Neoarchaean Clastic Rocks of the Kaapvaal Craton –

Provenance Analyses and Geotectonic Implications

Eva Anita Schneiderhan Neoarchaean Clastic Rocks on the Kaapvaal Craton – Provenance Analyses and Geotectonic Implications

by

Eva Anita Schneiderhan

THESIS presented in fulfilment of the requirements for the degree of PHILOSOPHIAE DOCTOR in

GEOLOGY

in the

FACULTY OF SCIENCE

of the

UNIVERSITY OF JOHANNESBURG

Supervisor Dr. U. Zimmermann

Co-Supervisor Prof. J. Gutzmer

May 2007

Declaration

I declare that this thesis is my own original work, conducted under the supervision of Dr. Udo Zimmermann and Prof. Jens Gutzmer. It is submitted for the degree Doctor of Philosophy at the Faculty of Science at the University of Johannesburg. No part of this research has been submitted in the past, or is being submitted, for a degree or examination at any other university.

E. A. Schneiderhan

I Acknowledgements

There are a number of people whom I would like to thank, because without them I would not have been able to succeed with this thesis. I appreciate your time and your support that you expressed in various ways, especially during the difficult last months.

First of all, I am thankful to my supervisor Dr. Udo Zimmermann and co-supervisor Prof. Jens Gutzmer for providing me with the topic of this thesis. I am grateful for their support, constructive criticism and discussions, as well as for the bursary and research funding granted.

I am indebted to the German Academic Exchange Service (DAAD) for providing a generous scholarship under grant number D/02/28359 for the first two years.

Avgold Ltd., Mr. Rod Tucker and Mr. Dries Ferreira are thanked for giving access to their exploration drill core and their hospitality in Allanridge. Furthermore, Dr. Herman Dorland is thanked for helping me with the sample collection in the field. In this respect, I would also like to thank Mr. Bill Wheater who gave us access to his Farm Kransberg, from where some of the studied samples originate.

Regarding the logistics and sample analyses, the support by the staff of the Department of Geology, University of Johannesburg, and SPECTRAU is thankfully appreciated.

I greatly appreciate the support by Dr. Richard Armstrong, Australian National University, Canberra, who performed SHRIMP II analyses on a large number of zircons.

Furthermore, I gratefully acknowledge that Prof. Klaus Mezger granted me access to the facilities of the Zentrallaboratorium für Geochronologie of the Westfälische Wilhelms Universität Münster, Germany, to carry out analyses on Sm-Nd and Rb-Sr isotopic systems. Also Miss Heidi Baier and Prof. Erik Scherer are thanked for their support in acquiring a lot of numbers and processing them into useful data. I appreciate their great hospitality during my research visit, as well as fruitful discussion of the data later on.

II Also Prof. Harald Strauss and Mr. Artur Fugman from the Stabile Isotopen Labor of the Geologisch-Paläontologisches Institut of the Westfälische Wilhelms Universität Münster are thanked for taking care of the analyses of stable isotopes of carbonates and organic carbon of selected samples.

I appreciate the support of all those who helped me get going in my first weeks in South Africa.

Thanks to André. You made me a much stronger person, taught me how to catch life’s curveballs, and proved a lot of patience. Jy is my beste maatjie!

Lastly, I am eternally indebted to my parents for their unconditional support during the last years. I appreciate that they let me go - which brought me closer to them again. This thesis is dedicated to my mother, and especially to my father who won’t see the result of my studies anymore. I know that he would have seriously tried to read this thesis entirely, and I appreciate it.

III Abstract

The provenance of the Neoarchaean Ventersdorp Supergroup and several age-related supracrustal successions was analysed to gain insight into the geotectonic evolution of the Kaapvaal Craton during the transition from the Archaean to Proterozoic Eras. The studied successions include, besides the siliciclastic formations of the Ventersdorp Supergroup, the upper Wolkberg and Buffelsfontein Groups, the Godwan Formation and the Schmidtsdrift Subgroup of the basal Transvaal Supergroup in Griqualand West. Petrographic, whole rock geochemical and Sm-Nd isotopic analyses were combined with SHRIMP U-Pb age dating of detrital zircons. Furthermore, Rb-Sr isotopic studies were carried out on carefully selected suites of samples from surface exposure or, wherever possible, on deep diamond drill core.

The Ventersdorp Supergroup is an up to 5 km thick undeformed, only slightly metamorphosed volcano-sedimentary succession deposited on the Kaapvaal Craton between 2714 Ma and 2665 Ma. A lack of major time hiati to the underlying Mesoarchaean Witwatersrand Supergroup and covering Neoarchaean to Palaeoproterozoic Transvaal Supergroup render the Ventersdorp Supergroup very well suited for the investigation of the geotectonic evolution of the Kaapvaal Craton near the Archaean-Proterozoic boundary. This is supported by its excellent preservation, which also allowed detailed studies of sedimentological structures, such as seismites indicating Neoarchaean earthquakes.

The provenance analyses carried out on the clastic formations of the Ventersdorp Supergroup point to a gradual change in tectonic evolution from typically Archaean to post-Archaean processes rather than a drastic, unique transition in the case of the Kaapvaal Craton. Texturally immature wackes of the Kameeldoorns Formation, representing the oldest clastic units of the Ventersdorp Supergroup, are derived mainly from Mesoarchaean source rocks, whereas the stratigraphically younger Bothaville Formation displays geochemical signatures comparable with Archaean trondhjemite-tonalite granodiorite-suites (TTGs), thus suggesting crustal addition in the so-called ‘Archaean-style’.

The extension of provenance analyses to supracrustal successions that are tentatively correlated with the Bothaville Formation, revealed contributions from granitoid

IV sources that formed under post-Archaean and Archaean conditions. Furthermore, the geochemical data for all analysed formations support a passive margin setting. Arc settings, as indicated in some samples, are due to the input of less fractionated volcanic material that provides evidence of distal volcanism. Analyses of Nd-isotopic systematics and U-Pb ages of detrital zircons revealed a Mesoarchaean age for the source rocks of the formations. U-Pb age dating of detrital zircons of the Godwan Formation suggests that this formation is of Mesoarchaean age, and therefore not a correlative of the other Neoarchaean successions.

Hence, the results suggest that the continental crust of the Kaapvaal Craton was thick enough since the Mesoarchaean (2.8 - 3.1 Ga) to allow long-term crustal recycling, and therefore modern plate tectonic processes could have operated earlier than on other well-studied cratonic blocks. During the Neoarchaean, however, crustal thickening of the Kaapvaal Craton took place by accretion of Archaean-style TTGs along the margins of the craton. Thus, Archaean and post-Archaean tectono-magmatic processes co-existed. Furthermore, the Neoarchaean supracrustal successions represent the first sedimentation events on an entirely stabilised and tectonically quiescent Kaapvaal Craton. Input from distal volcanic sources marks the last sign of volcanic activity prior to the craton-wide deposition of carbonate rocks of the Transvaal Supergroup. Geochronological data also imply a connection of the Neoarchaean Kaapvaal Craton to further cratonic blocks that may hold source rocks for the studied formations, as for some small age populations of older detrital zircons (ca. 3.1 - 3.4 Ga), no suitable source area could be identified on the Kaapvaal Craton itself.

However, it seems unlikely that the Zimbabwe Craton was one of these cratonic blocks, because the Rb-Sr whole rock ages of all studied formations yield a model age of 2092 ± 55 Ma, which is thought to correspond to a craton-wide influence of the 2.05 Ga old Bushveld Igneous Complex on the Rb-Sr isotope systematics of all analysed clastic successions. This influence is apparently missing in the Southern and Central Marginal Zones of the Limpopo Belt, suggesting that the collision between the Kaapvaal and Zimbabwe Cratons only took place after the emplacement of the Bushveld Igneous Complex, i.e. after 2.05 Ga.

V Table of Contents

Table of Contents

DECLARATION I ACKNOWLEDGEMENTS II ABSTRACT IV

1 INTRODUCTION 1 1.1 Objective 5 1.2 Analytical Methodology 6 1.3 Plate-tectonic Models and crustal Growth 7 1.3.1 Geotectonic models 8 1.3.1.1 Vertical geotectonic (plume) model 9 1.3.1.2 Horizontal geotectonic (arc) model 10 1.3.1.3 Supercontinents and Vaalbara 10 1.3.2 Archaean and post-Archaean crustal growth models 12 1.3.2.1 Archaean crustal growth 12 1.3.2.1.1 Greenstone Belts 13 1.3.2.1.2 Trondhjemite-Tonalite Granodiorites (TTG) 14 1.3.2.2 Post-Archaean crustal growth 15 1.3.2.3 Geochemical characteristics 15 1.3.3 Summary 16 1.4 References 17

2 GEOLOGICAL SETTING 29 2.1 Previous Work 38 2.2 Ventersdorp Supergroup 41 2.2.1 Lithostratigraphy 41 2.2.2 Palaeoenvironmental Significance 46 2.3 Transvaal proto-basinal Successions 47 2.3.1 Godwan Formation 47 2.3.1.1 Lithostratigraphy 47 2.3.1.2 Palaeoenvironmental significance 48 2.3.2 Wolkberg Group 49 2.3.2.1 Lithostratigraphy 49 2.3.2.2 Palaeoenvironmental Significance 50 2.3.3 Buffelsfontein Group 51 2.3.3.1 Lithostratigraphy 51 2.3.3.2 Palaeoenvironmental Significance 52 2.3.4 Schmidtsdrift Subgroup 52

VI Table of Contents

2.3.4.1 Lithostratigraphy 52 2.3.4.2 Palaeoenvironmental Significance 53 2.4 References 54

3 SEDIMENTARY PROVENANCE OF THE NEOARCHAEAN VENTERSDORP SUPERGORUP, SOUTH AFRICA – SHEDDING LIGHT ON THE EVOLUTION OF THE KAAPVAAL CRATON IN THE NEOARCHAEAN 62 3.1 Abstract 62 3.2 Introduction 63 3.3 Geological Setting 65 3.4 Lithostratigraphy 67 3.5 Samples and analytical Methods 68 3.6 Results 71 3.6.1 Petrography 71 3.6.1.1 Wackes of the KDF 71 3.6.1.2 Wackes of the BVF 74 3.6.2 Whole rock geochemistry 75 3.6.2.1 Major elements 75 3.6.2.2 Trace elements 76 3.6.2.3 Rare earth elements (REE) 76 3.6.3 147Sm/ 144Nd isotopes 83 3.6.4 Zircon geochronology 85 3.7 Discussion 91 3.7.1 Alteration 91 3.7.2 Provenance 94 3.7.3 Implications for the stabilization of the Kaapvaal Craton 105 3.7.4 Implications for plate tectonic processes 107 3.8 Conclusion 109 3.9 References 110

4 THE TRANSVAAL PROTO-BASINS – EMERGENCE OF STABLE SUPRACRUSTAL SUCCESSIONS ACROSS THE NEOARCHAEAN KAAPVAAL CRATON, SOUTH AFRICA 123 4.1 Abstract 123 4.2 Introduction 124

