Bokkeveld Group: Cape Supergroup), Western Cape Province, South Africa

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Facies analysis, palaeo – environmental successions and sequence stratigraphy of the Early to Middle Devonian Ceres Subgroup (Bokkeveld Group: Cape Supergroup), Western Cape Province, South Africa

Cameron Roy Penn-Clarke

Volume I

A dissertation submitted as fulfilment for the degree of Master of Science at the University of the Witwatersrand, Johannesburg

Johannesburg, 2013 I

Declaration

I, Cameron Roy Penn-Clarke, declare that this Dissertation is my own, unaided work. It is being submitted for the Degree of Master of Science at the University of the Witwatersrand, Johannesburg. It has not been submitted before for any degree or examination at any other University.

______

(Signature of candidate)

On the______day of______20 at______

Abstract

From fieldwork intensive research on the sedimentology, stratigraphy and palaeontology of the Ceres Subgroup (Bokkeveld Group: Cape Supergroup), a total of 12 lithofacies and 8 lithofacies associations have been recognised and described between the farm Grootrivierhoogte (GPS: 32°38'55"S; 19°24'43"E) and the town of Wupperthal (GPS: 32°18'16"S; 19°14'24.88"E) in the Cederberg, Western Cape, South Africa. Based on comparison with modern and ancient marginal – marine and shallow – marine depositional environments, the lithofacies associations recognised in this project are representative of sedimentation within 8 distinct sub – environments. These sub – environments, in turn, have been grouped into 2 larger environments of deposition, namely a wave and storm dominated shoreline and a wave and tidal influenced estuary.

Wave and storm dominated shoreline environments are restricted to the Gydo – Gamka and Voorstehoek – Hex River Systems only. Here Lithofacies Associations A – D represent sedimentation within Os, OT – dLSF, pLSF and USF – Beach Complex sub – environments respectively. Both systems are regressive and conformable with one another following transgression. Palaeontological finds within these systems indicate that Os sub – environments are the most species rich and are dominated by typical Malvinokaffric Realm Biota. Fossils found in interpreted Os sub – environments are usually disarticulated with a low degree of damage. This suggests a major benthic Os lifestyle for the Malvinokaffric Realm Biota in calm marine conditions below storm wave base. OT – dLSF and pLSF sub – environments contain a more restricted assemblage of fossils and are restricted to the brachiopods Australocoelia cf. tourteloti and Australospirifer sp. indet. as well as crinoids (based on disarticulated ossicles). Fossils found in OT – dLSF and pLSF sub – environments are always disarticulated and highly damaged suggesting storm and wave activity entraining fossils (prior to fossilisation) and damaging them at or above storm wave base within these sub – environments. Plant fossils within these sub – environments are all interpreted to be allocthonous and were brought in with offshore directed storm and wave activity. Dubious plant rootlet ichnofossils have III

been found in USF – Beach complex environments and possibly may be from a Bs sub – environment.

A wave and tide influenced estuarine environment is present only in the Tra-Tra – Boplaas System. Lithofacies Associations E – H represent sedimentation within respective outer estuary (channelised tidal flat), inner estuary (lagoon to washover flats) and inner estuary (estuary bay to subaqueous and subaerial bayhead estuary) sub – environments. Australospirifier sp. indet. fossils have been found, often within communities, in interpreted washover fan sub – environments in addition to plant fossils, plant rootlet ichnofossils and Skolithos sp. and surface cast ichnofossils suggesting a mixed salinity environment with marsh development in lagoons. Coalified plant fossils are most numerous in inner estuary environments.

The facies control of Malvinokaffric Realm fossil occurrences indicate that biostratigraphy for the Bokkeveld Group does not appear possible as the majority Malvinokaffric Realm Biota appear to have existed within an Os sub – environment.

A sequence stratigraphic analysis of the Ceres Subgroup has revealed a total of five parasequences. These are bound by M.R.S. sequence stratigraphic surfaces and have been associated with 3rd order T-R cycles. Parasequences 1.1 – 1.4 represent T-R cycles within the Gydo – Gamka and Voorstehoek – Hex River Systems. The Gydo – Gamka and Voorstehoek – Hex River Systems are regressive and conformable with each other with the Gamka-Voorstehoek transition representing a T.S.T. The Gydo – Gamka System is associated with the 1st 2nd order flooding event into the Cape Basin marking the start of Bokkeveld Group sedimentation during the late Emsian and is hypothesised to have started earlier into Rietvlei time sedimentation. A total of 3 sea level rises and falls appear to have been evident during Voorstehoek – Hex River time sedimentation as parasequences 1.2 – 1.4 are present within it. The contact between the Hex River and Tra-Tra Formations is paraconformable and appears to represent a S.R. – U associated with a large and 2nd 2nd order flooding event within the Cape Basin. The Tra-Tra – Boplaas System is a part of parasequence 2.1 and is overall transgressive representing a T.S.T. which is hypothesised to continue into the Waboomberg Formation and terminate at an M.R.S. somewhere in the Wupperthal Formation. IV

For my grandfather, George Melbourne Penn-Clarke 17 February 1927 – 01 February 2013

Acknowledgements

I would like to personally acknowledge the numerous people and organisations involved in this dissertation, without which it would not be a reality.

