Cretaceous reactivation and intensified erosion in the Archean–Proterozoic Limpopo Belt, demonstrated by apatite fission track thermochronology

Tectonophysics 480 (2010) 99–108

Contents lists available at ScienceDirect

Tectonophysics j o u r n a l h o m e p a g e : w w w. e l s ev i e r. c o m / l o c a t e / t e c t o

Cretaceous reactivation and intensified erosion in the Archean–Proterozoic Limpopo Belt, demonstrated by apatite fission track thermochronology David X. Belton a,1, Matthias J. Raab a,b,⁎ a b

The University of Melbourne, School of Earth Sciences, Melbourne 3010, Australia Department of Primary Industries, GeoScience Victoria, GPO Box 4440, Melbourne 3001, Australia

a r t i c l e

i n f o

Article history: Received 16 May 2008 Received in revised form 19 September 2009 Accepted 24 September 2009 Available online 3 October 2009 Keywords: Zimbabwe Thermochronology Limpopo Transpression Reactivation Apatite fission track

a b s t r a c t Cratons are generally assumed to be regions of long-lasting tectonic stability. In particular the study of the Phanerozoic exhumation history of cratons has been largely hampered by the scarcity of suitable stratigraphic controls onshore. This fact is even more pronounced in terranes lacking Mesozoic or younger penetrative structural fabrics and metamorphic overprinting. Our study in the Limpopo belt shows that modern apatite fission track thermochronology provides a hitherto unavailable perspective in the study of these rocks, and has profound implications for the crustal evolution of the Zimbabwe Craton. Apatite fission track data from 35 samples taken along two transects, in the southern edge of the Zimbabwe Craton and in the Central Zone of the Limpopo Belt, suggest that extensive regions experienced kilometerscale exhumation in two discrete events, as recently as the Cretaceous. The first occurred at around 130 Ma, and the second at around 90 Ma. Basin subsidence and sedimentation loads on the Mozambique margin support the timing of these events and provide strong indications of the source and pathways for the eroded material. Further, the results indicate that young and old “surfaces” (in a geomorphological sense) may be structurally juxtaposed in regions of high elevation in Zimbabwe. This is contrary to early ideas of surface chronologies based on summit accordances or invoking pediplanation. © 2009 Elsevier B.V. All rights reserved.

1. Introduction Cratons, considering their extraordinary age and preservation, are frequently assumed to be regimes of enduring tectonic stability (Park and Jaroszewski, 1994). Here, we present evidence that rebut ideas of absolute cratonic stability, particularly within the context of the Zimbabwe Craton. While work on the Archaean and Proterozoic formation and tectonic evolution of cratons has made remarkable advances (Brandt and de Wit, 1997; Reeves et al., 2004; Treloar et al., 1992; van Reenen, 1995), the study of Phanerozoic tectonic history of cratons has largely been hampered by the absence of suitable stratigraphic controls onshore. Many of these terranes also lack younger penetrative structural fabrics and have not been exposed during the Mesozoic to conditions likely to cause metamorphic overprinting (Marshak et al., 1999). This limitation is immediately obvious given the lack of studies on Cretaceous histories of Southern Africa's exposed cratonic basements. Nevertheless, the geomorphological community invested considerable effort in trying to under-

⁎ Corresponding author. Present address: Department of Primary Industries, GeoScience Victoria, GPO Box 4440, Melbourne 3001, Australia. Tel.: +61 3 9658 4585; fax: +61 3 9658 4555. E-mail addresses: [email protected] (D.X. Belton), [email protected] (M.J. Raab). 1 Fax: +61 3 8344 7761. 0040-1951/$ – see front matter © 2009 Elsevier B.V. All rights reserved. doi:10.1016/j.tecto.2009.09.018

stand the development of Southern Africa's landscape (Du Toit, 1933; King, 1967; Partridge and Maud, 1987). More recently the emphasis has been put on clarifying the development of ancient and modern drainage systems (de Wit et al., 2000; Moore, 1999) within the frame>work of erosion surfaces and flexural hinges. Partly for this reason, established landform evolution models established by King (1976, 1967) and more recently re-enforced by Partridge and Maud (1987) in southern Africa long after the underlying assumptions have elsewhere been shown to lack validity (Belton et al., 2004). Within cratons, the problems are further complicated by the difficulty of identifying the presence, let alone the timing, of even modest structural offset within granitic and gneissic lithologies. Until recently there has been little consensus about the timing and magnitude of uplift and denudation in southern Africa because researchers investigating southern Africa's geomorphic evolution traditionally placed great emphasis on field observations. Alexander Du Toit was famous for his remarkably detailed suite of notes (Partridge, 1998). King (1967) and those who followed, such as Lister (1979, 1987) and Partridge and Maud (1987) also placed great emphasis on the collection of robust data and field evidence. The research presented here follows a number of studies in southern Africa using onshore apatite fission track thermochronology (Brown et al., 2000; Gallagher and Brown, 1999; Raab et al., 2002, 2005; Tinker et al., 2008) and provides new insights into the crustal evolution of the Zimbabwe Craton.

100

D.X. Belton, M.J. Raab / Tectonophysics 480 (2010) 99–108

2. Geology — from craton to mobile belt The Limpopo Belt (Fig. 1) separates the Archaean Zimbabwe Craton from the Kaapvaal Craton in South Africa and is one of the most extensively studied mobile belts in the world (Reenen et al., 1992) High-grade metamorphic rocks (granulites and retrograde amphibolite facies) predominate and are intensely deformed. They are exposed in an area some 600 km by 300 km between the lower grade (typically greenschist facies) rocks of the Zimbabwe and Kaapvaal cratons. The Limpopo Belt's Archaean rocks have been subdivided into three structural zones (McCourt and Vearncombe, 1992) (Fig. 2). The Central Zone (CZ) is dominated by tonalitic and leucocratic gneisses with marbles, paragneisses and quartzites forming an Archaean cover sequence intruded by gabbroic and anorthosite units (Robertson and Du Toit, 1981). Parts of the Central Zone show a distinct horizontal lineation interpreted to be the result of strike–slip motion from ca. 2 Ga (Kamber et al., 1995). Large, mylonitic shear zones with a pronounced foliation differentiate the Central Zone from the adjacent Marginal Zones. In turn, the boundaries of the Northern and Southern Marginal Zones (NMZ and SMZ), within the cratons proper, are largely defined by shear zones with a thrust sense verging southeast in the SMZ and verging northwest on the Umlali Thrust zone (Mkweli et al., 1995) that marks the extremity of the NMZ (Fig. 2). Thrusting on the Kaapvaal Craton was dated by Kreissig et al. (2001) at 2.6 Ga for the NMZ and 2.7 Ga for the SMZ. Phanerozoic rocks that may have provided a more detailed and regional consistent stratigraphic record about periods of enhanced denudation are either absent or restricted to localised basins where exposure is severely limited. Much of the Palaeozoic and Mesozoic stratigraphy in these basins is obscured by what is essentially a diachronous unit — the post-Karoo basalts. For the cratons bordering the Limpopo Belt, little or no sedimentary record is preserved between the Late Proterozoic and the early Mesozoic, and over large areas, even the latter units are absent. The same basins are characterised by Karoo sediments of Permo-Triassic age, developed in rift basins prior to the Gondwana breakup (Visser, 1989). The lowermost sediments are Permian-age diamictites or conglomerates of apparent glacial origin (Dwyka Beds) resting on Precambrian granulite basement. The Lower Karoo sequence develops as a series of argillaceous and carbonaceous shales and mudstones, the older layers hosting

Fig. 2. Geology of southern Zimbabwe with selection of sample locations. See also crosssections for Profile “A” and “B” in subsequent Fig. 4 (adapted from Stagman, 1978).

economic coal seams. These are separated from the Upper Karoo sequence by a minor unconformity identified at some locations — elsewhere the contact is conformable (Stagman, 1978). The upper

Fig. 1. The tectonic history of the Zimbabwe Craton is intimately linked to the surrounding mobile belts that focused much of the deformation during successive episodes of orogenesis throughout the Archaean and Proterozoic.