VII Table of Contents

4.3 Geological Setting 126 4.4 Lithostratigraphy 128 4.5 Samples and analytical Methods 131 4.6 Results 135 4.6.1 Petrography 135 4.6.2 Whole Rock Geochemistry 139 4.6.2.1 Major elements 139 4.6.2.2 Large Ion Lithophile Elements (LILE) 139 4.6.2.3 Transition and high field strength elements (HFSE) 140 4.6.2.4 Rare Earth Elements (REE) 141 4.6.3 147Sm/ 144Nd Isotopes 150 4.6.4 Detrital zircon geochronology 152 4.7 Discussion 166 4.7.1 Alteration 165 4.7.2 Provenance 167 4.7.3 Implications for the evolution of the Kaapvaal Craton 178 4.8 Conclusion 181 4.9 References 181 5 Rb-Sr ISOTOPIC SYSTEMATICS OF NEOARCHAEAN SILICICLASTIC ROCKS OF THE KAAPVAAL CRATON – EVIDENCE OF CRATON-WIDE HYDROTHERMAL FLUID FLOW EVENTS ASSOCIATED WITH THE BUSHVELD IGNEOUS COMPLEX 190 5.1 Abstract 190 5.2 Introduction 190 5.3 Geological Setting 193 5.4 Analytical Methods 196 5.5 Results 197 5.6 Discussion and Conclusion 200 5.7 References 201

6 ARCHAEAN SEISMITES OF THE VENTERSDORP SUPERGROUP, SOUTH AFRICA 206 6.1 Abstract 206 6.2 Introduction 206 6.3 Geological Setting 207

VIII Table of Contents

6.4 Petrography 209 6.5 Indications for basinal seismicity 210 6.5.1 Syn-sedimentary faults 211 6.5.2 Convolute lamination 211 6.5.3 Graben-like down-sagging structures 212 6.5.4 Syn-sedimentary breccia 212 6.5.5 Syneresis cracks 213 6.6 Discussion 213 6.7 Conclusion 214 6.8 References 215

7 SUMMARY AND CONCLUSION 218 7.1 Provenance of Neoarchaean siliciclastic successions 218 7.2 Geotectonic Implications for the Kaapvaal Craton 219 7.3 References 225

APPENDIX I: Analytical techniques (on CD) 228

APPENDIX II: Petrographic and geochemical descriptions (on CD) 239

APPENDIX III: Zircon geochronology data (on CD) 381

APPENDIX IV: Stable isotope data (on CD) 400

APPENDIX V: Drill core logs (on CD) 402

IX Chapter 1 Introduction

1 Introduction

The Neoarchaean is an important period of time in Earth history, during which several irreversible and important changes took place. Firstly, the Neoarchaean is the most important phase of continental growth (Windley, 1984). Secondly, the Neoarchaean marks the transition from Palaeoarchaean to modern plate tectonic processes (e.g. Abbott et al., 1994; Kerrich and Polat, 2006; Smithies et al., 2007). Finally, closely related to the changing plate tectonic processes are transformations of the modes of crustal evolution (e.g. Hollings and Wyman, 1999; Davidson and Arculus, 2006).

The Archaean-Proterozoic boundary, in particular, is associated with dramatic changes in the crustal rock record regarding its composition (Taylor and McLennan, 1985; McLennan et al., 2006), as well as the advent of modern plate tectonics (Windley, 1984). Yet, where the transition can be traced in the geological record, it appears gradual instead of ‘dramatic’ (McLennan et al. 1979). Although the Archaean-Proterozoic boundary is arbitrarily placed at 2.5 Ga, the comparison of rock records from different cratons suggests that the changes typically associated with the transition between these two time periods is diachronous, e.g. at 3.2 Ga on the Kaapvaal Craton (e.g. Diener et al., 2005; Moyen et al., 2006), at 3.1 Ga on the Pilbara Craton (Smithies et al., 2007), and at 2.7 - 2.5 Ga in most other Precambrian regions (McLennan, 1989). This obviously implies that some cratons stabilised earlier than others.

The timing of the final change of plate tectonics from Archaean processes to modern ones is still a matter of contention. Hamilton (1998) and Stern (2005) believe that modern plate tectonic processes have only been active since the Meso- and Neoproterozoic, respectively, whilst others believe that they were already operational by ca. 3.3 - 3.2 Ga (Kato et al., 1998, Vearncombe and Kerrich, 1999), or prior to 3.4 Ga as proposed for the Kaapvaal Craton (Dziggel et al., 2002; Diener et al., 2005; Clemens et al., 2006). This debate is linked to controversial opinions regarding the underlying driving mechanisms for Archaean tectonics, such as plumes (e.g. Tomlinson et al., 1999; Condie, 2005) versus wetspots (Smith and Lewis, 1999), versus lateral, lithosphere-controlled temperature variations (Anderson, 1994; Hamilton, 1998). At present, the majority of authors prefers plumes over the above- mentioned alternatives. However, plumes may in fact not be an alternative mechanism

1 Chapter 1 Introduction to modern plate tectonic processes during the Archaean, but rather a complementary one (Davies, 1992). Furthermore, it is widely accepted that the transition from flat to steep subduction, representing the onset of modern plate tectonic processes, was well established at around 2.5 Ga (e.g. Davies, 1992; Abbott et al., 1994; Smithies, 2000; Griffin et al., 2003; Peschler et al., 2004; Kerrich and Polat, 2006). Before 2.5 Ga, plumes and convergent margins with flat subduction may have dominated, with modern processes playing only a subordinate role (Kröner, 1991; Kerrich and Polat, 2006; Smithies et al., 2007; Ernst, 2007).

With a well-preserved and almost complete stratigraphic record from Early Archaean (3.6 Ga) to Late Palaeoproterozoic supracrustal sedimentation, the Kaapvaal Craton of southern Africa is a prime example for the study of the transition from Archaean to post-Archaean tectonic processes and associated sedimentary basin evolution.

The present study centres on the Neoarchaean (2.8 - 2.65 Ga) Ventersdorp Supergroup (Fig. 1.1), widely recognized as one of the least deformed and metamorphosed (lower greenschist facies, Cornell, 1978; Tyler 1979 a) volcano- sedimentary successions of Neoarchaean age worldwide, which even allows detailed sedimentological studies (Chapter 6). The Ventersdorp Supergroup formed in response to craton-wide tectonic processes affecting the Neoarchaean Kaapvaal Craton. No major time gap exists between the deposition of the Ventersdorp Contact Reef (VCR) of the Witwatersrand Supergroup and the eruption of the Klipriviersberg (KRB) flood basalts at the base of the Ventersdorp Supergroup (Hall and Els, 2002). There appears to be also no major time gap between the top of the Ventersdorp Supergroup and the overlying Neoarchaean to Palaeoproterozoic (2.68 - 2.1 Ga) Transvaal Supergroup, i.e. there is a virtually complete record of volcanism and sedimentation across the Archaean-Proterozoic boundary. The Ventersdorp Supergroup is comprised of volcanic rocks associated with volcanically-derived and siliciclastic sedimentary rocks that are expected to reveal clues to igneous activity and the possible style of plate tectonic processes. Therefore, a conclusive provenance study of clastic rocks of the Ventersdorp Supergroup provides a rare opportunity to gain insight in the crustal composition and hence the evolution of the Kaapvaal Craton in the Neoarchaean.

2 Chapter 1 Introduction

Fig. 1.1: Geological map indicating the exposed areas of the Ventersdorp and Transvaal Supergroups, and the pre-Transvaal successions (modified after Eriksson et al., 2002; Sumner and Beukes, 2006). Furthermore, the sampling locations are shown. KFN1: exploration drill core that comprises the entire Ventersdorp Supergroup; GWF: Godwan Formation; STF: Selati Formation; TKF: Tygerkloof Formation; BT-VBF: Vryburg Formation near Bothithong; PK-VBF: Vryburg Formation near Prieska. Models for the tectonic setting of the Ventersdorp Supergroup have been proposed by numerous authors who studied the volcanic rocks of the Ventersdorp Supergroup. Some support modern subduction-related rifting (Burke et al., 1985; Stanistreet et al., 1986; Clendenin et al., 1988; Crow and Condie, 1988; Silver et al., 2006) or an Atlantic-type continental margin (Tinker et al., 2002), others prefer an Archaean plume-related setting (e.g. Marsh et al., 1992; Nelson et al., 1992; Hatton, 1995; White, 1997; Arndt, 1999; Eriksson et al., 2002; Ernst and Buchan, 2002; Barley et al., 2005; Ernst et al., 2005) or even compare the Ventersdorp Supergroup to greenstone belts, an interpretation that is based on the geochemistry of volcanic rocks of the Ventersdorp Supergroup (Schweitzer and Kröner, 1985; Cheney et al., 1990; McIver et al., 1982; Eriksson et al., 2001).

The maximum stratigraphic thickness of the Ventersdorp Supergroup is approximately 5100 m, half of which is composed of sedimentary rock units, including volcanogenic and siliciclastic conglomerates, wackes, reworked tuff beds,

3 Chapter 1 Introduction and minor carbonate rocks (van der Westhuizen and de Bruiyn, 2000). Sedimentary deposits are intercalated with thick successions of volcanic rocks of bimodal composition (Fig. 1.2). The base is composed of mafic to ultra-mafic volcanic rocks of the KRB Group, followed by conglomerates, wackes and stromatolitic carbonates of the Kameeldoorns Formation (KDF), and rhyolitic to intermediate lavas of the Makwassie and Rietgat Formations, which belong to the Platberg Group. The top is composed of conglomerates and wackes of the Bothaville Formation (BVF), and mafic volcanic rocks of the Allanridge Formation, which together constitute the Pniel Group (Winter, 1976).

Age-equivalent to the Ventersdorp Supergroup are several supracrustal successions of limited lateral distribution on the Kaapvaal Craton. The successions investigated in this study include the Godwan Formation, the Wolkberg and Buffelsfontein Groups, and the Vryburg Formation of the lower Transvaal Supergroup (Figs. 1.1, 1.2). Based on sequence stratigraphic analyses, the Pniel Group of the Ventersdorp Supergroup has been correlated with the upper Wolkberg Group, the upper Buffelsfontein Group and the upper Godwan Formation (Fig. 1.2; Cheney et al., 1990; Cheney and Winter, 1995). This is supported by findings of Tyler (1979 b), who used stratigraphic and geochemical analyses in combination with analyses of depositional environments to propose correlation between the Wolkberg and Buffelsfontein Groups. Beukes et al. (2002) and Dorland (2004) suggest the correlation of the Vryburg Formation of the Schmidtsdrift Subgroup with the Wolkberg and Buffelsfontein Groups. Geochronological data to support such correlation is presented by Dorland (2004). Barton et al. (1995) suggested an age relationship between the upper Buffelsfontein Group and Pniel Group by using whole rock Rb-Sr isotopic systematics. The relevance of these Rb-Sr data may be called into question (Chapter 5). The least agreement exists with regards to the correlation of the Godwan Formation, since Button (1978) related it to the upper Witwatersrand Supergroup, and Myers (1990) presented results that correlated the lavas of the Godwan Formation with the Nsuze Group of the Pongola Supergroup. In addition, the stratigraphic position of the BVF below Allanridge lavas renders its correlation to the VBF, which is located on top of the Allanridge lavas, questionable.

4 Chapter 1 Introduction

Fig. 1.2: Schematic lithostratigraphy and tentative correlation of the studied successions. The sampled formations are in bold italic; depositional ages are indicated. Fm: Formation. Schmidtsdrift Subgroup after Beukes (1979); Buffelsfontein Group after Barton et al. (1995); Wolkberg Group after Dorland (2004); Godwan Formation after Myers (1990). 1: U-Pb zircon ages, Gutzmer and Beukes (1998); 2: U-Pb zircon age, Martin et al. (1998); 3: U-Pb zircon ages, Armstrong et al. (1991); 4: U-Pb zircon age, Barton et al. (1995); 5: ID-TIMS zircon age, Anhaeusser and Poujol (2004); 6: ID-TIMS zircon age, Kamo and Davis (1994).