• My supervisors, Prof. Bruce Rubidge, Dr. Zubair Jinnah and Dr. John Almond for their supervision, support, corrections of various drafts and helpful discussions. • The National Research Fund (NRF) and the Palaeontological Scientific Trust (PAST) for financial support. • The South African Heritage Resources Agency for granting me a fossil collecting permit. • The farmers and shareholders of the Cederberg Wilderness Area for allowing access onto their properties and keeping me sane. Special mention must be made to the Marais and Vorster families at Mount Ceder for accommodation. • Ms. Rika du Plessis, manager of Matjiesrivier Reserve for facilitating my stay at the Reserve and for allowing ranger Mr. Willem Titus to help me during fieldwork exercises in Wupperthal. • Profs. Arthur Boucot (Oregon State University), Peter Isaacson (University of Idaho), Dr. Norton Hiller (University of Canterbury) and Dr. John Maisey (American Museum of Natural History) for insightful discussions on the Malvinokaffric Realm Biota and furnishing of information. • Dr. Anthony Tankard (Tankard Industries) for insightful discussions on Cape Basin Geology and information. • Mr. Eugene Bergh of Iziko Museum, Cape Town for allowing me access to the invertebrate collections. • Ms. Claire Browning from Council for Geoscience Bellville, for discussions on the Bokkeveld Group and facilitating access to the palaeontological collections at CGS Bellville. • Dr. John Almond and his wife, Mrs. Madelon Tusenius for giving me a warm meal and a place to stay whilst in Cape Town. • My family who have supported my geological passions and endeavours since a child. • My girlfriend, Ms. Natalie Brand, for her support during the best and worst of times. VI

Contents Page

Declaration I Abstract II Dedication IV Acknowledgements V

Chapter 1 Introduction 1

1.1 Palaeotectonic Setting of the Cape Basin and the Cape Supergroup 5

1.1.1 Base of the Cape Basin 6 1.1.2 Subsidence Analysis of the Cape Basin 7 1.1.3 Models for the Tectonic Setting of the Cape Basin 11 1.1.3.1 Back-arc Basin Model (Catuneanu, et al. (1998)) 12 1.1.3.2 Margin-sag-Interior-sag Basin (Broquet, (1992)) 13 1.1.3.3 Episutural Aulacogen Basin Model (Tankard, et al. (1982)) 15 1.1.3.4 Episutural Sag Basin Model (Tankard, et al. (2009)) 15

1.2 Structural Features Associated with the Cape Fold Belt 16

1.3 Geology of the Bokkeveld Group 19

1.3.1 Palaeoenvironmental Setting of the Bokkeveld Group 19 1.3.2 Allocyclic vs. Austocyclic Models for Megacycle Cyclicity 21

1.4 Lithostratigraphy of the Bokkeveld Group in the Western Cape Province 24 1.5 Palaeontology of the Bokkeveld Group 26

1.5.1 The Malvinokaffric Realm 27 1.5.2 The Faunal Assemblage and Attributes of the Malvinokaffric Realm 29 1.5.3 Previous Attempts at Establishing a Biostratigraphy for the 29 Bokkeveld Group VII

Chapter 2 Aims and Methods 31

2.1 Aims 31

2.2 Materials and Methods 31

Chapter 3 Geology of Study Area 33

3.1 General Geology and Geography of Study Area 33

3.2 Introduction to Sedimentary Facies Analysis 34

3.3 Current lithofacies model for the Bokkeveld Group (Theron, 1972, Tankard 35 and Barwis, 1982, Theron and Loock, 1988)

3.3.1 Storm/wave dominated Delta Model 35 (Theron, 1972, Theron and Loock, 1988) 3.3.2 Mixed wave/tidal dominated Delta Model 36 (Tankard and Barwis, 1982)

3.4 Lithofacies Descriptions 41

3.4.1 Lithofacies 1 (Clayshale Lithofacies) 42 3.4.2 Lithofacies 2 (Siltstone Lithofacies) 43 3.4.2.1 Sub-lithofacies 2.1 44 (Plane -laminated Siltstone Sub- lithofacies)

3.4.2.2 Sub-lithofacies 2.2 44 (Ripple bedded Siltstone Sub-lithofacies) 3.4.3 Lithofacies 3 46 (Heterolithic Silty Sandstone and Mudstone Lithofacies) 3.4.4 Lithofacies 4 49 (Hummocky and Swaley Cross Stratified Sandstone Lithofacies) 3.4.4.1 Palaeocurrent Indicators Associated with Lithofacies 4 53 VIII

3.4.5 Lithofacies 5 (Coarse Grained Ripple Sandstone Lithofacies) 54 3.4.5.1 Palaeocurrent Indicators Associated with Lithofacies 5 58 3.4.6 Lithofacies 6 (Lenticular Wacke Lithofacies) 59 3.4.7 Lithofacies 7 (Macroform Cross-bedded Sandstone Lithofacies) 62 3.4.7.1 Palaeocurrent Indicators Associated with Lithofacies 7 66 3.4.8 Lithofacies 8 (Planar Parallel-bedded Sandstone Lithofacies) 67 3.4.8.1 Palaeocurrent Indicators Associated with Lithofacies 8 69 3.4.9 Lithofacies 9 (Massive Bioturbated Sandstone Lithofacies) 70 3.4.10 Lithofacies 10 (Epsilon Cross-bedded Sandstone Bar Lithofacies) 72 3.4.11 Lithofacies 11 74 (Lenticular and Tabular Cross-bedded Sandstone Bar Lithofacies) 3.4.11.1 Palaeocurrent Indicators Associated with Lithofacies 11 77 3.4.12 Lithofacies 12 79 (Carbonaceous Mudstone Lithofacies)

3.5 Lithofacies Associations 79

3.5.1 Lithofacies Association A 80 3.5.2 Lithofacies Association B 80 3.5.3 Lithofacies Association C 81 3.5.4 Lithofacies Association D 82 3.5.5 Lithofacies Association E 83 3.5.6 Lithofacies Association F 83 3.5.7 Lithofacies Association G 84 3.5.8 Lithofacies Association H 85