D.X. Belton, M.J. Raab / Tectonophysics 480 (2010) 99–108

units are characterised by coarser sediments including grits at the base, a series of red beds and fine grained, creamy-coloured sandstone at the top (the Clarens Sandstone). These sediments are generally intercalated with basaltic volcanics of the Stormberg Group (183 Ma– 179 Ma) (White, 1997) associated with the breakup of Gondwana. On the Zimbabwe craton proper, only a single, very restricted outcrop of the Karoo volcanics remains. Near Featherstone, 90 km south of Harare (Fig. 3), a small outcrop of vesicular basalt lies for the most part, directly on the granite, and where present, the Karoo sediments are reduced to a thin, intermittent, metre-scale unit of poorly indurated siltstone. Widespread normal faulting postdates the volcanism and, in the Limpopo, is responsible for the confinement of Karoo units to a series of rhomb-shaped basins. In the east, along the border with Mozambique, these basins are unconformably overlain by extensive sheets of the Malonga Formation, a sequence of undeformed clastic sediments of mid-Cretaceous age (Botha and de Wit, 1996). 3. Geomorphology of the study area In contrast to the drainage divide of central Zimbabwe which lies above 1250 m, most of the region south of Masvingo drops from the granite–greenstone landscape at around 900 m, down to the Northern Marginal Zone and the Limpopo basin, at an elevation of between 300 m and 600 m. Large inselbergs or whalebacks are abundant and result from the weathering of a granitic unit segmented by sets of orthogonal joints which have markedly influenced the local drainage. One of the primary joint/fault orientations parallels the trend of the Limpopo Belt. The most abrupt relief in the Masvingo area tends to be localised in areas of resistant lithologies within the greenstone belts. Further southeast, in the area of the Triangle Shear Zone (Fig. 2), local relief is very subdued, largely due to the development of mylonitic fabrics during Precambrian tectonics (Kamber et al., 1995), with only occasional rises or kopjes providing opportunities for sampling basement exposure. In the Limpopo Valley, a steeper topography is found

101

in areas of late-stage, post-Karoo intrusives (∼ 180 Ma) throughout the Nuanetsi Province including the ring-complexes of Marangudzi, Mateke Hills and Mutandawhe. On the South African side of the border, relief in areas of exposed basement is still subdued, but good exposure is found in river channels. Around Messina, and elsewhere in the intensely deformed Tchipise region, basement rocks dominated by Archaean quartzites produced isolated but prominent hills of several hundred metres. Much of the central and southern Limpopo valley is characterised by extensive fault scarps of Triassic sandstone (the Clarens Sandstone) running west–south–west. Further south, elevations rise rapidly from 750 m typical of the Limpopo Valley, up to some 1600 m in the Soutpansberg Ranges (Fig. 4). 4. Sampling strategy The apatite fission track samples from the southern Zimbabwe craton include two suites of samples (Fig. 2). The first suite consists of 12 samples from a traverse running south–southeast from the Mwenezi Range near Featherstone on the central Zimbabwe Craton, to the Limpopo River in the Central Zone (CZ). The traverse crosses the Northern Marginal Zone (NMZ) orthogonally. The sampling of the thrusted cratonic margin of the NMZ and the boundary between the NMZ and the Central Zone to the south was undertaken for two reasons, 1) to establish if the margin shared a common history with the central craton and 2) to identify Mesozoic reactivation (if any) on these Precambrian structures delineating the zones. The second suite of 23 samples of the Central Limpopo traverse was focused primarily in the Central Zone (CZ) of the Belt, running roughly east to west in a direction parallel to the structural trend of the mobile belt. These samples were selected to document the timing and magnitude of exhumation of the Limpopo Belt. This was aimed specifically at regional responses to global tectonic events during and since the Mesozoic. The strategy was intended to identify variations (if any) between the core of the craton and its margins and at the same time, examine the potential for reactivation of major structural

Fig. 3. Shaded relief topography of Zimbabwe indicating the locations of samples and traverses discussed.

102

D.X. Belton, M.J. Raab / Tectonophysics 480 (2010) 99–108

Fig. 4. Geological cross-sections of the geology of southern Zimbabwe and South Africa with the apatite fission track and TASC results for Profile “A” and “B” (Fig. 2) shown in their structural and lithological context. SMZ (Southern Margin Zone), CZ (Central Zone), and NMZ (Northern Margin Zone).

elements of the Archaean Limpopo Mobile Belt (Robertson and Du Toit, 1981) and of the younger sub-vertical faults that characterise the region. Within the Central Zone of the Limpopo Belt, the sampling strategy was aimed at determining the thermal history of a series of faultbounded Karoo basins in the Limpopo valley. Sampling within the Karoo sedimentary units in the Limpopo Basin was undertaken primarily in argillaceous and carbonaceous units since these were most likely to yield apatite. 5. Methodology 5.1. Apatite fission track thermochronology (AFTT) AFTT is a well established technique for estimating broad, regionalscale patterns of denudation (e.g. Gleadow et al., 2002). It is most useful for detecting and measuring amounts of denudation of at least 3–5 km of the Earth's surface over periods of the order of several million years. The thermal history of a sample is determined by analysing the preserved spontaneous fission tracks of 238U in apatite. These tracks, which are continuously produced, shorten as a function of temperature. Tracks are fully annealed above 110 °C and the chronometer is reset; whereas below 60 °C, the annealing process slows to the extent that the distribution of track lengths, reflecting the 110–60 °C cooling history is preserved (Gleadow et al., 2002). The temperature range between 60 °C and 110 °C is termed the partial annealing zone (PAZ). AFTT is most closely related to the manifestation of tectonic movements and the rates of surface denudation.