1.1 Objective

The major objectives of this study are to provide a detailed provenance analysis of the two main sedimentary formations (BVF and KDF) of the Ventersdorp Supergroup, and the correlation of these to other age-related supracrustal successions of the Kaapvaal Craton. The results obtained yield a more detailed understanding of the relevant plate-tectonic processes involved during the stabilization, i.e. crustal thickening, of the Kaapvaal Craton. It is suggested that the volcanism and sedimentation associated with the Ventersdorp Supergroup reflect the transition from a tectonically unstable to stable Kaapvaal Craton.

In addition, a better understanding of the relationship of the Bothaville Formation to its age-equivalents is obtained by examining their litho- and sequence-stratigraphy, geochronology, provenance, and geochemical characteristics (e.g. Tyler, 1979 b; Cheney et al., 1990; Barton et al., 1995). Finally, the data provide evidence for the processes that influenced the growth, stabilization and evolution of the Kaapvaal

5 Chapter 1 Introduction

Craton during the Neoarchaean. The study illustrates that the correlation of supracrustal successions provides an important means to understanding the tectonic evolution of an Archaean craton (Taylor and McLennan, 1985; McLennan et al., 1990, 1993) and it is expected that this approach may be applicable to other cratons.

1.2 Analytical Methodology

Provenance analyses aim at models that outline the character of the source rocks of the studied formations, as well as their palaeotectonic setting. In this study, a combination of several methodological approaches was chosen to obtain a meaningful, comprehensive provenance analysis of all examined sedimentary formations (Appendix I).

Samples were collected from suitable drill cores and surface exposures. Prior to sampling, the drill cores were logged and profiles of outcrops were drawn to illustrate the lithological content and understand the stratigraphic context of the different formations. Drill core samples chosen were at least 20 cm long, to assure representation. In outcrops, the samples were selected in such a way that after removing weathering rims, they were large enough to mill a fist-sized portion, make a thin section, and keep a hand specimen.

For petrographical studies, thin sections were prepared and fist-sized portions of the samples were milled to analytical fineness (< 25 µm). The thin sections were investigated under transmitted and polarised light to establish mineralogical composition and textural characteristics (results in Appendix II). Sample powders were used to determine the mineralogical composition using X-ray powder diffractometry, which is especially important for lithologies that are too fine-grained for light microscopic studies. Such fine-grained lithologies were also examined by scanning electron microscopy.

Sample powders were also used for geochemical whole rock analyses. Most major and trace element concentrations were determined using X-ray fluorescence spectrometry. Selected trace elements and rare earth element concentrations were analysed using either instrumental neutron activation analysis (INAA) or inductively coupled plasma mass-spectrometry (ICP-MS) analysis. The latter data were supplied by ACME Analytical Laboratories Ltd., Vancouver, Canada (results in Appendix II). Reproducibility of the XRF analyses was established at less than 5 % deviation

6 Chapter 1 Introduction through repeated measurements, while the accuracy deviated between 0.39 % and 11.44 % as established through measurements of standard JR-1, which presented the closest geochemical composition compared to the analysed samples.

For the 147Sm/ 144Nd isotopic ratio and 87Rb/ 86Sr isotopic ratios, sample powder aliquots were prepared and analysed at the Centre for Geochronology of the Westfälische Wilhelms Universität, Münster, Germany.

Lastly, detrital zircons were separated and then analysed by Dr. Richard Armstrong using the sensitive high resolution ion microprobe (SHRIMP II) at the Australian National University, Canberra (results in Appendix III).

Appendix IV presents results of carbon and oxygen stable isotope analyses of selected stromatolitic carbonate samples.

1.3 Plate-tectonic models and crustal growth

The presence or importance of plate tectonic processes driven by mantle convection, as they are operating in the Phanerozoic, is controversial for the Archaean (Kerrich and Polat, 2006). Alternative mechanisms proposed for the Early Archaean include vertical non-plate tectonic processes, such as foundering and up-welling of basalts initiated by thermochemical mantle convection (van Thienen et al., 2004), delamination (Zegers and van Keken, 2001; Bédard, 2006) and diapirism (Hamilton, 1998), or flat and shallow subduction with a change to modern (horizontal) plate tectonics at 2.5 Ga (Abbott et al., 1994). As underlying driving mechanisms for vertical none-plate tectonic processes, plumes (e.g. Tomlinson et al., 1999; Condie, 2005), wetspots (Smith and Lewis, 1999) and lateral, lithosphere-controlled temperature variations (Anderson, 1994; Hamilton, 1998) are being discussed, of which the plume mechanism is favoured by most authors. However, plumes are not an alternative mechanism to modern plate tectonic processes, but rather a complementary one, especially with regards to crustal evolution (Davies, 1992; Rollinson, 2006). According to Davies (1992), in early Earth history plate tectonic processes operated, but at a much slower rate, such that a complementary mechanism was necessary to achieve the required rate of heat loss from the Earth’s interior.

It is widely accepted that the thermal evolution of the Earth played a crucial role in the development and activity of plate tectonic and none-plate tectonic processes. On

7 Chapter 1 Introduction one side, it is believed that heat loss was encouraged by plate tectonic processes involving a thin, komatiitic oceanic crust (Nisbet and Fowler, 1983; Davies, 2006). On the other side, Hamilton (1998) suggested that voluminous magmatism without plate tectonic processes was the primary heat loss mechanism in the time period between 3.6 Ga and 2.6 Ga. Thus, it is pivotal to understand, which mechanisms prevailed, to shed light not only on the Earth’s crustal evolution, but also on its thermal evolution during the Early Archaean.

In addition to the process itself, also the timing of the transition from tectonic processes limited to Early Earth history to modern plate tectonics is a matter of contention. Hamilton (1998) and Stern (2005) believe that modern plate tectonic processes have only been active since the Meso- and Neoproterozoic, respectively, whilst others believe that they were already functional by ca 3.3 - 3.2 Ga (Kato et al., 1998, Vearncombe and Kerrich, 1999). Furthermore, it is widely accepted that the transition from flat to steep subduction, representing the onset of modern plate processes as dominant control on plate tectonics, was well established at around 2.5 Ga (e.g. Davies, 1992; Abbott et al., 1994; Smithies, 2000; Griffin et al., 2003; Peschler et al., 2004; Kerrich and Polat, 2006). Before 2.5 Ga, plumes and convergent margins with flat subduction may have dominated, with modern plate tectonic processes playing only a subordinate role (Kröner, 1991; Kerrich and Polat, 2006; Ernst, 2007; Smithies et al., 2006). However, the Archaean-Proterozoic boundary is accepted by many authors as the time at which modern plate tectonic processes were fully functional (Windley, 1984; Moores, 2002).

1.3.1 Geotectonic models

Different models have been proposed for the evolution and dynamics of the Earth’s crust during the time prior to 2.5 Ga. The major difference between Archaean and post-Archaean processes lies in the fact that Archaean non-plate tectonic processes may have involved mostly vertical tectonic movements, whilst modern post-Archaean plate tectonic processes are directed mostly horizontally.

Geochemical studies of Archaean greenstone belts identified convergent margins, involving oceanic crust, continental and oceanic island arcs, and smaller continental blocks, as possible plate tectonic settings for the Archaean (Kerrich and Polat, 2006). Non-plate tectonic plume settings and delamination, which are related to the Earth’s

8 Chapter 1 Introduction hotter thermal state in the Early Archaean, are commonly cited alternatives (e.g. Davies, 1992; Zegers and van Keken, 2001; Tomlinson and Condie, 2001).

Plumes and plate tectonics operate largely independently from each other, and are complementary modes of mantle convection (Davies, 1992; Hill et al., 1992). Mantle plumes arise from the hot thermal boundary layer at the base of the Earth’s mantle (bottom-up), whilst plates are part of the cool thermal boundary layer at the top of the convecting mantle (top-down) (Hill et al., 1992). Driving forces are different, namely heat loss from the mantle for the plates, versus heat loss from the core to trigger plume upwelling (Hill et al., 1992).

1.3.1.1 Vertical geotectonic (plume) model

Vertical tectonic processes are bottom-up controls on mantle convection, such as ascending plumes and mantle overturn events, which are linked to higher temperatures of the Earth’s Interior (Ernst, 2007). During the Archaean, mantle plumes may have been much more frequent with stronger intensities and larger volumes than in the post-Archaean (Tomlinson and Condie, 2001; Abbott and Isley, 2002). However, it appears as if no difference existed between the duration of Archaean and post-Archaean superplume events (Abbott and Isley, 2002).

Plumes are related to the formation of large igneous provinces, whereas superplumes are often named as the cause for the breakup of supercontinents (e.g. Dalziel et al., 2000). Upwelling asthenosphere or a plume causes extension and lithospheric thinning associated with rifting (Bott, 1995), which can culminate in continental breakup. Of particular importance to this study is one important superplume event that produced large igneous provinces at ca. 2.7 Ga, and that is either related to the breakup of the Neoarchaean supercontinent Kenorland or to a global catastrophic event (Nelson, 1998; Eriksson et al., 2002; Ernst and Buchan, 2002; Condie, 2004; Ernst et al., 2005; Kerrich and Polat, 2006).

Opposing the plume model, Anderson (1994) stated that the presence of hot mantle or massive volcanism is not sufficient to prove the existence of a deep mantle plume. According to this author, the plume hypothesis is only one of several possibilities for regional variations in mantle chemistry and temperature related to continental breakup.

9 Chapter 1 Introduction

In Phanerozoic settings, plumes often exist below back-arc regions above subducting oceanic lithosphere (Bott, 1995). Most Precambrian arc-like rocks are thought by some to be the result of hotspots and mantle plumes instead of the results of island arc processes (Abbott, 1996). According to studies by Tomlinson and Condie (2001), approximately 35 % of Neoarchaean (3.0 - 2.5 Ga) greenstones and ca. 80 % of Palaeoarchaean (3.0 Ga) greenstones may have geochemical plume affinities, which express themselves in ‘pseudo-subduction signatures’ such as Ti, Ta and Nb depletion and Th and Pb enrichment (Hofmann, 1988).

1.3.1.2 Horizontal geotectonic (arc) model

Horizontal tectonic processes, i.e. modern plate tectonics, are top-down controls of convection, where cool, buoyant oceanic crust dominates the convection of the less hot mantle (Ernst, 2007). Some authors have traced back shallow (non-uniformitarian) subduction and the generation of convergent margins into the Mesoarchaean (Kerrich and Polat, 2006; Smithies et al., 2007). However, modern plate tectonics started off with subduction along isolated patches, and formed linear belts only later in the Mesoarchaean (Ricou, 2004; Smithies et al., 2007). Structures related to modern plate tectonic processes such as thrusting, imbrication, strike-slip faulting, continental rifting and orogenic collapse were recognised in greenstone belts of different Archaean cratons (Chen et al., 2006; Kerrich and Polat, 2006).