Chapter 4 Systematic Palaeontology 86

Chapter 5 Discussion 101

5.1Introduction to Marginal and Shallow Marine Environments 101

5.2 Wave and Storm Dominated Shoreline 103

5.2.1 Lithofacies Association A 104 (Middle to Outer Shelf/Offshore) 5.2.2 Lithofacies Association B 105 (Lower Offshore Transition Zone to Upper distal Lower Shoreface) 5.2.3 Lithofacies Association C 106 (proximal Lower Shoreface) 5.2.4 Lithofacies Association D 111 (Upper Shoreface –Beach Complex)

5.3 Wave and Tide Influenced Estuary 113

5.3.1 Lithofacies Association E 113 (Channelised Tidal Flat) 5.3.2 Lithofacies Association F 114 (Lagoon to Washover Flat) 5.3.3 Lithofacies Association G 116 (Estuary Bay to Subaqueous Bayhead Estuary) 5.3.4 Lithofacies Association H 117 (Subaerial Bayhead Estuary)

5.4 Palaeoflow directions, progradational sense and basin morphology 118 for the Gydo-Gamka and Voorstehoek-Hex River Systems

5.4.1 Shoreline orientation data from Lithofacies Association B and C 118 5.4.1.1 Shoreline orientation from wave ripple crest trends 118 5.4.1.2 Shoreline orientation from parting lineation trends 119 5.4.1.3 Shoreline orientation from asymmetric current 119 X

ripples 5.4.2 Shoreline orientation data from Lithofacies Association D 120 5.4.2.1 Shoreline orientation from cross bed lee slope dip directions 120 5.4.2.2 Shoreline orientation from asymmetric current Ripples 120 5.4.2.3 Shoreline orientation from parting lineation trends 120 5.4.3 Combined approach to shoreline orientation and progradational sense of the Gydo – Gamka and Voorstehoek – Hex River Systems 121 5.4.3.1 Northeast – southwest striking shoreline model 121 (Fig. 5.18) 5.4.3.2 Northwest - southeast striking shoreline model 121 (Fig. 5.19)

5.5 Palaeoflow directions, progradational sense and basin morphology 122 for the Tra-Tra – Boplaas System

5.5.1 Shoreline orientation from lee slope dip direction of downstream 123 accretion bars 5.5.2 Shoreline orientation from apparent gutter trends 123 5.5.3 Shoreline orientation from rib and furrow structure trends 123 5.5.4 Shoreline orientation from asymmetric current ripples 123

5.6. Palaeontology 124

5.6.1 Preservation and Taphonomy 124 5.6.2 Palaeo - ecology and Biostratigraphy 126

5.7 Sequence Stratigraphy and Basinal Correlation 127

5.7.1 Introduction to Sequence Stratigraphy 127 5.7.2 Sequence stratigraphy of the Ceres Subgroup 130 5.7.3 Towards a Devonian Sea Level Curve for South Africa 132

Chapter 6 Conclusions 133

References 137

Chapter 1: Introduction

The Devonian Period (416.0 ± 2.8 – 359.2 ± 2.5 Ma) (Gradstein, et al., 2004) spans a critical time in Earth’s history. Two seemingly independent and major evolutionary radiations within the Eukaryota occurred during this Period. Broadly, the Devonian world consisted of two large supercontinents, Euramerica to the north and Gondwana to the south, separated by the Iapetus Sea (House and Gradstein, 2004). The Devonian was a Period of inferred global sea level high (House and Gradstein, 2004).

On land, the first true terrestrial environments and their accompanying ecosystems evolved via the en masse colonisation and diversification of terrestrial embryophytes the so called ‘Devonian Explosion’ (White, 1986, Thomas and Spicer, 1987, Edwards, 1990, Anderson, et al., 1999c, Kenrick and Davies, 2004). During the Devonian Period, vascular embryophytes diversified from their simple, low lying and herbaceous Late Silurian antecessors, the rhyniopsids and zosterophyllopsids (White, 1986, Thomas and Spicer, 1987, Anderson, et al., 1999c, Kenrick and Davies, 2004). The Devonian fossil record tentatively demonstrates that the zosterophyllopsids gave rise to the lycopods (White, 1986, Thomas and Spicer, 1987, Anderson, et al., 1999c, Kenrick and Davies, 2004). Rhyniopsids are thought to have given rise to all other higher vascular embryophytes and include: progymnosperms, pteridophytes and sphenopsids by the turn of the Period (White, 1986, Thomas and Spicer, 1987, Anderson, et al., 1999c, Kenrick and Davies, 2004). Some pteridophytes and lycopods reached arborsecent proportions and are thought to have been instrumental in the afforestation of these earliest terrestrial landscapes (Scheckler, 2003).

From a geomorphological perspective, embryophytes have been instrumental in shaping terrestrial environments owing to their massive biological contribution to geomorphological processes on continental landscapes that were once only governed by physical and chemical weathering (Davies and Gibling, 2010). Such biogeomorphological activities that embryophytes have been, and are, responsible for include: pedogenesis, by providing a means to stabilise unconsolidated sediment

forming soils, via root systems as well as adding to the nutrient content of soils (i.e. carbon (humic) content via plant decomposition) as well as regulating salt content in soils (Scheckler, 2003, Kenrick and Davies, 2004, Huggett, 2007). As weathering agents, the mechanical action of roots may create pore spaces in soils (encouraging aeration of soils) or expand joints and bedding planes in rocks (encouraging chemical weathering processes) (Huggett, 2007). Humic and fulvic acids, produced by plant decay, encourages chelation in soils and promotes metal transfer between rock and soil via creating organic matter-metal complexes, and thus too chemically weather and alter rock (Kenrick and Davies, 2004, Huggett, 2007). Stabilising hillslope, coastal, alluvial and fluvial landforms, by decreasing the amount of sediment that would otherwise be lost to sheet flow run off, or gravitational (mass movement) processes via roots or rhizoids (Burgoyne, et al., 2005, Huggett, 2007, Davies and Gibling, 2010). This encouraged a change from previous braidplain, sand sheet wash fluvial systems and aeolian environments to meandering fluvial systems with channelised beds and the development of muddy floodplains (Davies and Gibling, 2010).

