from rightlateral motion between the plates. The second deformational event, occurring 3150 Ma ago, and the third deformational event, occurring sometime

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201

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between 2700 Ma and 2600 Ma ago, involved refolding of the northsouth structures about, and thrust faulting along, approximately northwestsoutheast-trending axial planes and resulted from primarily left-lateral motion between the plates. It is improbable that in these three instances the motions of the plates were exactly parallel. Consequently, there was almost certainly a component of crustal shortening or extension occurring across the Central Zone during each of these deformational events (see e.g. Coward, 1976; Coward et al., 1976a, b). Components of crustal shortening or extension are responsible for the parallelism in the deformational histories of the Central and Marginal Zones (see e.g. Key and Hutton, 1976; Key, 1977). The fourth deformational event occurred 2600 Ma ago and was caused by a convergence of the plates across the eastern end of the belt, essentially perpendicular t o the Central Zone. This resulted in the thrusting of rocks of the Central Zone over those of the Northern Marginal Zone and of the rocks of the Southern Marginal Zone over those of the Central Zone. This thrust faulting occurred along the pre-existing fault zones bordering the Central Zone. Because, however, the thrust faulting dies out t o the westsouthwest, a clockwise rotation of the plate including the Rhodesian Craton with respect t o that including the Kaapvaal Craton of 15" is deduced, apparently occurring around the portion of the belt now exposed in eastern Botswana. This rotation and resulting thrust faulting caused the slight divergence of earlier fold axes and thrust faults in the Messina area with respect t o those in the Selebi-Pikwe area. It also explains why the last deformation in the Messina area is not manifested in the rocks of the SelebiPikwe area. Part of this rotation may have caused or resulted from the opening of the rift into which the Great Dyke Complex was intruded and the crustal thinning during the formation of the Witwatersrand Basin. It may also have caused or resulted from the deformation along northwestsoutheast-trending axes between the Matsitama greenstone belt and Pikwe. It may even have resulted in the fracturing of the southwestern part of the Rhodesian Craton along west-northwest axes. These fractures were reactivated during late Karoo times (Reeves, 197813). The rotation produced, however, both the convergence and bending of the fault zones bordering the Central Zone and a general tilting of the rocks of the Central and Southern Marginal Zones to the south and to the west. In both cases the crust was turned slightly on edge and deeper levels were exposed by subsequent erosion going northwards. In the Central Zone the basement complex was exposed. In the Southern Marginal Zone the orthopyroxene-quartz bearing granulites were exposed. The blanket 2600 Ma whole-rock ages from these rocks reflect their passing during thrust faulting into conditions of pressure and temperature where the radiometric clocks could start. Minor thrust faulting probably occurred in the Northern Marginal Zone t o accommodate some crustal shortening and this faulting may have broughtrocks of granulite grade t o the surface.

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202

It is important to note that even though the Limpopo Mobile Belt may appear surficially symmetrical with regard to the Central and Marginal Zones and even though i t may have been symmetrical before the last deformational event, this event imposed a major structural asymmetry on the belt. The Rhodesian and Kaapvaal cratons as well as the Limpopo Mobile Belt have acted as a reasonably stable plate since 2600 Ma ago and have formed a major structural domain termed by Clifford (1966) the “Kalahari Shield”. A cartoon depicting the plate-tectonic model presented here for the evolution of the Limpopo Mobile Belt is shown in Fig. 8-8 and a summary of the tectonic development is presented in Table 8-1.

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IMPLICATIONS F O R CRUSTAL EVOLUTION O F SOUTHERN AFRICA

Plate tectonics as a mechanism implies a global interaction in tectonic activity. The pattern of activity in any one location should not be viewed

A. >3570Ma

B. 3570 Ma C. 3570- 3350 Ma

D. 3350-3150 Ma E. 3150 Ma 8 2700-2600 Ma F. 2700-2600 Ma G. 1770 Ma 8i 200 Ma Fig. 8-8. A cartoon showing schematic block diagrams depicting the major tectonic events in the evolution of the Limpopo Mobile Belt (see also Table 8-1).N = north; R = Rhodesian Craton; L =Central Zone of t h e Limpopo Mobile Belt; K = Kaapvaal Craton. Horizontal pattern in Fig.8-8G shows t h e rocks of t h e Karoo Supergroup and t h e Soutpansberg Group. The parallel lines in Figs. 8-8D and 8-8E show t h e primary axes of folding and thrust faulting during the first three deformational events affecting the belt. Solid arrows denote directions of relative motion.