1.3.1.3 Supercontinents and Vaalbara

Several times in Earth history, all continents amalgamated to one supercontinent. The most prominent ones are Rodinia (1100 - 1000 Ma), Gondwana (900 - 500 Ma) and Pangaea (450 - 250 Ma) (Condie, 2002). Another supercontinent called Kenorland (Williams et al., 1991) formed most probably in the Neoarchaean at 2.7 Ga (Heaman, 1997; Condie, 1998; Condie, 2000; Kerrich and Polat, 2006). The configuration and existence of Kenorland are not undisputed, as some authors prefer a model in which several supracratons co-existed instead of one supercontinent (Aspler and Chiarenzelli, 1998; Bleeker, 2003). In contrast to a supercontinent that comprises all continents combined into one, a supracraton is composed of only a few continents that collide with each other. Aspler and Chiarenzelli (1998) limit the amount of supracratons to two, namely Kenorland and Zimvaalbara. The latter is thought to be

10 Chapter 1 Introduction composed of the Zimbabwe, Kaapvaal and Pilbara Cratons, which started to break up before all continental fragments could amalgamate to one supercontinent.

However, the existence of such a supercontinent is controversial. Nelson et al. (1992) and Nelson et al. (1999) relate the presence of 2.7 Ga old large igneous provinces to a global event, and not to the breakup of a supercontinent. Yet, the breakup of a supercontinent, or supracraton, might be an explanation for the plume activity that can play a major part in supercontinental breakup (e.g. see the example of Pangea; Dalziel et al., 2000).

The correlation of the Kaapvaal and Pilbara Cratons to form the Neoarchaean Vaalbara (Cheney, 1996) supports the existence of at least one supracraton at 2.7 Ga. Correlation is based on the excellent comparison of the lithostratigraphy (e.g. Grobler et al., 1989; Thorne and Trendall, 2001), and geochronology (Martin et al., 1998). Only few authors doubt the common history of the Kaapvaal and Pilbara Cratons during the Late Archaean (Nelson et al., 1992; Nelson et al., 1998; Wingate, 1998).

The excellent correlation of at least three major lithostratigraphic successions supports the existence of Vaalbara. These are firstly the Ventersdorp Supergroup and the Transvaal proto-basin of the Kaapvaal Craton with the Fortescue Group of the Pilbara Craton; secondly the Black Reef Formation and the Chuniespoort and Ghaap Groups of the Kaapvaal Craton with the Jeerinah Formation and the Hamersley Group of the Pilbara Craton; thirdly the lower Pretoria and Postmasburg Groups of the Kaapvaal Craton with the Turee Creek Group of the Pilbara Craton (Trendall et al., 1990; Arndt et al., 1991; Eriksson et al., 1993, 1995; Eriksson and Reczko, 1995; Walraven and Martini, 1995; Aspler and Chiarenzelli, 1998).

Close similarities of the Ventersdorp Supergroup and the Fortescue Group of the Pilbara Craton (Western Australia) have been established in numerous studies (Grobler et al., 1989; Arndt, 1999; Thorne and Trendall, 2001; Eriksson et al., 2002). They include the emplacement mechanism of the lavas, described as subaerial suites of continental flood basalts (Arndt, 1999), as well as similar lithologies and geochemical affinities (Grobler et al., 1989).

11 Chapter 1 Introduction

1.3.2 Archaean and post-Archaean crustal growth models

Different thermal state and heat transfer mechanisms resulted in different tectonic processes dominating the Archaean and post-Archaean time, respectively (Hamilton, 1998). These gave rise to different crustal growth processes. Since the onset of modern plate tectonic processes, crustal growth is dominated by island arc volcanism at convergent plate margins (Davidson and Arculus, 2006; McLennan et al., 2006), while Archaean crustal growth appears to be mainly, but not exclusively, controlled by vertical means of heat transfer (e.g. Hollings and Wyman, 1999; Tomlinson et al., 1999; Ernst, 2007). The most important phase of continental growth seems to have happened in the Neoarchaean (Fig. 1.3) (Windley, 1984).

Fig. 1.3: Diagram presenting a crustal growth curve versus time in relation to crust forming processes. 1: Ernst (2007); 2: Tomlinson et al. (1999); 3: Drummond and Defant (1990); 4: Hollings and Wyman (1999); 5: de Wit and Ashwal (1997); 6; e.g. Peschler et al. (2004); 7: Rollinson (2006). TTG: trondhjemite-tonalite granodiorites; GSB: greenstone belt.

1.3.2.1 Archaean crustal growth

The Earth’s thermal evolution is closely linked to plate tectonic processes and the stabilization of the early Earth’s crust (Nisbet and Fowler, 1983; Zegers and van Keken, 2001). Some authors believed that due to elevated mantle temperatures during the Eoarchaean Earth, asthenospheric circulation was so dynamic that continental crust could only be preserved by ca. 3.8 Ga when asthenospheric circulation had slowed down sufficiently (Ernst, 2007 and references therein). Even during the Mesoarchaean, radiogenic heat production was three times higher than at present day

12 Chapter 1 Introduction as discussed by Zegers and van Keken (2001). Thus, a hotter mantle allowed pressure-released melting at greater depths, so that thick basaltic crust could form on a stratified harzburgite residue (Zegers and van Keken, 2001). The resulting strong compositional stratification suppressed gravitational instabilities that are necessary for modern-day subduction (Zegers and van Keken, 2001). As a consequence, Zegers and van Keken (2001) came to the conclusion that early crust stabilised through delamination of an eclogitic layer at the base of mafic proto-crust, which led to rapid uplift, extension and voluminous magmatism that produced trondhjemite-tonalite granodiorites (TTGs). This model is in contrast to the suggestion that the genesis of Archaean TTGs is associated with subduction zones (e.g. Drummond and Defant, 1990; Hollings and Wyman, 1999).

The addition of material to the continental crust in an episodic or semicontinuous fashion reached its maximum in the Neorchaean and Palaeoproterozoic, so that by then ca. 70 % of the continental crust was generated (Fig.1.3) (Ernst, 2007 and references therein; McLennan et al., 2006). Since then, the production of continental crust has declined (Ernst, 2007). Towards the end of the Archaean, thickened continental cratons and emerging landmasses appeared more frequently, which resulted in increased importance of erosion, sedimentation and alteration or weathering (Ernst, 2007). At the same time, modern plate tectonic processes are thought to dominate (e.g. Davies, 1992; Kerrich and Polat, 2006).

The Archaean crustal basement that evolved through plume and delamination processes (Tomlinson et al., 1999; Zegers and van Keken, 2001) is therefore thought to be composed of two major components, which are greenstone belts and TTGs. The geochemical composition of both components is unique to the Archaean and has no equivalent post-Archaean counterpart. Both greenstone belts and TTGs are thought to reflect the unique heat flow and tectonic processes that dominated the Earth’s evolution until the Mesoarchaean.

1.3.2.1.1 Greenstone Belts

The lithostratigraphy of Archaean greenstone belts can be broadly subdivided into three units. Tholeiitic basalts and komatiites constitute the basal unit, followed by tholeiitic to calc-alkaline bimodal basalts and rhyolites (Lowe, 1994; Polat and Kerrich, 2001), and compositionally mature sedimentary rocks at the top. The

13 Chapter 1 Introduction volcanic rocks that dominate the lower stratigraphic units are interpreted as oceanic plateaus derived from mantle plumes (Abbott, 1996; Polat and Kerrich, 2001; Tomlinson and Condie, 2001). These are covered by immature sedimentary formations, which may be associated with arc-trench systems (Polat and Kerrich, 2001). The sedimentary successions are intruded by syn- to post-kinematic TTG plutons (Polat and Kerrich, 2001). More mature sedimentary rocks, which characterise the youngest portion of greenstone belt successions, indicate stabilisation of the crust and the onset of intracrustal recycling (Stevenson, 1995). However, according to Fralick and Miall (1989), sedimentation of more mature sedimentary rocks on the Superior Craton began only in the Palaeoproterozoic, which underlines the diachronous character development of the various cratons. Accretion of single greenstone belts onto older continental blocks, together with underplating and the formation of anatectic granitoids, resulted in crustal growth and stabilization (Lowe, 1994; Kusky and Polat, 1999).

Deformation in association with plutonism is a typical feature of greenstone belts (de Wit and Ashwal, 1997 and references therein), and gives evidence of two different plate tectonic regimes. The first one can be observed in Palaeoarchaean greenstone belts, where vertical diapiric deformation related to TTG plutonism affected these belts. The second one, observed in Neoarchaean greenstone belts, is characterised by deformation through horizontal compression as in modern plate tectonic processes (de Wit, 1998; Hickmann, 2004; Peschler et al., 2004; Chen et al., 2006). Hence, Neoarchaean greenstone belts resemble rock successions formed in arc-trench settings and subduction zones (Langford and Morin, 1976; Polat and Kerrich, 2001).

1.3.2.1.2 Trondhjemite-Tonalite Granodiorites (TTG)

Archaean TTGs are sodium-rich, high-Al2O3 granitoids that are associated with greenstone belts (Drummond and Defant, 1990). Geochemically, they are characterised by depletion of heavy rare earth elements and negative Nb-Ta and Ti anomalies (Drummond and Defant, 1990). Archaean gneiss terranes comprise more than 80 % TTGs, which also contribute fundamentally to the generation of the present mass of continental crust (Drummond and Defant, 1990). The large volume of TTGs produced in the Archaean is associated with the elevated thermal state of the Palaeoarchaean Earth (Drummond and Defant, 1990).

14 Chapter 1 Introduction

However, TTGs older than 3.0 Ga are not related to modern subduction processes (Smithies, 2000; Condie, 2005), unlike their Phanerozoic counterparts, the so-called adakites (Drummond and Defant, 1990). Instead, Archaean TTGs are believed to be generated by delamination and melting of hydrous basaltic material (sub-continental lithospheric mantle) at the base of the thickened crust that formed above plumes (Smithies, 2000; Zegers and van Keken, 2001; Condie, 2005; Bédard, 2006). Their geochemical characteristics imply that they must have been derived from mafic sources with low Nb/Ta ratios (Xiong et al., 2005). In addition to this, the emplacement of TTGs is interpreted to be linked to major mantle plume events, e.g. at 2.7 Ga, which may be the consequence of catastrophic mantle overturn (Condie, 2005).

1.3.2.2 Post-Archaean crustal growth

Since the domination of modern plate tectonic processes, crustal growth mainly takes place at convergent plate margins, and to a smaller extent at within-plate and rifted margin settings (Rollinson, 2006). At convergent margins, dehydration processes in the subducting slab produce basaltic mantle melt, which fractionates in magma chambers to andesitic composition, and which leads to thickening of the continental crust through volcanism and plutonism (Rollinson, 2006). Eventually, collision and accretion of oceanic island arcs to one another and to other continental crust lead to crustal growth (Rollinson, 2006).

The extrusion of flood basalts, as well as the formation of oceanic plateaus on oceanic crust in within-plate settings, contribute to a minor extent to the growth of the crust through crustal thickening (Rollinson, 2006). Finally, at rifted margins, the extrusion of basalts adds to the crust through accretion along the rifted margins, however not to a large extent (Rollinson, 2006).

1.3.2.3 Geochemical characteristics

Archaean and post-Archaean upper continental crust (UCC) are distinctly different in their geochemical compositions (McLennan and Taylor, 1984; Taylor and McLennan, 1985; McLennan et al., 2006). According to McLennan and Taylor (1984), early Archaean rocks are characterised by a wider range of element concentrations and ratios, and have lower abundances in trace and rare earth elements in general, as compared to post-Archaean UCC, although certain compatible elements can be

15 Chapter 1 Introduction enriched, such as Cr and Ni (Table 1.1). Ratios that are different in the Archaean and post-Archaean UCC are for example LaN/YbN, GdN/YbN and Eu/Eu*, and differences exist in trace element concentrations, such as Th, U, and Nb and Ta (Table 1.1) and compatible elements.