Terrestrial embryophytes thus created new ecosystems and niches that seemingly encouraged concomitant colonisation by the earliest evolving terrestrial metazoans, and thus, the start of complex terrestrial food webs between plants and animals (Selden, 2003). Body fossils, as well as ichnofossils indicate that by the Late Devonian to Early Carboniferous, fully terrestrial adapted arthropods (arachnids, isopods, myriapods and hexapods had already made their appearance from the Late Silurian, but become more numerous in the Devonian), as well as tetrapods, i.e. amphibians (progenitors of the later reptilian amniotes), were possibly venturing onto land (Clack, 1997, Selden, 2003, Benton, 2005, Grimaldi and Engel, 2005). Thus, by the end of the Devonian Period, the first forests and terrestrial ecosystems were already entrenched (Scheckler, 2003, House and Gradstein, 2005).

The Devonian Period has been referred to as the “Age of the Fish”. In the oceans, fish grade gnathostomes showed the greatest diversity, with derived forms existing concomitantly with the “agnatha” (‘jawless fish’) and with other primitive gnathostome classes. Among the ancient gnathostome fish classes that would reach their maximum diversity in the Devonian were the placoderms and acanthodians (Benton,

2005, Kardong, 2008). Both classes diversified to include, as of then, unoccupied ecological niches, including freshwater habit forms, bottom dwellers, filter feeding and active predatory forms (Benton, 2005, Kardong, 2008). Placoderms have their genesis in the Late Silurian and are considered to be basal gnathostomes and may represent a separate radiation within the Gnathostomata, their relationship to other gnathostomes and the agnatha is poorly understood (Benton, 2005, Kardong, 2008). Acanthodians have their genesis in the Ordovician and share several cranial and skeletal synapomorphies with the osteichthyes (‘bony fish’) and may be united with the Tetrapoda within a larger group, the Teleostomi (Benton, 2005, Kardong, 2008). Alternatively, recent cladistic analysis suggests that acanthodians may be stem group chondrichthyians or stem group gnathostomes based on shared ‘primitive’ cranial synapomorphies shared with chondrichthyians (Davis, et al., 2012).

These orders existed alongside the earliest actinopterygiian (‘ray finned fish’) fish, the palaeonisciformes and sarcopterygii (‘fleshy finned fish’) fish, together grouped as the osteichthyes (‘bony fish’) as well as true sharks and chimaeras (chondrichthyes: elasmobranchii and holocephali respectively) (Holmer, 2000, Benton, 2005, Kardong, 2008, Elliot, et al., 2010, Gess, 2011). Ammonoid cephalopods too are widely considered to have made their first appearance during the Devonian Period, and are thought to have originated from a Silurian Bactritid ancestor (Holland, 2003).

By the end of the Devonian Period, the “ostradcodermi” sensu lato which includes the Pteraspidomorpha, Thelodonts Anaspids and Osteostracans would go entirely extinct, whilst the Conodonta and Cylclostomata: Myxinoidea (hagfishes) and Hyperoartia (lampreys) persisted (Benton, 2005, Kardong, 2008). Placoderms too go entirely extinct at the end of the Devonian (Benton, 2005, Kardong, 2008).

The Bokkeveld Group (< 407.0 ± 2.8 - > 385.3 ± 2.6 Ma.) is an Early to Middle Devonian sedimentary rock succession of the Cape Supergroup comprising alternating mudstone, siltstone and sandstone lithologies that crop out over a substantial area of the Eastern and Western Cape Provinces of South Africa (Theron and Johnson, 1991). The rocks of the Bokkeveld Group contain a rich fossil fauna of endemic south-western Gondwana gnathostome fish and Malvinokaffric Realm

marine invertebrates as well as some of the earliest known terrestrial embryophytes (Seward, 1932, Plumstead, 1969, Boucot, et al., 1983, Oosthuizen, 1984, Boucot, et al., 1986, Anderson, et al., 1999a,b,c, Boucot, 1999).

Rocks of the Bokkeveld Group record the last major (2nd order) marine transgression into the Cape Basin (Cooper, 1986, Cloetingh, et al., 1992, Tankard, et al., 2009). The Group comprises rocks that record sedimentation within a series of up to 6 coarsening upward megacycles, the last of which continues into the overlying Witteberg Group (Theron, 1972, Johnson, et al., 2006, Theron, pers comms, 2012). Each coarsening upward megacycle represents smaller (3rd order) marine transgressions and regressions within the major transgressive event (Theron, 1972, Cooper, 1986, Cloetingh, et al., 1992, Tankard, et al., 2009). The mechanism controlling the initiation of both 2nd and 3rd order transgressions is unknown; however, autocyclic and allocyclic models have been suggested by Cooper (1986) and Theron (1972), Tankard, et al. (1982), Theron and Loock, (1988), Tankard, et al. (2009) respectively. Each coarsening upward megacycle is thought to represent sedimentation as a series of wave/storm dominated deltas (Theron, 1972, Tankard and Barwis, 1982). In outcrop, lithologies of the Bokkeveld Group are present as continuous arenaceous and argillaceous units that form the basis for formational scale stratigraphic division of the Group (Theron, 1972, Theron and Johnson, 1991). The lateral continuity of these sedimentary bodies has been explained by secondary reworking and amalgamation of sediment at time of depostion by wave/storm action (Theron, 1972, Tankard and Barwis, 1982).