203 TABLE 8-1 The tectonic development of the Limpopo Mobile Belt Age *

* 180Ma * 730Ma f

1770Ma

* 1900Ma * 1950Ma

Description Reactivation of the Tuli and Soutpansberg troughs and the deposition of the rocks of the Karoo Supergroup (Fig. 8-8G). Intrusion of mafic dykes (Reeves, 1978b) Emplacement of the Beit Bridge kimberlite pipes (Allsopp and Kramers, 1977) Formation of the Tuli and Soutpansberg troughs and deposition of the rocks of the Soutpansberg Group (Fig. 8-8G) Deposition of the rocks of the Waterberg Group in the western portion of the belt and the emplacement of postkinematic gabbroic dykes in the Central and Marginal Zones Establishment of ubiquitous mineral radiometric ages possibly related t o the termination of uplift and erosion of the entire region

k 2220 Ma

Emplacement of post-kinematic gabbroic dykes in the Central Zone and intrusion of the Mahalapye batholith

2550 Ma

Emplacement of post-kinematic granitic plutons in the Southern Marginal Zone and of the Satellite Dykes of the Great Dyke Complex into the Northern Marginal Zone

f

2600-1950 Ma

Uplift and erosion t o nearly the present surface level, the resulting detritus filling the contemporaneously existing basins on the Rhodesian and Kaapvaal cratons. Possibly some deformation occurring in the Botswana segment of the belt

2700-2600Ma

The third period of deformation affecting the belt resulting from a left lateral motion between the Rhodesian and Kaapvaal cratons (Fig. 8-8E). The fourth period of deformation affecting the belt resulting from a clockwise rotation of the Rhodesian Craton with respect t o the Kaapvaal Craton (Fig. 8-8F)

* 2700 Ma

Emplacement of the Bulai pluton in the Central Zone, perhaps syntectonically with the third period of deformation affecting the belt

3050-2850 Ma

Metamorphic event of unknown significance and emplacement of mafic dykes

k 3150Ma

The second period of deformation and major fabric forming event in the supracrustal rocks. This deformation resulted from left lateral motion between the Rhodesian and Kaapvaal cratons (Fig. 8-8E)

3270-3150 Ma

The first period of deformation affecting the belt resulting from a right lateral motion between the Rhodesian and Kaapvaal cratons (Fig. 8-8D)

f

3270 Ma

3570-3350 Ma

* 3570Ma

Emplacement of the Messina Layered Intrusion in the Central Zone Formation of a fault bounded basin and emplacement of the sequence of supracrustal rocks characteristic of the Central Zone (Fig. 8-8C) Emplacement of gabbroic dykes in the basement complex, possibly as

204 Table 8-1 continued Age*

Description a result of formation of a continental rise in the Rhodesian-Kaapvaal Craton originating from a mantle plume (Figs. 8-8A and 8-8B).The establishment of the Limpopo Mobile Belt

f

3790Ma

> 3790 Ma

A major period of deformation of the Rhodesian-Kaapvaal Craton Deposition of a sequence of greywacke as a result of erosion of volcanic terrain. This volcanic terrain contained a large component of rocks formed by partial melting of the upper mantle and oceanic crust

* Based on the decay constants recommended

in Steiger and Jager (1977).

singly but instead as part of a total earth pattern. Unfortunately, with the present state of knowledge, this is difficult to recognize in detail for the past 200Ma of earth history and is nearly impossible to visualize for Precambrian time except in the very broadest terms. Nevertheless, if the model presented here is correct, it may provide some clues to the nature of the development of the Rhodesian and Kaapvaal cratons and some evidence about the style of tectonic activity during Archaean time. In order to form an aulacogen strong, reasonably thick continental crust must exist. Therefore, it may be inferred that the Limpopo Mobile Belt formed in a continental nucleus containing rocks a t least 3800 Ma old. The formation of the aulacogen resulted in the creation of two plates each containing some of the old continental crust. These remnants are now probably incorporated in the Rhodesian and Kaapvaal cratons although they have yet to be identified. The stresses that caused the deformation of the aulacogen were generated elsewhere within or adjacent to the plates encompassing the Rhodesian and Kaapvaal cratons. It is therefore useful t o look at the tectonic activity on these cratons that accompanied each deformational event in the Central Zone of the Limpopo Mobile Belt. The postulated formation of the triple junction and the consequent generation of a continental rise followed by an aulacogen or graben occurred at approximately the same time as the formation of the early granite-greenstone terrains of both cratons (see e.g. Hawkesworth e t al., 1977; Wilson et al., 1978; Hamilton e t al., 1979; Barton, 1980). Most of the greenstones contain mafic and ultramafic rocks that are believed to represent oceanic crust although in order t o explain the formation of high-magnesium basaltic and peridotitic magmas a close genetic relationship between these rocks and sialic rocks of a continental crust must be inferred (see e.g. Anhaeusser, 1973; Hunter, 1974a; McIver, 1975; Key e t al., 1976; Hawkesworth and O’Nions, 1977;,Groves et al., 1978). This suggests the probability that greenstone belts are actually remnants of oceanic crust formed during advanced rifting of continental crust (Hunter,