These changes in the overall geochemical composition of UCC are related to the increasing importance of intracrustal differentiation since the Neoarchaean, which entails the evolution of a more felsic crustal composition with enrichment in incompatible trace elements (McLennan et al., 2006).

Table 1.1 Selection of important trace element ratios and concentrations that differ distinctly for Archaean and post-Archaean UCC. Archaean UCC post-Archaean UCC

LaN/YbN 6.8 9.2

GdN/YbN > 2.0 1.4 Eu/Eu* > 0.85 0.42 Sm/Nd 0.20 0.17 Th/Sc 0.43 0.79 La/Sc 1.30 2.21 Cr 180 83 Ni 105 44 Nb 4.58 12.0 Ta 0.34 1.0 Th 5.7 10.7 U 1.5 2.8 Trace element concentrations are reported in ppm. Data from McLennan et al. (2006).

1.3.3 Summary

The change from Palaeoarchaean vertical tectonic processes to modern horizontal plate tectonic processes that dominate since the Neoarchaean is related to the changing thermal state of the Earth. This gives way to a decrease in the importance of plume activity, which was the main initiator for crustal growth during the Palaeo- and Mesoarchaean, when most of the crust was formed through amalgamation of greenstone belts to proto-continents and intrusions of TTGs. Since modern plate tectonics became dominant in the Neoarchaean, most of the post-Archaean crustal growth occurs at convergent plate margins.

16 Chapter 1 Introduction

Crustal stabilization and the change in tectonic processes resulted in a change of the UCC composition, as the post-Archaean UCC is characterised by intracrustal fractionation compared to the less differentiated Archaean UCC.

A particular feature of the Neoarchaean is the formation of a supercontinent that either comprised all existent cratonal blocks or consisted of at least two supracratons, of which one may have been represented by the amalgamated Kaapvaal and Pilbara Cratons.

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Martin D. M., Clendenin C. W., Krapez B. and McNaughton N. J. (1998) Tectonic and geochronological constraints on late Archean and Paleoproterozoic stratigraphy correlation within and between the Kaapvaal and Pilbara Cratons. Journal of the Geological Society, London 155, 311 - 322.

McIver J. R., Cawthorn R.G. and Wyatt B.A. (1982) The Ventersdorp Supergroup - the youngest komatiitic sequence in South Africa. In: Arndt N. T. and Nisbet E. G. (eds.) Komatiites. George Allen & Unwin, London 81 - 90.

McLennan S. M. (1989) Rare earth elements in sedimentary rocks: influence of provenance and sedimentary processes. In: Lipin B. R. and McKay G. A. (eds.) Geochemistry and mineralogy of rare earth elements. Mineralogical Society of America 21, 169 - 200.

McLennan S. M. and Taylor S. R. (1984) Archaean Sedimentary Rocks and Their Relation to the Composition of the Archaean Continental Crust. In: Kröner A.,

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Hanson G. N. and Goodwin A. M. (eds.) Archaean Geochemistry. Springer- Verlag, Berlin Heidelberg, 47 - 72.

McLennan S. M., Fryer B. J. and Young G. M. (1979) Rare earth elements in Huronian (lower Proterozoic) sedimentary rocks; composition and evolution of the post-Kenoran upper crust. Geochimica et Cosmochimica Acta 43, 375 - 388.

McLennan S. M., Taylor S. R. and Hemming S. R. (2006) Composition, differentiation, and evolution of continental crust; constraints from sedimentary rocks and heat flow. In: Brown M. and Rushmer T. (eds.) Evolution and differentiation of the continental crust. Cambridge University Press, New York, NY, United States, 92 - 134.

McLennan S. M., Taylor S. R., McCulloch M. T. and Maynard J. B. (1990) Geochemical and Nd-Sr isotopic composition of deep-sea turbidites: Crustal evolution and plate tectonic association. Geochimica et Cosmochimica Acta 54, 2015 - 2050.

McLennan S. M., Hemming S., McDaniel D. K. and Hanson G. N. (1993) Geochemical approaches to sedimentation, provenance, and tectonics. In: Johnsson M. J. and Basu A. (eds.) Processes Controlling the Composition of Clastic Sediments. Geological Society of America Special Paper 284, 21 - 40.

Moores E. M. (2002) Pre-1 Ga (pre-Rodinia) ophiolites: their tectonic and environmental implications. GSA Bulletin 114, 80 - 95.

Moyen J.-F., Stevens G. and Kisters A. (2006) Record of mid-Archaean subduction from metamorphism in the Barberton terrain, South Africa. Nature 442, 559 - 562.

Myers R. E. (1990) The geology of the Godwan Basin, Eastern Transvaal. Unpublished PhD-Thesis, University of the Witwatersrand, 320 pp.

Nelson D. R., Trendall A.F., de Laeter J.R., Grobler N.J. and Fletcher I.R. (1992) A comparative study of the geochemical and isotopic systematics of late Archaean flood basalts from the Pilbara and Kaapvaal Cratons. Precambrian Research 54, 231 - 256.

Nelson D. R. (1998) Granite-greenstone crust formation on the Archean Earth: a consequence of two superimposed processes. Earth and Planetary Science

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Letters 158, 109 - 119.

Nelson D. R., Trendall A. F. and Altermann W. (1999) Chronological correlations between the Pilbara and Kaapvaal Cratons. Precambrian Research 97, 165 - 189.

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Peschler A. P., Benn K. and Roest W. R. (2004) Insights on Archean continental geodynamics from gravity modelling of granite-greenstone terranes. Journal of Geodynamics 38, 185 - 207.

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Rollinson H. (2006) Crustal generation in the Archean. In: Brown M. and Rushmer T. (eds.) Evolution and differentiation of the continental crust. Cambridge University Press, New York, NY, United States, 173 - 230.

Schweitzer J. and Kröner A. (1985) Geochemistry and petrogenesis of early Proterozoic intracratonic volcanic rocks of the Ventersdorp Supergroup, South Africa. Chemical Geology 51, 265 - 288.

Silver P. G., Behn M. D., Kelley K., Schmitz M. and Savage B. (2006) Understanding cratonic flood basalts. Earth and Planetary Science Letters 245, 190 - 201.

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Tinker J., de Wit M. and Grotzinger J. (2002) Seismic Stratigraphic Constraints on Neoarchean - Paleoproterozoic Evolution of the Western Margin of the Kaapvaal Craton, South Africa. South African Journal of Geology 105, 107 - 134.

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26 Chapter 1 Introduction

Kaapvaal and Pilbara Cratons between about 3.1 and 2.4 Ga: Abstract, 3rd International Archaean Symposium, Perth, p. 81.

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27 Chapter 1 Introduction

Winter H. de la R. (1976) A lithostratigraphic classification of the Ventersdorp succession. Transactions of the Geological Society of South Africa 79, 31 - 48.

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28 Chapter 2 Geological Setting

2 Geological Setting

The Neoarchaean (2.85 - 2.65 Ga) volcano-sedimentary Ventersdorp Supergroup extends today over an area of approximately 200000 km2 from southern Botswana into central South Africa (Fig. 2.1). Despite the large extent, most of the rocks of the Ventersdorp Supergroup are not exposed and fresh material is often only accessible in drill core.

Fig. 2.1: Geological map showing the present day distribution of the Ventersdorp Supergroup on the Kaapvaal Craton in relation to the Witwatersrand basin (modified after Eriksson et al. (2002) and Poujol et al. (2003)). KFN1: location of drill core KFN1 containing the Ventersdorp Supergroup. In its northeastern portion, the Ventersdorp Supergroup unconformably covers Au-U- bearing conglomerates of the Witwatersrand Supergroup, whereas an angular unconformity characterises the contact to the underlying Archaean basement in its southwestern portion (Tankard et al., 1982; Stanistreet et al., 1986). Sedimentary rocks of either the Neoarchaean Transvaal Supergroup or the Permian to Jurassic Karoo Supergroup unconformably overlie the Ventersdorp Supergroup. The Ventersdorp Supergroup attains a maximum thickness of approximately 5100 m, of which ca. 2900 m are sedimentary rocks, including volcanogenic and siliciclastic

29 Chapter 2 Geological Setting conglomerates, wackes, reworked tuffs and minor carbonate rocks (Winter, 1976). The sedimentary deposits are intercalated with thick successions of bimodal volcanic rocks. SHRIMP U-Pb age dating of magmatic zircons of the volcanic rocks of the Klipriviersberg Group and Makwassie Formation of the Ventersdorp Supergroup revealed ages of 2714 ± 8 Ma and 2709 ± 4 Ma, respectively (Armstrong et al., 1991). The top of the Ventersdorp Supergroup is not dated, but tuff beds at the base of the Transvaal Supergroup in Griqualand West reveal a maximum age of 2669 ± 5 Ma (Gutzmer and Beukes, 1998).

Despite its Neoarchaean age, the Ventersdorp Supergroup is hardly deformed and only weakly metamorphosed to lower greenschist-facies (Cornell, 1978; Tyler, 1979 a). Two distinct hydrothermal mineralization events have been recognized in the Allanridge lavas in the southwestern portion of the Ventersdorp Supergroup (de Bruiyn et al., 2002; Whitelaw et al., 2005). Both of these are tentatively related to metamorphic fluid flow (de Bruiyn et al., 2002; Whitelaw et al., 2005).

The distribution of the Ventersdorp Supergroup is influenced by various structural elements. McCarthy et al. (1990) describe feeder dykes for the eruption of the Klipriviersberg lavas that trend in a northeasterly direction. Clendenin et al. (1988 a) state that depositional environments of the Ventersdorp Supergroup are influenced by normal simple shear down-to-the-west that was followed by block faulting and the formation of graben structures that developed during the extension of the Platberg volcanic rocks. Extension occurred in both north-south and east-west directions (McCarthy et al., 1990). Furthermore, old existent faults were reactivated and new extensional, left-lateral, oblique slip faults initiated (McCarthy et al., 1990).

Rocks tentatively associated with sedimentary basins postdating the Ventersdorp but predating the Transvaal Supergroup, i.e. the Transvaal proto-basins, are preserved in the Transvaal and Griqualand West Basins (South Africa) and in the Kanye Basin of Botswana (Fig. 2.2; Eriksson and Altermann, 1998). In the Transvaal Basin, the remnants of the Transvaal proto-basins crop out along an arcuate belt along its northeastern margin striking from Swaziland to Polokwane in the Limpopo Province to Botswana, unconformably covering Archaean basement (Bosch et al., 1993).

30 Chapter 2 Geological Setting

Fig. 2.2: Outcrop of Ventersdorp Supergroup and outlines of Griqualand West and Transvaal Basins with outcrops of the Godwan Formation (GWF), Wolkberg (STF) and Buffelsfontein (TKF) Groups and Vryburg Formation (VBF) (modified after Eriksson et al. (2002) and Sumner and Beukes (2006)). KFN1: exploration drill core intersecting the entire Ventersdorp Supergroup; BT-VBF: Vryburg Formation near Bothithong; PK-VBF: Vryburg Formation near Prieska. The easternmost part of the Transvaal proto-basin is possibly represented by the ca. 1500 m thick volcano-sedimentary Godwan Formation, with outcrops limited to the Kaapsehoop area, Mpumalanga Province (Fig. 2.3). It presents a basin that could so far not be conclusively correlated to either the Transvaal, or the Witwatersrand Basin (Button, 1978). Its basal mafic volcanic rocks unconformably cover Archaean basement, whilst the Black Reef Formation of the lower Transvaal Supergroup unconformably covers the uppermost sedimentary unit of the Godwan Formation (Visser, 1956; Button, 1978). A Pb-Pb whole rock age of 2325 +69 / -73 Ma (Armstrong, 1987) is the only available geochronological data for the Godwan Formation. However, given the field relationship with the unconformably overlying 2588 ± 7 Ma Black Reef Formation (U-Pb zircon age, Martin et al., 1998), the Pb-Pb whole rock age of Armstrong (1987) must be discredited as being too young to record deposition.