Considering the aerial extent of the Bokkeveld Group, relatively little research has been undertaken on the sedimentology, stratigraphy and palaeontology of this very important aspect of the South African fossil record. In depth sedimentological studies and palaeoenvironmental reconstructions on the Bokkeveld Group have been performed by Theron (1972) and Tankard and Barwis (1982) in addition to palaeoecologic reconstructions by Hiller and Theron (1988). More recent work by Fourie (2010) focused on the provenance of clastic material and tectonic setting for the Bokkeveld Group whilst under active sedimentation assuming current accepted models by previous workers. Re-interpretation of the palaeoenvironments of the Bokkeveld Group, as well as the use of sequence stratigraphic methods, coupled

with palaeontological evidence, may effectively be used to track basin-wide changes that occurred during deposition of the Bokkeveld Group, leading up into the Witteberg Group.

This study is restricted to the first three upward coarsening megacycles (Gydo- Gamka, Voorstehoek-Hex Rivier and Tra-Tra- Boplaas Systems) of the Ceres Subgroup, which were deposited from Emsian to Eifelian time, based on available palaeontological data (Theron, 1972, Theron and Johnson, 1991). These rocks have been studied within the north-south trending limb of the Cape Fold Belt, the Cederberg, Western Cape, South Africa. Here, a reassessment of the sedimentology and an intensive facies analysis has been conducted in order to determine the most likely palaeoenvironment (s) of deposition as well as whether any degree of cyclicity exists between the two coarsening upward megacycles. Cyclicity has been addressed by placing sequences and parasequences within the upward coarsening megacycles within a sequence stratigraphic framework. The stratigraphic ranges and distributions of fossil fauna within each megacycle have been taken into account in order to constrain palaeoenvironments. An attempt at erecting a biostratigraphic and ecostratigraphic framework has been attempted in this study to discuss whether any controls on fossil assemblages existed, i.e. if these controls were temporal or spatial.

1.1 Palaeotectonic Setting of the Cape Basin and the Cape Supergroup

The Cape Basin houses the 6 – 10 km thick sedimentary succession of the Cape Supergroup which was deposited in the Aghulas Sea under a divergent tectonic regime over a near continuous period of ~ 160 - 170 million years, extending from the Early Ordovician to the Early Carboniferous (Broquet, 1992, Cloetingh, et al., 1992, McCarthy and Rubidge, 2005, Thamm and Johnson, 2006, Tankard, et al., 2009).

The Cape Supergroup consists of, in stratigraphic order, the Table Mountain Group (Early Ordovician to Earl Devonian), the Bokkeveld Group (Early to Middle

Devonian), the Witteberg Group and its tentative lateral correlate, the Msikaba Formation (Middle Devonian to Early Carboniferous) (Thamm and Johnson, 2006, Kingsley and Marshall, 2009). Outcrop of the Table Mountain, Bokkeveld and Witteberg Groups is present along the entire extent of the Cape Fold Belt (CFB) and covers some 200,000 km2 (Fig. 1.1) (Thamm and Johnson, 2006, Tankard, et al., 2009). The Cape Supergroup is estimated to cover an offshore area of at least 100 km offshore (Tankard, et al. 2009). Subcrop of the Cape Supergroup has been demonstrated to extend northwards underneath the Karoo Supergroup from exploration borehole and teleseismic data (Fig. 1.1) (Boucot, et al., 1986, Tankard, et al., 2009). The Msikaba Formation appears to be separate from rocks of the main Cape Supergroup within the CFB and crops out in the Eastern Cape and Kwa-Zulu Natal Provinces, roughly between the towns Port Shepstone and Port St. Johns (Kingsley and Marshall, 2009).

1.1.1 Base of the Cape Basin

The Cape Basin is floored by several lithostratigraphic entities. The Cape Supergroup largely unconformably overlies the Namaqua-Natal and Saldanian mobile belts (Thamm and Johnson, 2006, Tankard, et al. 2009, Booth, 2011). These mobile belts include various metavolcano-sedimentary and associated igneous intrusive rocks of the Mesoproterozoic Namaqua Natal Belt and the Pan-African (Neoproterozoic – Early to Middle Cambrian) Gariep and Saldania Belts (Broquet, 1992, Gresse, et al.., 2006, Thamm and Johnson, 2006, Tankard, et al., 2009). The topmost Vanrhynsdorp Group (Ediacaran – Early Cambrian) is truncated by the Cape Supergroup (Gresse, et al., 2006). The exact geology of these terranes is beyond the scope of this study and shall not be discussed.

Lowermost successions of the Cape Supergroup (Piekenierskloof Formation: Table Mountain Group) were deposited within a series of syn-rift grabens where these deposits are conformable with the underlying Klipheuwel Group (Fig. 1.2) (Broquet, 1992, Gresse, et al., 2006, Thamm and Johnson, 2006, Tankard, et al., 2009). Rocks of the Klipheuwel Group are thought to represent deposition within an earlier rift, possibly related to the Cape Basin

(Tankard, et al., 1982) or a molasse related to Pan-African orogenesis (Gresse, et al., 2006). The Msikaba Formation demonstrably nonconformably overlies the Namaqua Natal Province in its southern most outcrop area (Kingsley and Marshall, 2009). To the North of its outcrop area (from Hibberdene northwards) the Msikaba Formation unconcofmably overlies the Natal Group (Kingsley and Marshall, 2009).