5.2. Constraining cooling onset ages with the track age spectrum calculator (TASC) The Track Age Spectrum Calculator (TASC) approach developed by Belton et al. (a contribution in Ehlers et al., 2005) is an advanced tool to extract additional information from individual confined track length data and the apparent fission track age of a sample. TASC derives the age of onset of cooling, which reflects the time a sample passed through the 110 °C isotherm. Essentially, this temperature defines the time the apatite fission track system begins to retain tracks and starts to record thermal history information. By recalculating track densities and length-dependent probabilities, each track can be allocated to an equivalent age and a “track age spectrum” can be derived. The track age spectrum is independent of chemistry and mineralogy and the ages of onset of cooling in combination with other indicators such as cooling style and timing provide robust time– temperature nodes as input parameters for further thermal (inverse) modelling. This is particularly useful in areas with limited stratigraphic and structural controls, such as deeply eroded terranes. 5.3. Thermal modelling We used inverse modelling of cooling histories on fission track data to provide quantitative constraints on time–temperature. We used the maximum likelihood approach by Gallagher (1995) to test and quantify possible thermal histories using single grain age distributions and track length distributions. This approach approximates the thermal history by an unknown series of points in the time–

D.X. Belton, M.J. Raab / Tectonophysics 480 (2010) 99–108

temperature space and, using a genetic algorithm, searches this space to find the thermal histories which provide good fits to the observed data (as measured by the likelihood function).

103

11.49 µm and 12.39 µm. Only one other sample had such a short mean track length. Sample RWB96-11 which was collected in close proximity to the RWB96-12 and -13, had a young age (89.7 ± 4.1 Ma) but with a mean length of 11.77 µm.

6. AFT results 7. TASC analysis AFT results for the southern Zimbabwe Craton traverse are presented in Table 1 and shown in the cross-sections in Fig. 4. With the exception of one sample (Z96-003), ages from the mid-craton are typically old (280 ± 22 Ma at ± 1σ for Z96-011) with short, mean confined track lengths (12.00 ± 0.14 µm) and trends younger towards the mobile belt with ages down to 89 ± 3.7 Ma (Z96-116) in the basement of the Central Zone (CZ). Confined mean track lengths in this part of the CZ remained relatively short at 12.80 ± 0.11 µm for Z96-116. One basement sample (Z96-115) adjacent to the faultbounded Tuli basalts of the Karoo basin had an age of 141 ± 15 Ma with slightly longer lengths of 13.34 ± 0.2 µm. The source rocks for most of the analysed apatites were basement granites and gneisses. The age trend follows a pattern of age increase with elevation, from 280 Ma at an elevation of 1479 m (on the central watershed) to 89 Ma at an elevation of 482 m in the Limpopo Valley. In general the short mean track lengths with a multi-phased cooling history are consistent with a protracted and potentially complex history. In contrast, the mid-craton sample, Z96-003 gave a surprisingly young age of 116 ± 22 Ma with a mean track length of 11.67 ± 0.39 µm (based on only 13 tracks from 13 grains). This sample came from a moderately lithified sandy sediment stratigraphically below one of the few rare outcrops of Stormberg age basalt (ca. 170 Ma) close to Featherstone. The crystalline basement sample, Z96-011, was recovered from a slightly higher elevation about 100 km further south but records a fission track age of 280 ± 22 Ma. AFT results for the Central Limpopo traverse are presented in Table 2. While maintaining the trend observed in the southern craton samples, all but two of the CZ samples returned post-Karoo ages between 80 Ma and 150 Ma with longer, mean confined track lengths on the order of 12.9 µm to 13.9 µm. Only three samples (98RWB-27, -28 and -29) were recovered from the sedimentary remnants of the Karoo formation. These were invariably poor quality samples (cf. Z96003 above) with only 3–5 grains each and no confined fission tracks were observed. All the remaining samples were sourced from basement gneisses and granitic rocks and had sufficient grains to generate robust ages and provide abundant confined tracks for length analysis. Within the CZ, two samples RWB96-12 and -13 gave older ages (228 Ma and 213 Ma respectively) with shorter mean lengths of

Track age spectra were calculated for all samples and these results are summarised in Tables 3 and 4. For the smallest of the two suites of samples collected from the central Craton towards the mobile belt in the south, the TASC analysis produced a maximum cooling onset age of 464 ± 42 Ma for sample Z96-112 on the central watershed, trending down to a minimum cooling onset age of 111 ± 4 Ma for sample Z96-116 within the Central Zone. The track density multiplier (ρs) is the factor by which the original fission track densities are adjusted to account for each track's probability of intersecting the polished surface of a sample. It is used to generate a measured age of passage through the 110 °C isotherm (ie. the cooling onset age). For the Central Zone, track density multipliers typically lie between 1.3 and 1.6, reflecting the complex track length histograms for these samples. This implies that the samples towards the central craton may have passed through the bottom of the PAZ (nominally the 110 °C) before the Permo-Triassic sediments of the Karoo sequences began accumulating, whereas those in the Central zone only entered the PAZ in the early Cretaceous. The track age spectrum for Z96-115 is shown in Fig. 5. This sample from the Northern Marginal Zone (NMZ) has characteristics of both the protracted history of the inner craton and the more recent Cretaceous events. As noted above, sample Z96-003 is somewhat enigmatic. TASC analysis of this sample is subject to large error because of the low reliability of fission track age. Nevertheless, it is interesting that the cooling onset age it records (157 ± 30 Ma) is only marginally younger than the eruption age of the Stormberg basalts (ca. 182 Ma–White, 1997) — remnants of which lie immediately above this site. This may indicate thermal resetting of the apatite fission track age as a direct result of the basalt emplacement. The nearby sample Z96-011 records a more typical craton cooling onset age of 397 ± 30 Ma, suggesting that the immediate thermal effects of such an eruption are likely to be restricted to rocks in close proximity to the emplaced magma. This would be in line with predictions from the thermal modelling of Brown et al. (1994). The event spectra are used to summarise all the cooling steps identified in individual track age spectra (Ehlers et al., 2005). Fig. 6A is the event spectrum for the NMZ, and indicates that two of the older

Table 1 Apatite fission track results for the Southern Zimbabwe Craton Traverse. Sample

Elevation (m)

No. of grains

ρs (x 106 cm− 2)

Ns

ρi (× 106 cm− 2)

Ni

P(χ2) %

U (ppm)

Mean ρs/ρi ± 1σ

ρd (× 106 cm− 2)

Nd

Age ± 1σ (Ma)

Mean track length (µm)

Std. dev (µm)

No. of lengths

Z96-03 Z96-11 Z96-112 Z96-113 Z96-114 Z96-115 Z96-116 Z96-117 Z96-123 Z96-127 Z96-128 Z96-146

1388 1479 828 818 706 712 482 492 451 691 721 977

9 21 20 20 19 20 20 20 16 15 20 19

1.755 1.773 1.844 2.781 2.655 4.698 3.172 1.7 1.172 1.873 1.548 1.192

86 1220 1103 1788 1314 140 1913 471 674 384 1000 1337

3.551 1.442 1.196 2.401 2.875 7.617 8.337 3.675 2.419 4.463 1.653 2.61

174 992 715 1544 1423 227 5027 1018 1391 915 1068 2928

5.3 0 5.8 1.2 0 87.3 0.3 6.9 2 0 6 19

32.7 14.3 12.4 24.5 28.7 7.4 79.9 34.5 25.9 46.9 17 26.4

0.491 ± 0.103 1.350 ± 0.129 1.716 ± 0.118 1.200 ± 0.062 1.028 ± 0.079 0.667 ± 0.061 0.384 ± 0.015 0.456 ± 0.032 0.466 ± 0.035 0.455 ± 0.045 0.953 ± 0.070 0.521 ± 0.062