31 Chapter 2 Geological Setting

Similar to the Ventersdorp Supergroup, the Godwan Formation does not show any sign of any significant metamorphic overprint. The lavas reveal epidotization, carbonatisation and K-metasomatism that partly obliterate the original igneous textures (Myers, 1990).

The Godwan Formation strikes northeast with a dip of 15- 80°, is partly folded and cut by faults that trend into three different directions: northeast, northwest and approximately east-west (Button, 1978). According to Visser (1956), most faults are of pre-Black Reef age. Moreover, pre-Transvaal diabase sills cross-cut the Godwan Formation (Visser, 1956). The contact to the unconformably overlying Black Reef Formation is characterised by an angular unconformity, as the Black Reef Formation dips westward at an angle of only 5° (Button, 1978).

Fig. 2.3: Geological map of the Godwan Formation in the vicinity of Kaapsehoop indicating the location of the samples collected in outcrops on the Farm Berlin (BL-GWF), as well as the location of the Farm Groot Geluk on which the sampled drill core GG1 is situated (modified after Walraven and Hartzer, 1986). Along the northeastern and northern margins of the Transvaal Basin, the Wolkberg and Buffelsfontein Groups constitute infills of proposed Transvaal proto-basins (Figs. 2.4 and 2.5; Button, 1973 a, b; Bosch, 1993; Barton et al., 1995). Both groups unconformably cover Archaean basement, and are in turn unconformably overlain by the Black Reef Formation of the basal Transvaal Supergroup (Bosch, 1993; Barton et al., 1995). The Wolkberg Group varies in its stratigraphic thickness, as it wedges out against palaeohighs (Button, 1973 a, b), but in palaeovalleys, it can reach up to

32 Chapter 2 Geological Setting

2600 m thickness (Button, 1973 a, b). The base of the Wolkberg Group consists of immature sedimentary rocks intercalated with basaltic lavas, while the upper Wolkberg Group comprises of mature arenites (Button, 1973 a, b).

The metamorphic overprint of the Wolkberg Group is negligible; the Abel Erasmus lavas display effects of silicification, carbonatisation and chloritisation. These are signs of hydrothermal alteration rather than metamorphism. The beds of the Wolkberg Group dip with about 10 - 15° to the west and southwest, or with up to 35- 48° to the south (de Villiers, 1976). Folding and pre- and post-Transvaal doming occurred with fold axes striking and plunging to the southwest, and fault systems of post-Transvaal age trend in three different directions, which are northeast to north, northwest and east (de Villiers, 1976).

Fig. 2.4: Geological map indicating the location of the outcrop of the Selati Formation, Wolkberg Group, from which samples were collected for this study (modified after du Plessis, 1978). The Buffelsfontein Group is preserved as infill in a ca. 2000 m deep palaeovalley that is bound by palaeohighs in the north and south (Tyler, 1979 b). Similar to the Wolkberg Group it is composed of bimodal volcanic and immature sedimentary rocks at the base and mature arenites towards the top (Tyler, 1979 b; Barton et al., 1995).

Conventional U-Pb age data exist for zircons from the Witfonteinrand lava of the Buffelsfontein Group, which were dated at 2664 ± 0.7 Ma (Barton et al., 1995)

33 Chapter 2 Geological Setting

(Fig. 1.2). This age may also apply to the lower Wolkberg Group, since the volcanic Abel Erasmus Formation is thought to be a lateral equivalent of the Witfonteinrand lava (Tyler, 1979 b; Barton et al., 1995) (Fig. 1.2).

The Witfonteinrand lava of the Buffelsfontein Group is hardly deformed and shows alteration to low-temperature mineral assemblages only (Barton et al., 1995). Minerals like feldspar and clinopyroxene have been replaced by sericite, carbonates and chlorite (Barton et al., 1995).

The strata of the Buffelsfontein Group dip southeast and are cross-cut by various mafic dykes and sills of ca. 1.3 Ga age (Barton et al., 1995). Furthermore, a post- Transvaal, steeply dipping, left-lateral wrench fault with displacement of ca. 15 km cross-cuts the Buffelsfontein Group, and is probably related to reactivation of the Thabazimbi-Murchison-Lineament (Barton et al., 1995).

Fig. 2.5: Geological map of the Buffelsfontein Group in the vicinity of Thabazimbi with the sampling site of the Tygerkloof Formation (KB-TKF). Fieldwork was carried out on the Farm Kransberg, which provides best access to the little exposed Tygerkloof Formation (modified after Jansen et al., 1974). The Schmidtsdrift Subgroup is the lowermost unit of the Griqualand West Basin (Figs. 2.6, 2.7, 2.8). It is assigned to the lower Transvaal Supergroup and unconformably covers the Allanridge Formation of the Ventersdorp Supergroup (Beukes, 1979). The stratigraphic thickness does not exceed 200 m and the succession is composed of a siliciclastic-carbonate section associated with tuffaceous beds at the

34 Chapter 2 Geological Setting base, followed by carbonaceous to argillitic and carbonate rocks (Beukes, 1979). Surface outcrops exist along the northern, southwestern and southern margins of the Griqualand West Basin (Beukes, 1979). Zircon-analyses from tuffaceous beds of the Vryburg Formation revealed ages of 2669 ± 5 Ma and 2650 ± 8 Ma (Gutzmer and Beukes, 1998).

Similar to the other studied successions the Schmidtsdrift Subgroup is hardly affected by deformation and metamorphism. In the area close to Bothithong, the Schmidtsdrift Subgroup dips at 2 - 8° away from an underlying broad, domal structure.

Fig. 2.6: Geological map of the northern margin of the Griqualand West Basin with sample localities of the Vryburg Formation in the vicinity of Bothithong (BT-VBF; BT-I-VBF), Northern Cape Province (modified after Moen, 1979).

35 Chapter 2 Geological Setting

Fig. 2.7: Geological map of the eastern margin of the Griqualand West Basin in the vicinity of Schmidtsdrift, where the type locality of the Schmidtsdrift Subgroup that comprises the Vryburg Formation, is exposed. Sample locality SD-VBF presents the type locality, whilst SCH-VBF includes the contact of the underlying Allanridge lavas to the covering Vryburg arenites (modified after Bosch and Visser, 1993).

36 Chapter 2 Geological Setting

Fig. 2.8: Geological map of the southern margin of the Griqualand West Basin in the vicinity of Prieska. The sample site PK-VBF that represents the southernmost known extent of the Vryburg Formation (modified after Malherbe and Moen, 1996). The Wolkberg and Buffelsfontein Groups, as well as the Godwan Formation have tentatively been correlated with each other (Fig. 2.9; Cheney et al., 1990; Cheney and Winter, 1995), as well as with the Pniel Group of the upper Ventersdorp Supergroup (Cheney et al., 1990, Barton et al., 1995). This correlation is based on sequence stratigraphic characteristics, but also on the availability of some geochronological data for the volcanic rocks of the lower Wolkberg and Buffelsfontein Groups (Barton et al., 1995). Based on ages of 2.65 Ga obtained on zircons derived from the tuff layers, the Schmidtsdrift Subgroup has recently been identified as a lateral correlative of the Transvaal proto-basinal successions in the Griqualand West region (Beukes et al., 2002; Dorland, 2004).

37 Chapter 2 Geological Setting

Fig. 2.9: Schematic lithostratigraphy and tentative correlation of the studied successions. The sampled formations are in bold italic and depositional ages are indicated. Fm: Formation. Schmidtsdrift Subgroup after Beukes (1979); Buffelsfontein Group after Barton et al. (1995); Wolkberg Group after Dorland (2004); Godwan Formation after Myers (1990). 1: U-Pb zircon ages, Gutzmer and Beukes (1998); 2: U-Pb zircon age, Martin et al. (1998); 3: U-Pb zircon ages, Armstrong et al. (1991); 4: U-Pb zircon age, Barton et al. (1995); 5: ID-TIMS zircon age, Anhaeusser and Poujol (2004); 6: ID-TIMS zircon age, Kamo and Davis (1994).

2.1 Previous Work

Winter (1965, 1976) was the first studying the Ventersdorp Supergroup in its entire context, dividing it into the formations as they are being used today. Other studies concentrated mostly on the geochemistry of the volcanic rocks of the Ventersdorp Supergroup, especially the ultramafic lavas of the basal Klipriviersberg Group (McIver et al., 1982; Schweitzer and Kröner, 1985; Linton et al., 1990; Myers et al., 1990 a; Marsh et al., 1992; Nelson et al., 1992; Hatton, 1995; White, 1997), and the other volcanic units (Schweitzer and Kröner, 1985; Bowen et al., 1986; Crow and Condie, 1988; Nelson et al., 1992; Linton and McCarthy, 1993; Keyser, 1998; de Bruiyn et al. 2002; Gehring et al., 2005, 2006; Whitelaw et al., 2005), as well as on special aspects, such as stromatolites, and structural and facies analyses of single formations (Buck, 1980; Visser and Grobler, 1985; Clendenin et al., 1988 b; Karpeta, 1989; van der Westhuizen et al., 1989; McCarthy et al., 1990; Eriksson et al., 1998; van der Westhuizen and de Bruiyn, 2000; Hall and Els, 2002). Cornell (1978) and

38 Chapter 2 Geological Setting

Tyler (1979 a) studied the alteration of the Ventersdorp Supergroup, and established that the Ventersdorp Supergroup is one of the least deformed and metamorphosed succession of Neoarchaean age on Earth.

Previous studies of the Wolkberg and Buffelsfontein Groups, Schmidtsdrift Subgroup and Godwan Formation are limited and concentrate on petrographic, sedimentological and rarely on geochemical analyses. Lithostratigraphic descriptions of the Godwan Formation, Wolkberg Group and Schmidtsdrift Subgroup were carried out by Visser (1956), de Villiers (1976) and Beukes (1979), respectively. The focus of several other studies was the correlation between the different formations of the Transvaal proto- basins, and their relation to the upper Ventersdorp Supergroup. Tyler (1979 b) used stratigraphic and geochemical analyses of lavas in combination with characteristics of depositional environments to establish a correlation between the Wolkberg and Buffelsfontein Groups, which are geographically relatively closely related (Fig. 2.2). Based on sequence stratigraphic analyses of unconformity-bound sequences, the Pniel Group of the Ventersdorp Supergroup has been correlated with the upper Wolkberg Group, the upper Buffelsfontein Group and the upper Godwan Formation (Cheney et al., 1990; Cheney and Winter, 1995). Barton et al. (1995) proposed a relationship between the upper Buffelsfontein Group and Pniel Group of the upper Ventersdorp Supergroup, a suggestion based on Rb-Sr isotopic systematics. However, the correlation of the Godwan Formation seems to be the most difficult, since Button (1978) related it to the upper Witwatersrand Supergroup, whilst Myers (1990) suggested a correlation of the Godwan lavas with the Nsuze Group of the Pongola Supergroup. Beukes et al. (2002) extended the proposed correlation of the Transvaal proto-basins to the Schmidtsdrift Subgroup of the Griqualand West Basin.