Accommodation of the Cape Basin was partly due to pre-existing structural features present within the terranes over which the Basin formed. These structural features are present in both the Namaqua-Natal and the Saldanian Mobile Belts. From geophysical data, Tankard, et al. (2009) noted a general northward dipping lineament fabric which they attributed to be intra-plate sutures between the Saldanian and Namaqua-Natal terranes (Fig. 1.3). Tankard, et al. (2009) demonstrates that the formation of the Saldanian Orogen by dextral transpression created a series of northwest-southeast trending shears and faults with dextral stepping and offset between the Rio de la Plata craton and the Namaqua Natal Mobile Belt (Fig. 1.3). Post Saldanian orogenic collapse by brittle failure along these northwest-southeast trending faults and shears of the Saldanian Orogen created a series of rift basins that accommodated syn-rift sediments of what was to become the Klipheuwel Group and the Piekenierskloof Formation, thus ushering in the birth of the Cape Basin (Tankard, et al., 2009).

1.1.2 Subsidence and Basin Analysis of the Cape Basin

Few attempts have been made in the past at calculating subsidence rates and basin histories for the Cape Basin. Shone and Booth (2005) have outlined several factors hindering attempts at establishing a basin history for the Cape Basin

Earliest subsidence estimates for the Cape Basin were calculated by Theron (1972), however, these subsidence rates were calculated only for the time of Bokkeveld sedimentation with loose age constraints based on palaeontological data. Theron (1972) estimated a subsidence estimate of 2.54

cm.1,000 yrs-1 for the Clanwilliam Sub-basin and 6.63 cm. 1,000 yrs-1 for the Agulhas Sub-basin. It was assumed that sedimentation within the Bokkeveld Group was continuous and extended from Early to Late Devonian time (~ 60 Ma.). Alternatively, Theron (1972) considered that the Bokkeveld Group could possibly be more restricted to the Early Devonian based on the presence of Emsian brachiopods. According to Theron (1972) this would imply that sedimentation was continuous over a period of 15Ma. Subsidence rates were recomputed to be 10.16 cm.1,000 yrs-1 and26.50 cm. 1,000 yrs-1 for the Clanwilliam and Agulhas Sub-basins respectively. All subsidence estimates further, were calculated under the assumption that measured thicknesses for the Group were correct. These were ~ 1524 m for the Clanwilliam Sub-basin and ~3962.4 m for the Agulhas Sub-basin. Theron (1972) also mentions a discrete difference between the Table Mountain and Bokkeveld Group subsidence rates in that the subsidence rate for the Table Mountain Group, based on the Nardouw rate of subsidence, was ~3.05 cm. 1,000 yrs-1.

Later work by Rust (1973) made assumptions that sedimentation within the Cape Basin was continuous and uninterrupted by any major regional unconformities. It was assumed that early accommodation and sedimentation in the Cape Basin (i.e. Klipheuwel Group and Piekenierskloof-Graafwater Formation sedimentation) was strongly aided by brittle deformation and rifting and thought to represent terrestrial rift valley alluvial plain type environments. The Peninsula Formation to Nardouw Group sedimentation represents a period of downwarping followed by tectonic quiescence. The rocks of the Bokkeveld Group rare combined to epresent a renewed period of tectonic activity with alternate periods of quiescence creating periods of repeated transgressive and regressive sequences within an overall shallowing upward marginal marine (deltaic) environment until the time of Witteberg Group.

Cloetingh et al. (1992) focused on quantifying subsidence rates of the Cape and Karoo Basins by backstripping; this has partially revealed the subsidence and thermal history of both basins during the time of Cape and Karoo Supergroup deposition (Fig. 1.4). This method calculates the amount of decompaction of a given stratum by removing strata up sequence from it

(Cloetingh, et al., 1992, Allen and Allen, 2005c). Variables considered when calculating decompaction are standardised porosity and depth (compaction) relations for specific lithologies whilst assuming homogenous isostatic behaviour of the lithosphere (Cloetingh, et al., 1992, Allen and Allen, 2005c). Cloetingh, et al. (1992), however, did not account for allocyclic controls such as glacioeustacy in their calculations.

Only the subsidence history of the Cape Basin will be discussed since it is within the scope of this project. Basin subsidence was not uniform across the Cape Basin at the time of sedimentation, possibly as a result of a, then, active tectonic divide, the so-called “Syntaxial Domian”, which was persistent from the start of Cape Supergroup sedimentation within the Cape Basin (Fig. 1.4) (Theron, 1972, Cloetingh, et al., 1992). Two broad sub-basins have been recognised to have been active during sedimentation of the Cape Supergroup, based on differences in accumulated sedimentary packages and differences in assumed correlative strata (Theron, 1972, Cloetingh, et al., 1992, Thamm and Johnson, 2006) (Fig. 1.4). Differences in the nomenclature used for these two sub-basins exist. Cloetingh referred to these basins in terms of their region and thus used the terms “Western Cape Province” and “Southeastern Cape Province” for their respective western and eastern sub- basins of the Cape Basin. Theron (1972) used the terms “Clanwilliam Basin” and “Agulhas Basin” for the respective western and eastern sub-basins of the Cape Basin. For both cases, the division between the sub-basins is drawn at ~ 20º E, the occurrence of the Syntaxial Domain (Fig. 1.4). With the creation of new provincial boundaries in South Africa in 1994 (Chapter 6, Constitution of the Republic of South Africa, 1996), the nomenclature of Cloetingh, et al. (1992) could lead to confusion. The author therefore favours nomenclature originally used by Theron (1972), it is suggested that these basins should be called the “Clanwilliam Sub-basin and the Agulhas Sub-basin”. Reported average thicknesses for the Cape Supergroup, according to Thamm and Johnson (2006) are in the order of 6,925 m for the Clanwilliam Sub-basin and 11,316 m for the Agulhas Sub-basin. A geochemically distinct third “southern basin” sensu Fourie (2010) has been hinted at in the vicinity of the syntaxial

9 domain, but this may reflect a difference in sedimentary sourcing and not be an actual separate basin.