1.357 1.261 1.201 1.227 1.253 1.279 1.305 1.331 1.167 1.19 1.214 1.237

3887 3770 3887 3887 3887 3887 3887 3887 3770 3770 3770 3770

115.7 ± 22.2 279.6 ± 22.1 331.8 ± 20.5 252.2 ± 12.6 211.9 ± 14.7 140.8 ± 15.3 89.0 ± 3.7 108.4 ± 7.7 98.1 ± 6.8 92.1 ± 11.3 201.2 ± 10.5 106.0 ± 7.7

11.67 12 11.36 10.75 11.57 13.34 12.8 12.84 12.24 11.51 10.96 12.64

1.42 1.47 1.84 1.92 2.24 1.65 1.43 1.61 2.11 2.41 1.76 1.94

13 104 101 103 103 66 155 89 33 29 88 57

Standard track densities (ρD) and induced track densities (ρI) were measured on mica external detectors and fossil track densities (ρS) on internal mica surfaces. Nd, Ns and Ni are the numbers of tracks counted. Age is the central age, %Var is percent variation from pooled age. The pooled age is quoted for samples where P(χ2) N 5%, otherwise the central age is given. Ages were calculated using a zeta of 362.4 ± 12 (DXB) using Corning CN5 dosimeter glass.

104

D.X. Belton, M.J. Raab / Tectonophysics 480 (2010) 99–108

Table 2 Apatite fission track results for the Central Limpopo Traverse. Sample

No. of grains

ρs (× 106 cm− 2)

Ns

ρi (× 106 cm− 2)

Ni

P(χ2) %

U (ppm)

Mean ρs/ρi ± 1σ

ρd (× 106 cm− 2)

Nd

Age ± 1σ (Ma)

Mean track length (mm)

Std. dev (mm)

No. of lengths

96RWB-07 96RWB-08 96RWB-09 96RWB-10 96RWB-11 96RWB-12 96RWB-13 96RWB-14 96RWB-15 96RWB-16 96RWB-17 96RWB-19 96RWB-27 96RWB-28 96RWB-29 96RWB-31 96RWB-36 96RWB-37 96RWB-39 96RWB-40a 96RWB-40c 96RWB-41 96RWB-42

20 20 20 20 20 20 10 21 20 13 20 20 5 3 2 20 21 20 20 28 20 20 21

4.948 1.345 0.8 1.55 9.954 2.712 3.194 8.7 0.576 1.555 0.364 1.035 1.542 1.577 0.633 0.917 0.711 0.997 1.457 0.714 1.928 0.547 0.831

2790 1509 1054 1786 (1179 1394 1392 1152 638 1373 427 1431 75 217 32 893 1033 1358 1952 649 1646 939 615

8.199 2.626 1.563 3.488 2.098 2.245 2.928 2.414 1.125 2.54 0.749 2.448 2.529 3.176 1.701 2.216 1.918 2.37 3.221 1.489 4.232 1.352 1.768

4623 2946 2059 4020 2485 1154 1276 3196 1246 2243 880 3428 123 437 86 2071 2787 3231 4316 1354 3616 2319 1308

0.4 20.5 6.2 0 11.8 0 0 0.3 17.9 36.1 99.6 46.6 2 24.7 52.3 5.3 15.2 9.5 17.5 27.8 68.8 61.7 0.3

106 33.2 19.4 42.3 24.5 26.1 63.2 27.7 12.8 28.5 8.3 27.3 27.5 34.2 19 23.3 20.7 25.2 33.6 15.3 42.8 13.5 17.3

0.624 ± 0.022 0.514 ± 0.018 0.549 ± 0.031 0.464 ± 0.021 0.485 ± 0.023 1.332 ± 0.123 1.173 ± 0.116 0.372 ± 0.021 0.531 ± 0.034 0.664 ± 0.036 0.472 ± 0.020 0.423 ± 0.014 0.871 ± 0.242 0.532 ± 0.061 0.376 ± 0.049 0.433 ± 0.023 0.382 ± 0.018 0.418 ± 0.019 0.494 ± 0.023 0.499 ± 0.029 0.469 ± 0.014 0.405 ± 0.017 0.477 ± 0.039

9.669 9.883 1.009 1.031 1.052 1.074 1.011 1.091 1.102 1.114 1.125 1.137 1.148 1.16 1.12 1.139 1.157 1.178 1.197 1.217 1.236 1.255 1.274

3607 3658 3607 3607 3607 3607 4189 4189 4189 4189 (4189 4189 4189 4189 4783 4783 4783 4783 4783 4783 4783 4783 4783

105.6 ± 4.0 90.6 ± 3.6 94.2 ± 4.5 83.9 ± 4.0 89.7 ± 4.1 228.0 ± 24.0 213.2 ± 17.8 75.5 ± 3.9 108.0 ± 6.2 129.6 ± 4.9 104.0 ± 6.3 90.5 ± 3.2 147.3 ± 35. 109.7 ± 9.3 79.5 ± 16.5 93.3 ± 4.8 81.9 ± 3.6 93.4 ± 3.8 104.4 ± 3.6 111.1 ± 5.7 107.2 ± 3.5 96.8 ± 4.0 113.9 ± 8.5

13.73 13.64 13.87 13.73 11.77 11.49 12.39 12.96 13.17 13.66 13.87 13.78

1.7 1.28 1.76 1.39 2.71 2.29 2.67 1.88 2.66 2 1.05 1.31

100 103 102 100 63 100 100 100 75 100 82 100

13.3 13.18 12.95 13.33 13.87 13.17 13.1 13.46

1.47 1.71 1.53 1.4 1.22 1.36 1.66 1.78

100 100 100 100 100 100 100 41

Standard track densities (ρD) and induced track densities (ρI) were measured on mica external detectors and fossil track densities (ρS) on internal mica surfaces. Nd, Ns and Ni are the numbers of tracks counted. Age is the central age, %Var is percent variation from pooled age. The pooled age is quoted for samples where P(χ2) N 5%, otherwise the central age is given. Ages were calculated using a zeta of 362.4 ± 12 (DXB) using Corning CN5 dosimeter glass.

samples show a cooling at around 300 Ma, however all samples show a period of cooling at around 123 Ma. The six samples closer to the mobile belt also record a cooling in the Late Cretaceous (mean around ∼ 83 Ma). Four of the samples show a very late Palaeogene cooling. Results of the TASC analysis for the second, larger suite of samples from within the Central Zone are shown in Table 4. With just three exceptions, all 20 samples record Early Cretaceous cooling onset ages ranging from 137 Ma to 94 Ma. These samples can be divided into roughly two groups, one with a cooling onset age centred around 130 Ma, the other nearer to 100 Ma. As an example, the track age spectrum for sample RWB96-07 (Fig. 5B) clearly indicates the onset Table 3 TASC results for the Southern Zimbabwe Craton Traverse. Sample