In addition to attempts to correlate the different successions with each other, studies focusing on the depositional history and palaeoenvironments (Button, 1973 a, b; Bosch et al., 1993), including the study of algal stromatolites (Button, 1973 c) of the Wolkberg Group have been performed. However, detailed palaeoenvironmental reconstructions exist only for the Schmidtsdrift Subgroup (Beukes, 1977).

Models for the tectonic setting of the Ventersdorp Supergroup have been presented by numerous authors. Some support a modern subduction-related rift setting (Burke et al., 1985; Stanistreet et al., 1986; Clendenin et al., 1988 b; Crow and Condie, 1988; Silver et al., 2006) or an Atlantic-type passive continental margin (Tinker et al.,

39 Chapter 2 Geological Setting

2002), whilst others prefer an Archaean plume-related setting (e.g. Marsh et al., 1992; Nelson et al., 1992; Hatton, 1995; White, 1997; Arndt, 1999; Eriksson et al., 2002; Ernst and Buchan, 2002; Barley et al., 2005; Ernst et al., 2005) or even compare the Ventersdorp Supergroup to greenstone belts (Schweitzer and Kröner, 1985; Cheney et al., 1990; McIver et al., 1992; Eriksson et al., 2001).

Studies regarding the character of the major sedimentary successions (Kameeldoorns and Bothaville Formations), and an evaluation of their plate tectonic positions are scarce. Wronkiewicz and Condie (1990) carried out preliminary petrographic and geochemical analyses on the Bothaville Formation only, while Jahn and Condie (1995) compared the Bothaville Formation to sedimentary units of the Witwatersrand and Transvaal Supergroups based on whole rock geochemical analyses and Nd- isotopic systematics. Only Visser et al. (1976) compared the depositional environments of the Kameeldoorns and Bothaville Formations in a concise facies model. A literature survey reveals that previous studies of the Ventersdorp Supergroup took into consideration only single formations or applied only single analytical methodologies. None of the studies presents a combination of different approaches that would result in a more comprehensive interpretation of the Ventersdorp Supergroup. Also, none of the previous studies used integrated information from volcanic and sedimentary units of the Ventersdorp Supergroup to create a convincing basin evolution model, and no study has elaborated on the implications of the evolution of the Ventersdorp Supergroup on the stabilization of the Kaapvaal Craton in the Neoarchaean.

It is thus the objective of this study to apply a combination of various provenance- indicating methods, e.g. detailed geochemistry, Nd-isotopic systematics, U-Pb detrital zircon age dating and petrographic observations to the Ventersdorp Supergroup and Transvaal proto-basin successions. The resulting data yield insight into the correlation between the different studied successions. It is expected that the comprehensive provenance analyses will reveal possible correlations more clearly than using specific analytical techniques only. Furthermore, the comprehensive approach shall also be used to give insight in the stabilization process of the Kaapvaal Craton and its palaeotectonic evolution in the Neoarchaean.

40 Chapter 2 Geological Setting

2.2 Ventersdorp Supergroup

2.2.1 Lithostratigraphy

As reference for the lithostratigraphy, drill core KFN1 was used, because it intersects the entire Ventersdorp Supergroup with all its formations as described by Winter (1976). It was drilled for exploration purposes in the northern Free State, a few kilometers northwest of the mining town of Allanridge on the Farm Kruidenfontein (Fig. 2.10). The Ventersdorp Supergroup (3610 m total thickness in KFN1) is composed of three groups, which are from the base to the top the Klipriviersberg (1785 m thick in KFN1), Platberg (1406 m thick in KFN1) and Pniel (418 m thick in KFN1) Groups (Fig. 2.11). These are further divided into formations according to characteristic lithofacies as described by Winter (1976).

Fig. 2.10: Location of KFN1 and other examined drill cores. The NS- and WE-sections are illustrated in Figs. 2.12 and 2.13. The X- and Y-coordinates are of the South African coordinate system, which is normally used to indicate the locations of exploration drill cores (map courtesy of Avgold Ltd., 2003).

41 Chapter 2 Geological Setting

Following Winter (1976), the Klipriviersberg Group is subdivided into six formations, based on macroscopic appearance of the lavas. These are talcose tuffs (Westonaria Formation), porphyritic lava (Alberton Formation), purple-topped ash-flow tuffs (Orkney Formation), agglomerates (Jeannette Formation), mafic effusive rocks with characteristic variolitic and spherulitic textures (Loraine Formation) and lavas rich in clusters of milky quartz amygdales (Edenville Formation; Winter, 1976). Myers et al. (1990 b), however, defined ten formations based on geochemical data. Since the detailed analysis of the Klipriviersberg lavas is not the scope of this thesis, the whole group was treated as one entity. In KFN1, two mafic dykes crosscut the Klipriviersberg lavas at 3368 m and 3023 m depth.

Fig. 2.11: Stratigraphy of the Ventersdorp Supergroup in drill core KFN1. The Kameeldoorns and Bothaville Formations (bold and italic) were the objectives of this study. 1: U-Pb zircon ages of Armstrong et al. (1991). Fm: Formation. The Platberg Group is subdivided into three formations: the Kameeldoorns, Makwassie and Rietgat Formations. The Kameeldoorns Formation consists of a variety of sedimentary rock types with a thickness of 435 m in KFN1, interrupted by 245 m of Klipriviersberg lavas at the base of the formation, so that the total thickness of the Kameeldoorns Formation amounts to 680 m in KFN1. Such interfingering of the uppermost lava flows of the Klipriviersberg Group with volcanogenic Kameeldoorns conglomerates is not unusual and has been described by Winter

42 Chapter 2 Geological Setting

(1976). The base of the Kameeldoorns Formation is composed of volcaniclastic conglomerates, which show vast changes in thickness in different drill cores (Figs. 2.12 and 2.13), and grade towards the top from clast-supported volcanogenic conglomerates into matrix-supported siliciclastic, i.e. non-volcaniclastic, conglomerates, until they finally give way to siliciclastic wackes. Single wacke horizons are intercalated with the conglomerates. Near the top, the wackes are texturally more mature and possess a higher amount of quartz than stratigraphic lower wacke horizons. Furthermore, they are intercalated with fine-grained reworked ash deposits that encase stromatolitic carbonate horizons.

In KFN1, the Kameeldoorns Formation is overlain by 396 m of felsic lava flows of the Makwassie Formation. They are interrupted by single, up to 0.5 m thick wacke horizons. The Rietgat Formation, which unconformably follows on top of the Makwassie Formation, also contains felsic lava flows (330 m in KFN1) and locally coarse volcanogenic conglomerates, as well as stromatolitic carbonates (Karpeta, 1989). In KFN1, however, the Rietgat Formation is entirely volcanic, consisting of greenish-grey lava flows that are locally strongly pale yellowish-green discoloured.

The uppermost Pniel Group unconformably overlies the Platberg Group and is divided into the Bothaville and Allanridge Formations. The Bothaville Formation, 362 m thick in KFN1, is composed of compositionally mature wackes of mostly medium sand grain size, framed by siliciclastic conglomerates at the base and the top. Single tuffaceous shale horizons with associated stromatolitic carbonates are intercalated with the wackes, which are partly volcaniclastic.

The Allanridge Formation concludes the Ventersdorp Supergroup and is represented by mafic lava flows. In KFN1, sedimentary rocks of the Phanerozoic Karoo Supergroup unconformably cover the Allanridge Formation.

43 Chapter 2 Geological Setting

Fig. 2.12: WE-section illustrating the lateral distribution of the Ventersdorp Supergroup in the vicinity of drill core KFN1. Vertical exaggeration is five times. Note strong thickness variations of the Kameeldoorns Formation, compared to the almost constant thickness of the Bothaville Formation. Fm: Formation. See Appendix V for detailed logs.

44 Chapter 2 Geological Setting

Fig. 2.13: NS-section illustrating the lateral distribution of the Ventersdorp Supergroup in the vicinity of drill core KFN1. Vertical exaggeration is nine times. Note the prominent thickness variations of the Kameeldoorns Formation, whilst the thickness of the Bothaville Formation remains almost constant. Fm: Formation. See Appendix V for detailed logs.

45 Chapter 2 Geological Setting

2.2.2 Palaeoenvironmental Significance

Various authors relate the mafic to ultramafic lavas at the base of the Klipriviersberg Group to a plume event and interpret them as flood basalts (e.g. Hatton, 1995; White and McKenzie, 1995; White, 1997), while others interpret them as a subduction zone- related volcanic suite (e.g. Marsh et al., 1989; Crow and Condie, 1990).

The rapidly changing lithologies and thicknesses of the Kameeldoorns Formation, as displayed in different drill cores in close proximity to each other (Figs. 2.12, 2.13), are interpreted as evidence for deposition in an active graben system. Extension (rifting) is unambiguously suggested by many authors (e.g. Winter, 1976; Buck, 1980; Burke et al., 1985; Clendenin et al., 1988 b), and sometimes thought to follow the plume event that may be responsible for the eruption of the Klipriviersberg lavas (Eriksson et al., 2002). Furthermore, the upper portion of the Kameeldoorns Formation with its stromatolitic carbonates and volcanic ash deposits is believed to have formed in lakes that were fed by braided river systems (Visser et al., 1976). Raindrop imprints, found in a Platberg correlative (Sodium Group; van der Westhuizen et al., 1989), support a terrigeneous setting for the Kameeldoorns Formation.

Several authors compare the lower portion of the Ventersdorp Supergroup, that is the Klipriviersberg and Platberg Groups, to Archaean greenstone belts, as they show a typical greenstone belt development with komatiitic lava flows at the base, followed by immature rift sedimentation and bimodal volcanism (McIver et al., 1982; Schweitzer and Kröner, 1985; Cheney et al., 1990).

The Makwassie lavas represent fractionated crustal melts, typical for later stages of plume-related flood basalt provinces (White, 1997; Eriksson et al., 2002). Characteristic lithofacies such as pumice breccias and ignimbrites imply a subaerial to shallow water setting (Gehring et al., 2006). Compared to the Makwassie Formation, the overlying Rietgat Formation records only episodic volcanism, where volcanic and few sedimentary rocks alternate (Visser and Grobler, 1985; Eriksson et al., 2002). Karpeta (1989) and Eriksson et al. (1998) assign a saline playa lake setting to the sparse stromatolitic carbonates related to penecontemporaneous volcanic activity in the Rietgat Formation.

46 Chapter 2 Geological Setting

The Bothaville Formation was deposited in a basin controlled by thermal subsidence (Clendenin et al., 1988 a) that pinches out against palaeohighs (Winter, 1976). Its facies is interpreted as continental fluvial and lacustrine, associated with terrestrial volcanism peripheral to the depositional basin (Eriksson et al., 2002). Buck (1980) suggested regional extensive alluvial plains dominated by braided rivers, where stromatolitic structures could develop in pools. In contrast, Visser et al. (1976) propose marine transgression with a beach or shallow marine environment for the Bothaville Formation. However, there seems to be no major time gap between the Rietgat volcanism and the Bothaville sedimentation (Visser and Grobler, 1985). Visser et al. (1976) describe pillowed lavas at the base of the Allanridge Formation, which clearly favours a subaqueous setting, either lacustrine or marine. It has been suggested that Allanridge volcanism, similar to Klipriviersberg volcanism, is plume- related (Eriksson et al., 2002).