Two discreet episodes of rapid tectonic subsidence were noted by Cloetingh, et al. (1992) within the Cape Basin (Fig. 1.5). The first of which is associated with the initial opening of the Cape Basin and extends from the time of initial Table Mountain Group (~ 464 Ma.) deposition to the base of the Nardouw Subgroup (Piekenierskloof to Cedarberg Formation interval) (~ 430 Ma.) (Fig. 1.5) (Cloetingh, et al., 1992). This initial rapid phase of tectonic subsidence was followed by a period of decelerated subsidence and quiescence until the start of the Bokkeveld Group (~ 390 Ma.) (Fig. 1.5) (Cloetingh, et al., 1992). The Bokkeveld Group represents the second phase of rapid, but constant tectonic subsidence and extending until the end of Bokkeveld Group sedimentation (~ 375 Ma.) (Fig. 1.5) (Cloetingh, et al., 1992). Deposits of the overlying Witteberg Group record a period of decelerated tectonic subsidence until the closure of the Cape Basin, the onset of the Cape Fold Belt Orogeny and creation of the Karoo Basin (Fig. 1.5).

More recent subsidence estimates for the Cape Basin have been calculated by Tankard, et al. (2009). Tankard, et al. (2009) has also attempted to relate accommodation and subsidence generation episodes within the Cape Basin to tectonic events associated with convergence and flat plate subduction between the palaeo-Pacific and Gondwanan Plates along the south western coastline of Gondwana, what is now northwest Argentina, during two discreet episodes associated within the Early Ordovician to Early Carboniferous Famatinian Orogenic Cycle in excess of 1000 km from the Cape Basin (Fig. 1.6).

The authors suggest a subsidence rate of 12 cm. 1000-1 yr. for Graafwater to Peninsula time sedimentation. This records the initial major (1st order) transgression and opening of the Cape Basin. The predominance of quartz sandstones within this interval has been suggested to represent base level basinal fill and a balance between accumulation and sedimentation rates. Tankard, et al., (2009) related the opening of the Cape Basin to subduction

10 and collision of the Precordillera terrane to south western Gondwana. The glacigene Pakhuis and its proglacial outwash Cedarberg Formation to Nardouw Subgroup sedimentation is thought to represent a period of basin stability and lower subsidence rates alluding to cessation in subduction along south western Gondwana and estimated to be 3 cm. 1000-1 yr. Bokkeveld Group – Weltevrede Subgroup (Witteberg Group) sedimentation occurred during a more tectonically active phase of the Cape Basin with subsidence rates calculated to have been between 18 - 21 cm. 1000-1 yr. during Bokkeveld Group sedimentation, decreasing to 4 - 7 cm. 1000-1 yr. during Weltevrede Subgroup sedimentation. This is thought to be related to magmatism and migmatisation associated with the Chanic Orogeny in Northwest Patagonia during Bokkeveld Group sedimentation with Weltevrede Subgroup sedimentation representative of tectonic quiescence after the Chanic Orogenic event. The Witpoort Formation and overlying Lake Mentz and Kommadagga Subgroups possibly represents basin stability and reworking with tectonic quiescence along the southwestern coast.

Booth (2011) has cautioned that thicknesses of certain units, the Peninsula Formation in particular, are dubious since thrust stacking may not have been taken into account when sections were measured, thus leading to misleading subsidence values.

1.1.3. Models for the Tectonic Setting of the Cape Basin

Although, a marginal (passive continental margin) “Atlantic – Type” basin (in earlier literature referred to as: “miogeoyncline”) is envisaged and favoured as likely candidate for the Cape Basin (Theron, 1972, Cloetingh, et al., 1992, Shone and Booth, 2005, Thamm and Johnson, 2006), other basin models have also been proposed. All models favour marginal sedimentation in a convergent tectonic setting via continental lithospheric extension of the Gondwana Plate as the mode for basin development via various mechanisms. All models agree that sedimentation within the Cape Basin commenced in the Early Ordovician and ceased in the Early Carboniferous.

1.1.3.1 Back-arc Basin Model (Catuneanu, et al., (1998))

A back-arc basin, under plate convergence, has been tentatively illustrated, although not implicitly stated by Catuneanu et al. (1998) as a potential Karoo Basin precursor. This model assumes that the Cape Basin formed due to extension in the over-riding Gondwana Plate in a convergent plate setting, due to subduction of the palaeo-Pacific Plate underneath the Gondwana Plate during the formation of the Pan Gondwanan Mobile Belt and the creation of the retro-arc Karoo foreland basin (Fig. 1.7). As illustrated by Catuneanu, et al. (1998), the locus for extension was located on a zone of weakness, the pre- existing Pan-African Belt. The particular Pan-African Belt in question appears be the late Neoproterozoic – early to middle Cambrian Saldanian Orogen that underlies the Cape Supergroup (Tankard, et al., 2009). Extension in the over-riding Gondwana Plate, although not stated by Catuneanu, et al. (1998), may have been aided by plate edge forces associated with subduction, viz. slab roll back, mantle flow, steep angle of subducting plate (Allen and Allen, 2005a, Nichols, 2009).