Elevation (m)

FT age (Ma)

Mean len. (µm)

Std. dev. (µm)

Cool onset age (Ma)

ρs multiplier

Z96-03 Z96-11 Z96-112 Z96-113 Z96-114 Z96-115 Z96-116 Z96-117 Z96-123 Z96-127 Z96-128 Z96-146

1388 1479 828 818 706 712 482 492 451 691 721 977

116 ± 22 280 ± 22 331.8 ± 20.5 252.2 ± 12.6 211.9 ± 14.7 140.8 ± 15.3 89.0 ± 3.7 108.4 ± 7.7 98.1 ± 6.8 92.1 ± 11.3 201.2 ± 10.5 106.0 ± 7.7

11.67 12 11.36 10.75 11.57 13.34 12.8 12.84 12.24 11.51 10.96 12.64

1.4 1.47 1.84 1.92 2.24 1.65 1.43 1.61 2.11 2.41 1.76 1.94

157 ± 30 379 ± 23 464 ± 42 395 ± 31 290 ± 29 169 ± 20 111 ± 4 134 ± 9 133 ± 8 155 ± 7 293 ± 18 137 ± 5

1.35 1.38 1.42 1.57 1.41 1.2 1.25 1.22 1.32 1.58 1.45 1.36

Mean length and standard deviation refer to confined track length measurements. Cooling onset age refers to calculated event spectra TASC (Ehlers et al., 2005). The track density multiplier (ρs) is the factor by which the original fission track densities are adjusted to account for each track's probability of intersecting the polished surface of a sample. For each sample the probability is calculated using the measured track length data displayed in the traditional length histogram. For example, if the original maximum (formation) track length is ∼16 µm, an 8 µm track has only a 50% chance of intersecting the surface and being sampled. Since we are only observing half of these short tracks, we need to multiply the counted number by 2 to find the actual number present. The amended track count or density is used to generate a measured age of passage through the 110 °C isotherm (ie. the cooling onset age).

of cooling at around the time of Gondwana breakup (∼ 130 Ma), and a younger (90 Ma) cooling phase. The exceptions are 96RWB-12 (340 ± 20 Ma), 96RWB-13 (285 ± 38 Ma) and 96RWB-16 (155 ± 9 Ma). The last of these may simply be an outlier of the Early Cretaceous population. However, the two oldest samples, both of which are located at the town of Platjens near the Botswana border, lie in deformed Archaean Basement that is indistinguishable from most of the other samples. The track age spectra for both of these samples 96RWB12 and 13 are very similar to those recorded in the more cratonic regions of the NMZ suggesting a comparable thermal history. Why they digress from neighbouring samples of the CZ is not clear,

Table 4 TASC results for the Central Limpopo Traverse. Sample

No. of FT age grains (Ma)

Mean len. Std. dev. Cool onset ρs (µm) (µm) age (Ma) multiplier

96RWB-07 96RWB-08 96RWB-09 96RWB-10 96RWB-11 96RWB-12 96RWB-13 96RWB-14 96RWB-15 96RWB-16 96RWB-17 96RWB-19 96RWB-31 96RWB-36 96RWB-37 96RWB-39 96RWB-40a 96RWB-40c 96RWB-41 96RWB-42

20 20 20 20 20 20 10 21 20 13 20 20 20 21 20 20 28 20 20 21

13.73 13.64 13.87 13.73 11.77 11.49 12.39 12.96 13.17 13.66 13.87 13.78 13.3 13.18 12.95 13.33 13.87 13.17 13.1 13.46

105.6 ± 4.0 90.6 ± 3.6 94.2 ± 4.5 83.9 ± 4.0 89.7 ± 4.1 228.0 ± 24.0 213.2 ± 17.8 75.5 ± 3.9 108.0 ± 6.2 129.6 ± 4.9 104.0 ± 6.3 90.5 ± 3.2 93.3 ± 4.8 81.9 ± 3.6 93.4 ± 3.8 104.4 ± 3.6 111.1 ± 5.7 107.2 ± 3.5 96.8 ± 4.0 113.9 ± 8.5

1.7 1.28 1.76 1.39 2.71 2.29 2.67 1.88 2.66 2 1.05 1.31 1.47 1.71 1.53 1.4 1.22 1.36 1.66 1.78

122 ± 4 106 ± 4 108 ± 5 95 ± 4 111 ± 9 340 ± 20 285 ± 38 94 ± 6 127 ± 18 155 ± 9 123 ± 8 104 ± 4 113 ± 5 97 ± 7 116 ± 4 124 ± 4 127 ± 7 137 ± 8 121 ± 5 136 ± 8

1.17 1.17 1.17 1.16 1.24 1.48 1.37 1.25 1.18 1.2 1.18 1.15 1.21 1.19 1.23 1.2 1.14 1.21 1.25 1.19

Mean length and standard deviation refer to confined track length measurements. Cooling onset age refers to calculated event spectra TASC (Ehlers et al., 2005).

D.X. Belton, M.J. Raab / Tectonophysics 480 (2010) 99–108

105

Fig. 5. Traditional fission track length histograms are given on the left with the fission track age together with standard parameters of mean length and standard deviation for the samples track length distribution. The plots on the right have two main components. First, the original histogram (left) has been reproduced — however the length axis (x) has been converted to time (Ma) — reflecting the span of time represented by each of the histograms bins. The bin frequencies are on the right axis (y2). The second component of the plot is a series of points calculated using the Track Age Spectrum Calculation (TASC, Ehlers et al., 2005) showing the time contribution of each bin (with error estimates) to the total age of the sample. The curve is added to guide the eye and highlight any bimodal distributions. The sample (A) 96Z-115 is broadly typical of those found in the NMZ, while (B) 96RWB-07 is more characteristic of the CZ samples. The MonteTrax fit (Gallagher, 1995) for 96RWB-07 is superimposed on the traditional length histogram.

but may indicate the presence of an unidentified fault block that has isolated the samples. The event spectrum for the Central Zone sample suite (Fig. 6B) also very clearly identifies two prominent cooling

events (∼ 121 Ma and ∼ 94 Ma), with most of the samples recording both. As was seen on the NMZ event spectrum, several samples suggest a late Palaeogene cooling.

Fig. 6. The nominal timing of such tectonic events as determined by the TASC approach can be informatively summarised by means of an “event spectrum”. This plot is analogous to the detrital zircon age-probability plots (Brandon, 1996) seen in tectonic provenance studies. It is constructed by extracting the approximate timing of changes in cooling styles from the track age spectra (Fig. 5). Since the timings are semiquantitative and generally imprecise, results from a suite of similar samples are grouped to produce a nominal central value with some indication of dispersion. For the NMZ (Northern Margin Zone) on the left, a suite of 11 samples were used, while for the Limpopo, a group of 20 samples were used to produce the figure. From each suite, a Gaussian curve is plotted using the calculated mean for each “event”, the standard deviation, and the total number of samples from which the same nominal timing has been observed provides the height of the curve. The intention of this plot is to indicate approximate time, with a crude determination of precision of this estimate and how widely this “event” might have been recorded in the overall sample set.