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205 1974a; Groves et al., 1978; see also Goodwin, this volume, Chapter 5, Kroner, this volume, Chapter 3, and Lambert, this volume, Chapter 18, ed.) or in back-arc basins between a continent and an island arc situated on either oceanic or continental crust (Tarney e t al., 1976). As such the Central Zone of the Limpopo Mobile Belt may have originally formed in much the same manner as did the greenstone belts, with the only difference that rifting ceased in the belt before oceanic crust was emplaced. This does not mean, however, that the evolution of the Limpopo Mobile Belt and of any given greenstone belt on the Rhodesian and Kaapvaal cratons were exactly coeval in every aspect. The only greenstone belt from which sufficient age data are available for comparison is the Barberton greenstone belt of the Kaapvaal Craton (see the synthesis of the tectonic evolution in Barton, 1980 and, for contrast, the discussion by Anhaeusser, this volume, Chapter 6). Here an oceanic crust composed of rocks of the Lower Onverwacht Group was formed 3550Ma ago (Hamilton e t al., 1979) in what we suggest t o have been a back-arc environment between an island arc and the continental nucleus ultimately t o contain the Limpopo Mobile Belt. The basin containing the oceanic crust was first filled with predominantly volcanic rocks (the Upper Onverwacht Group) and then by sediments of the > 3450Ma old Fig Tree and Moodies Groups (Barton, 1980). These sedimentary units were deposited in a narrow continental shelf-type environment as might exist along the margins of a fault bounded (rift) basin and they were derived from a rapidly eroding source area composed of both volcanic and continental rocks (Condie et al., 1970, Eriksson, 1979,1980). The amount of continental detritus increases upwards through the Fig Tree and Moodies Groups, suggesting that the source area was composed of continental crust overlain by volcanic rocks. This source area was located t o the south and, in a plate-tectonic context, is a good candidate for a volcanic terrain (island arc) situated on continental crust and derived from a subducting piece of oceanic crust. The basin was then compressed and intruded by plutons derived from the partial melting of oceanic crust similar t o that flooring the basin. During this deformational event the rocks of the basin were welded onto the continental nucleus. The elapsed time from the formation of the basin until welding of the basin rocks onto the continental nucleus was probably less than 100 (Barton, 1980). Subsequently, the deformed basin was intruded by plutons during three periods:, the first 3250 150Ma ago, the second 2900 k 150Ma ago and the third 2550 k 50Ma ago. In each instance the plutons were either derived directly from partial melting of oceanic crust or they represent remobilized material derived originally from partial melting of oceanic crust. The rocks exposed in the Barberton greenstone belt were never deeply buried and have been a t nearly their present crustal level for the past 3000 Ma. The formation of the Sebakwian greenstone belts of the Rhodesian Craton also occurred before 3500Ma ago (Moorbath e t al., 1976; Wilson et al.,

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206 1978; Wilson, 1979) but the detailed tectonic evolution of these rocks is poorly understood. However, the younger greenstone belts of the Rhodesian Craton appear t o have formed in a similar fashion t o the Barberton greenstone belt, and the span of time from the formation of a basin floored by oceanic crust until the destruction of that basin was only 100Ma from 2700Ma until 2600Ma ago (see e.g. Bickle e t al., 1975; Hawkesworth e t al., 1975; Coward e t al., 1976a, b). In addition, igneous and/or metamorphic activity, possibly associated with the formation of a granite2850 Ma ago greenstone terrain, occurred in the Rhodesian Craton (Hawkesworth e t al., 1979). The periods of destruction of these Archaean basins and the resulting formation of the greenstone belts coincide remarkably well with the periods of tectonic activity in the Central Zone of the Limpopo Mobile Belt, although the basin that ultimately became the Barberton greenstone belt may possibly have formed a t a slightly earlier time than that of the Central Zone, and the basin forming the Central Zone may have been created during the first deformation of the Barberton greenstone belt. This similar chronology suggests that the sialic crust of southern Africa was being stretched intermittently t o form basins from 3550 Ma until 2550 Ma ago. If the basins into which the rocks of the Witwatersrand triad, the Transvaal Supergroup, the Bushveld Complex and the Waterberg Group were deposited formed in a similar fashion, then this stretching continued until 1850 Ma ago. However, the early basins tended t o spread until the lithosphere fractured and oceanic crust was emplaced in them while the younger ones remained floored by continental crust, albeit thinner. In addition, the early basins were unstable and subject to destruction while the later ones were more permanent. The destruction of any given basin may have been a result of spreading t o form other basins.