2.3 Transvaal proto-basinal successions

2.3.1 Godwan Formation

2.3.1.1 Lithostratigraphy

The Godwan Formation unconformably covers Archaean basement, which consists of hornblende granite of the Kaap Valley pluton (3227 ± 1 Ma ID-TIMS zircon age; Kamo and Davis, 1994) in the Kaapsehoop area of the Eastern Kaapvaal Craton (Visser, 1956). Stratigraphically, it is divided from the base upwards into a Lower Sedimentary Stage, a Volcanic Stage, and an Upper Sedimentary Stage, which together reach a total thickness of ca. 1500 m (Visser, 1956) (Figs. 2.2 and 2.3).

Myers (1990) subdivided the Godwan Formation into seven formations: the Kaapsehoop and Amianthus Formations at the base, equivalent to Visser’s Lower Sedimentary Stage; the Berlin and Coetzeestrom Formations that correspond to Visser’s Volcanic Stage; the Battery Creek, Vlakplaats and Nooitgedacht Formations represent the above mentioned Upper Sedimentary Stage (Fig. 2.14). The Kaapsehoop Formation is between 50 m and 100 m thick and composed of heterogeneous sedimentary rocks, such as coarse diamictites, turbiditic arenites and quartz-arenites, whilst the Amianthus Formation is represented by a 20 - 30 m thick gritty, matrix- poor, feldspathic arenite (Myers, 1990). The 200 m thick Berlin Formation is entirely

47 Chapter 2 Geological Setting volcanic and composed of mafic amygdaloidal lava flows with intercalated horizons of sedimentary rocks, whereas the covering Coetzeestrom Formation comprises 50 - 200 m thick volcaniclastic rocks with intercalated lavas and rare quartz-rich arenites (Myers, 1990). The Coetzeestrom Formation is overlain by the basal conglomerate of the 200 - 250 m thick Battery Creek Formation that consists predominantly of medium to coarse grained quartz-arenites and wackes (Myers, 1990). The overlying Vlakplaats Formation is comprised of fine-grained wackes that are between 300 m and 500 m thick, whilst the uppermost Nooitgedacht Formation comprises 80 - 100 m of coarse, sandy and matrix-supported diamictites with a coarse, pebbly arenite at the base of the formation (Myers, 1990). The Black Reef Quartzite Formation of the Transvaal Supergroup unconformably covers the top of the Godwan Formation.

Fig. 2.14: Stratigraphy of the Godwan Formation after Myers (1990). The sampled formations are in bold and italic. 1: ID-TIMS zircon age, Kamo and Davis (1994); 2: U-Pb zircon age, Martin et al. (1998). Fm: Formation.

2.3.1.2 Palaeoenvironmental Significance

The Godwan Formation records several distinctly different depositional environments. The basal Kaapsehoop Formation is interpreted as shallow marine- marginal environment with lateral rapidly changing facies, whereas the overlying Amianthus Formation is thought to represent a fluvial setting with bed load deposits

48 Chapter 2 Geological Setting

(Myers, 1990). The volcanic Berlin and Coetzeestrom Formations are subaerial, whereas the sedimentary intercalations imply periods of volcanic quiescence and deposition in fluvial or shallow lacustrine environments. The volcaniclastic debris flow deposits of the Coetzeestrom Formation were deposited in alluvial fans, and were described by Myers (1990) as reflecting ‘catastrophic events’. The upper sedimentary formations suggest tectonic stability and shallow marine depositional settings, which include a transgressive event reflected by the fine-grained wackes of the Vlakplaats Formation (Myers, 1990). The topmost Nooitgedacht Formation, however, displays a change back to a fluvial depositional environment and suggests a marked marine regression (Myers, 1990). Palaeotectonic interpretations or models of this formation do not exist so far.

2.3.2 Wolkberg Group

2.3.2.1 Lithostratigraphy

The Wolkberg Group of the northeastern margin of the Transvaal Basin reaches locally up to 2600 m thickness (Bosch et al., 1993), but only 700 m in the area around Chuniespoort (de Villiers, 1976), which was the sample location for this study. It unconformably covers granitic Archaean basement and is subdivided into six formations (Fig. 2.15). These are in stratigraphic order from the base upwards the Sekororo, Abel Erasmus, Schelem, Selati, Mabin and Sadowa Formations (Button, 1973 a, b).

The Sekororo Formation at the base consists of feldspathic arenites and conglomerates and is laterally discontinuous (Button, 1973 b). It is covered by basaltic lavas and discontinuous layers of quartz-arenites and shales of the Abel Erasmus Formation (Button, 1973 b; Bosch et al., 1993). The overlying Schelem Formation is mainly composed of feldspathic arenites, quartz-arenites, conglomerates and shales (Button, 1973 b; Bosch et al., 1993; Dorland, 2004). The Selati Formation comprises mature quartz-arenites intercalated with horizons of tuffaceous shales, and is framed at the bottom and top by shale layers, and was sampled as the best mixed and representative formation for provenance purposes. The Mabin Formation is another succession of quartz-arenites with subordinate layers of shales (de Villiers, 1976), while the uppermost Sadowa Formation consists mainly of shales associated with only a minor amount of quartz-arenites (Button, 1973 b; de Villiers, 1976). An

49 Chapter 2 Geological Setting erosional unconformity marks the boundary to the overlying Black Reef Formation of the Transvaal Supergroup (Bosch et al., 1993).

Fig. 2.15: Stratigraphy of the Wolkberg Group after Dorland (2004). The Selati Formation (bold and italic) was sampled for this study. 1: U-Pb zircon age, Martin et al. (1998). Fm: Formation.

2.3.2.2 Palaeoenvironmental Significance

Different depositional environments are reflected by the lower and upper portions of the Wolkberg Group. The lower Wolkberg Group (Sekororo, Abel Erasumus and Schelem Formations) represents fluvial environments with deposition in braided streams on alluvial fans and braidplains (Button, 1973 b; Bosch et al., 1993). The volcanic rocks of the Abel Erasmus Formation are interpreted as continental within- plate lavas (Bosch et al., 1993). These were at least in parts extruded in subaqueous environments, as pillows are described by Button (1973 b). Stromatolitic carbonates and carbonaceous mudstones intercalated with the lavas imply periods of magmatic quiescence and deposition in a shallow lacustrine or marine environment (Button, 1973 a, b; Bosch et al., 1993). Furthermore, deposition of the lower Wolkberg Group took place on uneven surfaces, where it filled depressions. By the onset of the deposition of the upper Wolkberg Group, this palaeorelief was evened out (Button, 1973 b; Bosch et al., 1993). Thus, the upper Wolkberg Group is developed in larger basins, in which the Selati Formation developed first with characteristic pro-delta and

50 Chapter 2 Geological Setting delta-front deposits, and then changed to shallow water marine or lacustrine environments of the Sadowa Formation (Bosch et al., 1993). Beach deposits dominate the Mabin Formation of the topmost Wolkberg Group (Bosch et al., 1993).

2.3.3 Buffelsfontein Group

2.3.3.1 Lithostratigraphy

The Buffelsfontein Group reaches a thickness of ca. 2700 m in the area of the Farm Kransberg (Barton et al., 1995), which was also the sampling site for this study. It unconformably covers granitic Archaean basement (Makoppa Dome, 2886 +3/-2 Ma; ID-TIMS zircon age, Anhaeusser and Poujol, 2004) and is divided into four formations, namely the Hampton, Waterval, Witfonteinrand and Tygerkloof Formations (Fig. 2.16; Barton et al., 1995).

Fig. 2.16: Stratigraphy of the Buffelsfontein Group on the Farm Kransberg (after Barton et al., 1995). The Tygerkloof Formation (bold and italic) was selected for this study. 1: U-Pb zircon age dated by Barton et al. (1995); 2: U-Pb zircon age, Martin et al. (1998). Fm: Formation.

51 Chapter 2 Geological Setting

The Hampton Formation at the base of the Buffelsfontein Group is composed of lithic arenites and minor conglomerates, whereas the unconformably covering Waterval Formation consists entirely of lavas of intermediate composition (Barton et al., 1995). The Witfonteinrand Formation is framed by unconformities and consists of a variety of bimodal lavas, pyroclastic and sedimentary rocks, as well as stromatolitic chert (Barton et al., 1995), whilst the Tygerkloof Formation consists of quartz-rich arenites with single intercalated conglomerate horizons. The top of the Tygerkloof Formation is defined by an erosional unconformity that marks the contact to the overlying Black Reef Formation of the Transvaal Supergroup.

2.3.3.2 Palaeoenvironmental Significance

The strata of the Buffelsfontein Group represent depositional environments ranging from fluvial terrigeneous to high energy shallow marine (Tyler, 1979 b). The Hampton Formation formed in a meandering stream setting (Tyler, 1979 b), followed by a high-energy near-shore marine setting for the Hampton quartz-rich arenites. This shallow marine setting grades back into a fluvial braided river environment (Tyler, 1979 b). The overlying volcanic rocks extruded subaerially, as they display terrestrial flow-features and lack pillow structures (Tyler, 1979 b). During periods of volcanic quiescence, stromatolitic carbonates were precipitated in standing pools and clastic sedimentary rocks deposited in fluvial settings (Tyler, 1979 b). The mature quartz arenites of the topmost Tygerkloof Formation give evidence of a high-energy, tidal influenced near-shore marine environment (Tyler, 1979 b).

2.3.4 Schmidtsdrift Subgroup

2.3.4.1 Lithostratigraphy

The Schmidtsdrift Subgroup of the Ghaap Group in Griqualand West unconformably overlies lavas of the Allanridge Formation of the Ventersdorp Supergroup and is approximately 230 m thick (Fig. 2.17; Beukes, 1979). Its exposures along the northern, eastern and southern margins of the Griqualand West Basin were of interest of this study to outline possible differences across the Griqualand West outcrop region.

The Schmidtsdrift Subgroup is divided into the Vryburg, Boomplaas and Lokammona Formations, and is conformably overlain by carbonate rocks of the Campbellrand

52 Chapter 2 Geological Setting

Subgroup (Beukes, 1979). The Vryburg Formation at the base is composed of quartz- arenites that are interlayered with single tuff horizons, thin shales and carbonate beds. The Boomplaas Formation is comprised mostly of stromatolitic and oolithic carbonates and shales, whereas the Lokammona Formation at the top of the Schmidtsdrift Subgroup is entirely composed of carbonate-rich shales (Beukes, 1979; Dorland, 2004).

Fig. 2.17: Stratigraphy of the Schmidtsdrift Subgroup in the area of Bothithong after Beukes (1979). In this study, the basal Vryburg Formation (bold and italic) was examined. 1: U-Pb zircon ages of Gutzmer and Beukes (1998); 2: U-Pb zircon age by Martin et al. (1998). Fm: Formation.

2.3.4.2 Palaeoenvironmental Significance

The Schmidtsdrift Subgroup of the Griqualand West Basin displays an entirely marine environment (Beukes, 1977). The Vryburg Formation at the base represents a progradational subtidal-tidal flat, which turns into a transgressive tidal flat and prograding carbonate subtidal-tidal flat in the overlying Boomplaas Formation (Beukes, 1977). The topmost Lokammona Formation with its carbonaceous deposits implies a progradational tidal flat environment (Beukes, 1977).

53 Chapter 2 Geological Setting

2.4 References

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54 Chapter 2 Geological Setting

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55 Chapter 2 Geological Setting

the Witwatersrand, Economic Geology Research Unit 201, 20 pp.

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56 Chapter 2 Geological Setting

introduction to Precambrian basins: their characteristics and genesis. Sedimentary Geology 141-142, 1 - 35.

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57 Chapter 2 Geological Setting

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