106

D.X. Belton, M.J. Raab / Tectonophysics 480 (2010) 99–108

8. Inverse modelling Inverse modelling of both sample suites was carried out using the MonteTrax program (Gallagher, 1995). The information extracted by TASC analysis was used as a guide in determining both the number and the boundaries of time–temperature envelopes for each t–T point. However, to avoid influencing or forcing the modelling calculations and maximise the degrees of freedom, boundaries on the temperature selection were set to ±60 °C and the freedom on the time constraint was such that each t–T envelope had to partially overlap its neighbouring t–T envelope (Fig. 7). This approach proved very successful at generating rapid convergence to a solution with a high degree of fit for both length and ages parameters. Fig. 7B illustrates a thermal history (96RWB-007) replicated by many of the CZ samples, consisting of a deeply buried sample, cooling rapidly at around 130 Ma, pausing, then accelerating again to lower temperatures at ∼90 Ma. The minor cooling

seen in the Late Palaeogene is also a common feature of samples from the CZ. In contrast, samples from the western (Z96-115) and eastern (Z96-127) reaches of the NMZ (Fig. 7A and C) show a cooling beginning in the Late Jurassic after the cessation of Karoo volcanism. The modelling also suggests the possibility of modest reheating (ie. burial) before cooling again begins around the Late Cretaceous. 9. Discussion 9.1. Early Cretaceous cooling In the Limpopo Valley, Cretaceous cooling of 35 °C–50 °C in the CZ is estimated for the period 130 Ma to 90 Ma and, estimating a geothermal gradient of 26 °C/km, cooling equates to overburden removal of between 1.3 km and 2 km over the 40 My interval. This geothermal gradient is higher than the average geothermal gradient of other cratonic areas (13–18 °C/km) and is suggested on the basis that extensional regimes such as the Karoo Basins are a consequence of increased heatflow that reduces the flexural strength of the lithosphere (Jones, 1998). This, in turn, reduces the effective wavelength over which the lithosphere acts as a rigid plate (Christie-Blick and Biddle, 1985). Here, the bulk of exhumation is limited to regions south of the Umlali Thrust that bounds the NMZ, therefore isolating the central regions of the craton from much of the crustal response. In the Limpopo, there is also evidence for a strike–slip component in what appears to be a predominantly transpressional tectonic environment. While the bulk of the Early Cretaceous structural reactivation was accommodated on the Umlali Thrust, the eastern extension of the Bikita Shear Zone, the Mutare Greenstone Belt was also affected. 9.2. Transpression in the Limpopo region

Fig. 7. Inverse modelling of the fission track data (MonteTrax — Gallagher, 1995) for (A) the NMZ from sample Z96-115 and (B) the CZ from RWB96-07. Both show a prominent cooling in the Early Cretaceous, however almost all the CZ samples show this characteristic two stage cooling (with a second event at ∼90 Ma). (C) Inverse modelling for sample Z96-127, located further east along the NMZ from Z96-115, shows reburial prior to the Early Cretaceous cooling and has an additional prominent Late Cretaceous episode. Dashed boxes show time–temperature constraints for inverse modelling. We used large overlapping time–temperature ranges to maximise degrees of freedom in deriving thermal histories.

Within the Limpopo region there is a combination of linear to curvilinear, typically narrow, sub-vertical displacement zones (eg. Bospokpoort and Tshipise faults). As Christie-Blick and Biddle (1985) noted, this structural style, coupled with the restricted fault azimuth common to synthetic and antithetic shear distributions, suggests that a component of strike–slip motion may have accompanied phases of compression and extension in the Limpopo. Other authors (Cox, 1992; Cox et al., 1965) implied that strike–slip motion in this region might have played a significant role in the breakup of Gondwana. Indeed, a component of strike–slip movement has been interpreted for many Phanerozoic African Rift basins (Lambiase, 1989) and evidence for a similar minor transverse component can be seen in currently active basins such as the 280 Luangwa (Shudowski, 1985). Wilson et al. (1987) associated the emplacement of a major dyke swarm along fracture zones in the Limpopo Belt to a major stress field. De Wit (2007) associates those far-field effects at the end of the Mesozoic to a change and slowing in relative motion of the African and Eurasian plates, resulting in repeated regional uplift coinciding with two punctuated episodes of kimberlite intrusions at ∼120 Ma and ∼90 Ma. Additionally, the mechanism of transpression in a basement environment with faultbounded, rhomboid, sedimentary basins, provides a means of displacing relatively small pieces of crust upwards (or downwards) by magnitudes of up to several kilometers, while adjacent blocks remain static (Fig. 8). This mechanism may explain the preservation of the two, significantly older samples deep within the CZ (96RWB- 12 and -13). This interpretation for the exhumation and cooling of the NMZ and Limpopo regions is in accord with the results of recent gravity surveys and analysis by Gwavava et al. (1996) and Coward and Fairhead (1980). A number of conclusions can be drawn from the isostatic anomaly data of Gwavava et al. (1996). These are 1) the presence of NNE-orientated anomalously low gravity values (ca. −30 mGal to −40 mGal) coinciding with the NMZ in Zimbabwe (in the region of Z96-127). Gwavava et al. (1996) interpreted this to be the result of recent, rapid erosion where the crust has yet to achieve isostatic compensation. This was

D.X. Belton, M.J. Raab / Tectonophysics 480 (2010) 99–108

107

correlated with the proximal Sena Formation (Salman and Abdula, 1995). Further evidence is given by the mid-Cretaceous sedimentary Malonga Formation unconformably overlaying Karoo sediments along the border with Mozambique (Botha and de Wit, 1996). 9.3. Late Cretaceous cooling

Fig. 8. Model of the Central Limpopo Belt illustrating a possible transpressional mechanism for generating both strike–slip motion in small confined blocks and reverse faulting or thrust reactivation on suitably orientated, older slip planes.

consistent with an earlier survey (Gwavava et al., 1992) where a NNEorientated, positive Bouguer anomaly underlying the Northern Marginal Zone was attributed to recent crustal thinning; and 2) a large area of anomalously low response (−20 mGal) within the Central Zone and underlying the bulk of the CZ sample suite. This low anomaly was interpreted to be the result of a zone of thicker (lower-density) uppercrustal rocks producing a lower gravity response. This differed from surrounding areas where high-density lower-crustal rocks may be typically closer to the surface. These general conclusions are supported by evidence from seismic profiles by Stuart and Zengeni (1987), Nguuri et al. (2001), and other geophysical studies (de Beer and Stettler, 1992), that also corroborate the interpretation presented here based on the fission track analysis. Independent evidence from sediment deposition into sinks in Mozambique (Fig. 9) supports the timing of the Early Cretaceous denudation event seen on the Zimbabwe Craton (Dingle et al., 1983; Salman and Abdula, 1995). At this time, the Mozambique Basin records a major influx of coarse, terrigenous sediments, the Sena Formation (Dingle et al., 1983). The bulk of these sediments was delivered to the sub-basins and grabens within the Mozambique Basin, as well as offshore, via the Limpopo and Zambezi Rivers. Within the Maputo and Domo formations, sedimentation rates on the order of 20 m/Ma to 40 m/Ma are seen. These formations are distal facies