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CONCLUSION

The Archaean tectonic evolution of southern Africa may be viewed adequately in terms of intermittent plate-tectonic activity. The Limpopo Mobile Belt formed as an aulacogen in a nucleus of continental crust onto which deformed back arc basins were accreted in the form of granitegreenstone terrains. However, as most tectonic features such as the Limpopo Mobile Belt and the greenstone belts are only partially preserved and imperfectly exposed, it is difficult t o evaluate the time scale involved in each tectonic event and the aria1 extent and thickness of the plates involved. Therefore, in reality one can only speculate as t o whether application of uniformitarian processes is appropriate or whether some other paradigm involving gradual evolution of processes might fit the observed features better.

207 ACKNOWLEDGEMENTS

Financial support for J.M.B. came from the Council for Scientific and Industrial Research of South Africa as part of the South African contribution to the International Geodynamics Project. This paper is published with the permission of the Director of the Institute of Geological Sciences, Keyworth, England, Dr. G. M. Brown. We thank H. L. Allsopp, M. J. de Wit, R. E. P. Fripp, N. C. Gay, J. V. Hepworth, D. R. Hunter, M. Jackson, C. B. Smith and M. K. Watkeys for critically reviewing earlier versions of this manuscript. We also thank 6. R. Jones, Director of the Geological Survey of Botswana, for his encouragement of this study. REFERENCES Allsopp, H. L., 1961. Rb-Sr age measurements on total rock and separated mineral fractions of the Old Granite of the central Transvaal. J. Geophys. Res., 66: 1499-1508. Allsopp, H. L. and Kramers, J. D., 1977. Rb-Sr and U-Pb age determinations o n southern African kimberlite pipes. Extended Abstracts, 2nd Int. Kimberlite Conf. 1977, Cape Town. Anhaeusser, C. R., 1973. The evolution of the early Precambrian crust of southern Africa. Philos. Trans. R. Soc. London, Ser. A, 273: 359-388. Anhaeusser, C. R . , 1981. Geotectonic evolution of the Archaean successions in the Barberton Mountain Land, South Africa. In: A. Kroner (Editor), Precambrian Plate Tectonics. Elsevier, Amsterdam, pp. 137-160 (this volume). Bahnemann, K. P., 1971. In E. R . Morrison and J. F. Wilson (Editors), 1971. Symposium on the Granites, Gneisses and Related Rocks. Excursion Guidebook, Geol. Soc. S. Africa, 44 pp. Bahnemann, K. P., 1972. A Review of the Structure, the Stratigraphy and the Metamorphism of the Basement Rocks of the Messina District, North Transvaal. D S c . Thesis, Univ. Pretoria, South Africa, 1 5 6 pp. (unpubl.). Barton, J. M., Jr., 1979a. The chemical composition, Rb-Sr isotopic systematics and tectonic setting of certain post-kinematic mafic igneous rocks, Limpopo Mobile Belt, Southern Africa. Precambrian Res., 9 : 57-80. Barton, J. M., Jr., 1979b. Crustal evolution clues. Nuclear Active, 21: 16--19. Barton, J. M. Jr., 1980. The pattern of Archaean crustal evolution in southern Africa as deduced from the evolution of the Limpopo Mobile Belt and the Barberton granitegreenstone terrane. Spec. Publ., Geol. Soc. Aust., 7 : in press. Barton, J. M., Jr. and Ryan, B., 1977. A review of the geochronologic framework of the Limpopo Mobile Belt. Geol. Surv. Botswana, Bull., 1 2 : 183--200. Barton, J. M., Jr., Fripp, R. E. P. and Ryan, B., 1977. Rb/Sr ages and geological setting of ancient dykes in the Sand River area, Limpopo Mobile Belt, southeyn Africa. Nature, 267: 487-490. Barton, J. M., Jr., Ryan, B. and Fripp, R. E. P., 1978. The relationship between Rb-Sr and U-Th-Pb whole-rock and zircon systems in the greater than 3790m.y. old Sand River Gneisses, Limpopo Mobile Belt, southern Africa (Abstr.). U S . Geol. Surv. Open File Rep., 78-701: 27-28. Barton, J. M., Jr., Fripp, R . E. P., Horrocks, P. and McLean, N., 1979a. The geology, age and tectonic setting of the Messina Layered Intrusion, Limpopo Mobile Belt, southern Africa. Am. J. Sci., 279: 1108--1134.

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