During this episode, exhumation, as opposed to uplift as proposed by Molnar and England (1990), is focused across the southern regions of the Zimbabwe Craton. In the Limpopo cooling on the order of 20 °C–26 °C (0.8 km to 1 km of exhumation) is recorded by inverse modelling for the interval 90 Ma to 0 Ma. This younger period appears to have reactivated the Bikita Shear zone in preference to the Umlali Thrust, perhaps as a series of forward propagating imbricate thrusts as in a small scale version of the mechanism outlined by Allegre et al. (1984) to explain shortening in the Himalayas. Indeed, Himalayan style tectonics have previously been called upon to explain much older structural developments in the Limpopo (Treloar et al., 1992). 9.4. Hints of Palaeogene cooling TASC analysis of both data suites has consistently pointed to a Palaeogene cooling (Fig. 6). In general, this event seems to lie on or below the temperature sensitivity for reliable inverse modelling of the fission track data. Indeed, late-stage cooling in thermal histories derived by inverse modelling are often viewed with some caution because of the potential for artifacts as samples are “forced” to cool to their final surface temperature. The TASC analysis faces no such artifact so that a Palaeogene cooling observed in both the inverse modelling results and the TASC analysis is not to be dismissed lightly. Evidence for this cooling is most common in lower elevation samples in vertical profiles and may be attributed to the processes of valley incision and modest scarp retreat. This is more likely given that the samples for vertical profiles are frequently collected during valley descents rather than on ridgelines or the precipitous scarp faces. Further south, De Wit et al. (2000) described a period of river rejuvenation, in the Miocene that significantly altered South Africa's major drainage systems. Detailed analysis of this event lies beyond the current resolution of the apatite fission track technique and awaits more research with thermochronometers sensitive to a lower temperature regime.

Fig. 9. During the Early Cretaceous, the Sena Formation was fed by the major drainage pathways off the Zimbabwe Craton and dominated sedimentation into the Mozambique Basin. Sedimentation rates were in the order of 23 m/My to 44 m/My (Dingle et al., 1983; Salman and Abdula, 1995).

108

D.X. Belton, M.J. Raab / Tectonophysics 480 (2010) 99–108

10. Conclusions While the low magnitude Palaeogene cooling could not be attributed to a particular event in this study, this work demonstrates that major, predominantly late Archaean and early Proterozoic shear zones within the southern part of the Zimbabwe Craton have been reactivated, exposing adjacent rocks to kilometer-scale exhumation in two events during the Cretaceous. The first occurred at around 130 Ma in response to changing plate motions with the opening of the South Atlantic, and the second at around 90 Ma possibly resulting from the collision and obduction of the African plate in Oman. Both episodes also coincide with two intense episodes of kimberlite intrusions in southern Africa (Jelsma et al., 2004). Basin subsidence and sedimentation loads on the Mozambique margin support the timing of these events and provide strong indications of the source and pathways for the eroded material. Acknowledgements This work has been supported by grants from the Australian Research Council (ARC) and the Australian Institute of Nuclear Science and Engineering (AINSE). The authors thank Jan Kramers and Antoine Vernon for their thorough and helpful feedback on this manuscript. References Allegre, C.J., et al., 1984. Structure and evolution of the Himalaya–Tibet orogenic belt. Nature 307 (5946), 17–22. Belton, D.X., Brown, R.W., Kohn, B.P., Fink, D., Farley, K.A., 2004. Quantitative resolution of the debate over antiquity of the central Australian landscape; implications for the tectonic and geomorphic stability of cratonic interiors. Earth and Planetary Science Letters 219 (1–2), 21–34. Botha, G.A., de Wit, M.C.J., 1996. Post-Gondwanan continental sedimentation, Limpopo region, southeastern Africa. Journal of African Earth Sciences 23 (2), 163–187. Brandon, M.T., 1996. Probability density plot for fission track grain-age samples. Radiation Measurement 26, 663–676. Brandt, G., de Wit, M.J., 1997. The Kaapvaal Craton, South Africa. Oxford Monographs on Geology and Geophysics 35, 581–607. Brown, R.W., Summerfield, M.A., Gleadow, A.J.W., 1994. Apatite fission track analysis; its potential for the estimation of denudation rates and implications for models of long-term landscape development. In: Kirkby, M.J. (Ed.), John Wiley & Sons, Chichester. Brown, R.W., Gallagher, K., Gleadow, A.J.W., Summerfield, M.A., 2000. Morphotectonic evolution of the South Atlantic margins of Africa and South America. Christie-Blick, N., Biddle, K.T., 1985. Deformation and sedimentation along strike–slip faults. In: Biddle, K.T., Christie-Blick, N. (Eds.), Strike–Slip Deformation, Basin Formation and Sedimentation: SEPM Special Publications, pp. 1–37. Coward, M.P., Fairhead, J.D., 1980. Gravity and structural evidence for the deep structure of the Limpopo Belt, southern Africa. Tectonophysics 68 (1–2), 31–43. Cox, K.G., 1992. Karoo igneous activity, and the early stages of the break-up of Gondwanaland. Geological Society Special Publications 68, 137–148. Cox, K.G., et al., 1965. The geology of the Nuanetsi igneous province. Philosophical Transactions. Royal Society of London 257A, 76–215. de Beer, J.H., Stettler, E.H., 1992. The deep structure of the Limpopo Belt from geophysical studies. In: de Wit, M.J. (Ed.), Precambrian Research, pp. 173–186. de Wit, M.J., 2007. The Kalahari Epeirogeny and climate change: differentiating cause and effect from core to space. South African Journal of Geology 110, 367–392. de Wit, M.C.J., Marshall, T.R., Partridge, T.C., 2000. Fluvial deposits and drainage evolution. Oxford Monographs on Geology and Geophysics 40, 55–72. Dingle, R.V., Siesser, W.G., Newton, A.R., 1983. Mesozoic and Tertiary geology of Southern Africa. Du Toit, A.L., 1933. Crustal movement as a factor in the geographical evolution of South Africa (presidential address). Ehlers, T.A., et al., 2005. Computational tools for low-temperature thermochronometer interpretation. Low-Temperature Thermochronology: Techniques, Interpretations, and Applications 58, 589–622. Gallagher, K., 1995. Evolving temperature histories from apatite fission-track data. Earth and Planetary Science Letters 136 (3–4), 421–435. Gallagher, K., Brown, R., 1999. Denudation and uplift at passive margins; the record on the Atlantic margin of Southern Africa. Philosophical Transactions — Royal Society. Mathematical, Physical and. Engineering Sciences 357 (1753), 835–859. Gleadow, A.J.W., Belton, D.X., Kohn, B.P., Brown, R.W., 2002. Fission track dating of phosphate minerals and the thermochronology of apatite. Reviews in Mineralogy and Geochemistry 48, 579–630. Gwavava, O., Swain, C.J., Podmore, F., Fairhead, J.D., 1992. Evidence of crustal thinning beneath the Limpopo Belt and Lebombo monocline of Southern Africa based on regional gravity studies and implications for the reconstruction of Gondwana. In: von Frese, R.R.B., Taylor, P.T. (Eds.), Tectonophysics, pp. 1–20.

Gwavava, O., Swain, C.J., Podmore, F., 1996. Mechanisms of isostatic compensation of the Zimbabwe and Kaapvaal cratons, the Limpopo Belt and the Mozambique Basin. Geophysical Journal International 127 (3), 635–650. Jelsma, H.A., et al., 2004. Preferential distribution along transcontinental corridors of kimberlites and related rocks of Southern Africa. South African Journal of Geology 107 (1–2), 301–324. Jones, M.Q.W., 1998. A review of heat flow in southern Africa and the thermal structure of the lithosphere. Southern African Geophysical Review 2, 115–122. Kamber, B.S., Blenkinsop, T.G., Villa, I.M., Dahl, P.S., 1995. Proterozoic transpressive deformation in Northern Marginal Zone, Limpopo Belt, Zimbabwe. Journal of Geology 103 (5), 493–508. King, L.C., 1967. The morphology of the Earth; a study and synthesis of world scenery. King, L., 1976. Planation remnants upon high lands. Zeitschrift fuer Geomorphologie 20 (2), 133–148. Kreissig, K., et al., 2001. Geochronology of the Hout River shear zone and the metamorphism in the Southern marginal zone of the Limpopo Belt, Southern Africa. Precambrian Research 109 (1–2), 145–173 15 Jun 2001. Lambiase, J.J., 1989. The framework of African rifting during the Phanerozoic. Journal of African Earth Sciences 8 (2–4), 183–190. Lister, L., 1979. The geomorphic evolution of Zimbabwe Rhodesia. Verhandelinge van die Geologiese Vereniging van Suid Afrika, Transactions of the Geological Society of South Africa 82 (3), 363–370. Lister, L., 1987. The erosion surfaces of Zimbabwe. Zimbabwe Geological Survey Bulletin 90 163 pp. Marshak, S., van der Pluijm, B.A., Hamburger, M., 1999. The tectonics of continental interiors; preface. In: Hamburger, M. (Ed.), Tectonophysics, pp. vii–x. McCourt, S., Vearncombe, J.R., 1992. Shear zones of the Limpopo Belt and adjacent granitoid–greenstone terranes; implications for late Archaean collision tectonics in Southern Africa. In: de Wit, M.J. (Ed.), Precambrian Research, pp. 553–570. Mkweli, S., Kamber, B., Berger, M., 1995. Westward continuation of the craton-Limpopo Belt tectonic break in Zimbabwe and new age constraints on the timing of the thrusting. Journal of the Geological Society of London 152 (1), 77–83. Molnar, P., England, P.C., 1990. Late Cenozoic uplift of mountain ranges and global climate change; chicken or egg? Nature 346 (6279), 29–34. Moore, A.E., 1999. A reappraisal of epeirogenic flexure axes in Southern Africa. South African Journal of Geology, Suid-Afrikaanse Tydskrif vir Geologie 102 (4), 363–376. Nguuri, T.K., et al., 2001. Crustal structure beneath southern Africa and its implications for the formation and evolution of the Kaapvaal and Zimbabwe cratons. Geophysical Research Letters 28 (13), 2501–2504 01 Jul 2001. Park, R.G., Jaroszewski, W., 1994. Craton tectonics, stress and seismicity. In: Hancock, P.L. (Ed.), Pergamon Press, Tarrytown. Partridge, T.C., 1998. Of diamonds, dinosaurs and diastrophism; 150 million years of landscape evolution in southern Africa. South African Journal of Geology, SuidAfrikaanse Tydskrif vir Geologie 101 (3), 167–184. Partridge, T.C., Maud, R.R., 1987. Geomorphic evolution of Southern Africa since the Mesozoic. Verhandelinge van die Geologiese Vereniging van Suid Afrika = Transactions of the Geological Society of South Africa 90 (2), 179–208. Raab, M.J., Brown, R.W., Gallagher, K., Carter, A., Weber, K., 2002. Late Cretaceous reactivation of major crustal shear zones in northern Namibia: constraints from apatite fission track analysis. Tectonophysics 349 (1–4), 75–92. Raab, M.J., Brown, R.W., Gallagher, K., Weber, K., Gleadow, A.J.W., 2005. Denudational and thermal history of the Early Cretaceous Brandberg and Okenyenya igneous complexes on Namibia's Atlantic passive margin. Tectonics 24 (3). Reenen, v.D.D., Roering, C., Ashwal, L.D., de Wit, M.J., 1992. Special issue; the Archaean Limpopo granulite belt; tectonics and deep crustal processes. Precambrian Research 55 (1–4), 1–587 Mar 1992. Reeves, C.V., de Wit, M.J., Sahu, B.K., 2004. Tight reassembly of Gondwana exposes Phanerozoic shears in Africa as global tectonic players. Gondwana Research 7 (1), 7–19 Jan 2004. Robertson, I.D.M., Du Toit, A.L., 1981. The Limpopo Belt. In: Hunter, D.R. (Ed.), Precambrian of the Southern Hemisphere. Developments in Precambrian Geology. Elsevier, Amsterdam, pp. 641–671. Salman, G., Abdula, I., 1995. Development of the Mozambique and Ruvuma sedimentary basins, offshore Mozambique. Sedimentary Geology 96 (1–2), 7–41. Shudowski, G.N., 1985. Source mechanisms and focal depths of Eas African earthquakes using Rayleigh-wave inversion and body-wavemodelling. Geophysical Journal of the Royal Astronomical Society 83, 563–614. Stagman, J.G., 1978. An outline of the geology of Rhodesia. Stuart, G.W., Zengeni, T.G., 1987. Seismic crustal structure of the Limpopo mobile belt, Zimbabwe. Tectonophysics 144 (4), 323–335. Tinker, J., de Wit, M., Brown, R., 2008. Mesozoic exhumation of the southern Cape, South Africa, quantified using apatite fission-track thermochronology. Tectonophysics 455 (1–4), 77–93 18 Jul 2008. Treloar, P.J., Coward, M.P., Harris, N.B.W., 1992. Himalayan–Tibetan analogies for the evolution of the Zimbabwe Craton and Limpopo Belt. In: de Wit, M.J. (Ed.), Precambrian Research, pp. 571–587. van Reenen, D.D., 1995. Are the Southern and Northern Marginal Zones of the Limpopo Belt related to a single continental collisional event? South African Journal of Geology 98, 498–504. Visser, J.N.J., 1989. The Permo-Carboniferous Dwyka Formation of Southern Africa; deposition by a predominantly subpolar marine ice sheet. Palaeogeography, Palaeoclimatology, Palaeoecology 70 (4), 377–391. White, R.S., 1997. An active mantle mechanism for Gondwana breakup. South African Journal of Geology 100, 271–282. Wilson, J.F., Jones, D.L., Kramers, J.D., 1987. Mafic dyke swarms in Zimbabwe. Special Paper — Geological Association of Canada 34, 433–444.