Sedimentary profiles for back-arc basins are variable and loosely based on the lithosphere they are floored by, i.e. oceanic or continental, as well as the age and angle of the subducting plate (Allen and Allen, 2005a and b). These factors govern whether the back-arc basin is under net compression or extension (Allen and Allen, 2005a and b). Typical back-arc basin sedimentary profiles floored by oceanic lithosphere appear to be under-filled, and are characterised by marginal to shallow marine facies that grade into more dominant deep marine sedimentary facies generally (Ingersoll, 1988. Allen and Allen, 2005b, Nichols, 2009). Proximal volcano-clastics may be present as wedges that inter-finger with more distal finer grained clastic deposits derived from the arc source area (Ingersoll, 1988, Allen and Allen, 2005b, Nichols, 2009). Volcano-clastic deposits are common in early stages of back-arc basin development that gradually give way to more

pelagic to hemi-pelagic clastic deposits with basin maturity (Ingersoll, 1988, Allen and Allen, 2005b). Sediment is derived from the volcanic arc and the continental lithosphere (Allen and Allen, 2005b, Nichols, 2009) If extension is great enough, such that new oceanic lithosphere is created from an active rifting centre, new oceanic crust may be generated, thus ophiolites may be present in consolidated profiles (Ingersoll, 1988, Allen and Allen, 2005b, Nichols, 2009). Back-arc basins are generally short lived (Allen and Allen, 2005b, Nichols, 2009). Nichols (2009) has placed an estimate of ~ 20 Ma. for the Cenozoic Western Pacific back-arc basins between development and abandonment.

The back-arc basin model for the Cape Basin seems unlikely. Although, deposits of the Cape Basin reflect times of marginal, shallow and deep marine facies depositional environments, neither volcanoclasitites, nor ophiolites are present. The Cape Basin appears to have been active for 160 – 170 Ma. and thus dwarfs current estimates for back-arc basin longevity stated by Nichols (2009).

1.1.3.2 Margin-sag-Interior-sag basin Model (Broquet (1992)

Broquet (1992), sensu Kingston, et al. (1983): Global Basin Classification System, suggested a margin-sag-interior-sag (MSIS) basin model for the Cape Basin in dismissal of both the passive margin “Atlantic – Type” model of Theron (1972), Winter (1984, 1989) and Johnson (1991) and the aulacogen model of Tankard, et al. (1982). Broquet (1992), p. 170 feels that both models are simplistic since “...they take no cognizance of, or simply assume, the basin configuration, architecture (sequence stratigraphy) and conditions of sedimentation.” The margin-sag-interior-sag basin model sensu Kingston, et al. (1983) which Broquet (1992) presents is based on the subsidence curve values presented by Cloetingh, et al. (1992) (Fig. 1.5). This curve shows a punctuated, non exponential subsidence curve for Cape Basin. Broquet (1992) states that typical passive

margin basins would show an even and non punctuated exponential decrease in (thermal) subsidence with decaying heatflow. According to the basin classification scheme as presented in Kingston, et al. (1983), passive margin basins would be classified as continental-margin-sag (CMS) basins.

Broquet (1992) failed to explicitly explain how the Cape Basin may fit within this classification scheme. According to Kingston et al. (1983), MSIS basins consist of an a marginal marine basin (CMS) and an interior sag basin (alternatively an intracratonic sag) and thus form within the interior of continental margins and at continental margins where they are connected to a pre-existing ocean (Kingston, et al., 1983). Kingston et al. (1983), p. 2181 summarises MSIS basin deposits as “broad, gently dipping continental with sheet sands and carbonate deposits more closely comparable to those of interior basins than those of present day narrow continental margins”. Expected deposits thus resemble those of marginal marine systems. Kingston, et al. (1983) also state that no structural or stratigraphic evidence may exist for the other side of a given basin, due to the expected basinward fining of marginal marine systems. In essence, an MSIS basin would, therefore most closely would resemble an epicontinental/eperic sea connected to an open ocean. Broquet (1992) does mention, from isopach maps presented in Theron (1972), that the Aghulas Sub-basin would appear to be connected to the outboard palaeo-Pacific Ocean and that (although not explicitly stated) the Clanwilliam Sub-basin was bound by highboard non-orogenic highland source areas to the north and west. This may suggest that the Clanwilliam Sub-basin represents the Interior sag portion and the Aghulas Sub-basin represents the Margin sag portion.

Theron (1972 p. 136 - 137) favoured such a palaeogeogeographic construction of the Cape Basin, most certainly during Bokkeveld Group sedimentation, based on differences in sedimentary package thicknesses and palaeoflow datum. He envisioned the Clanwilliam

Sub-basin to be a semi restricted small, shallow marine basin situated between or within continental margins which he coined as a “Mediterranean Type Geosyncline” whilst the Aghulas Sub-basin represented a more open marine passive margin, continental shelf type basin, which he coined as an “Atlantic Type Geosyncline”.

1.1.3.3 Epistural Aulacogen Basin Model (Tankard, et al., (1982))

Tankard et. al. (1982) presented an aulacogenal model for the Cape Basin. This model favours incipient rifting between the triple junction of the African, South American and Antarctic Plates (Fig 1.8). This model also assumes major fault controlled subsidence of basement rocks as an accommodation space generating mechanism during Table Mountain Group sedimentation, followed by tectonically controlled downwarping and accommodation space creation during Bokkeveld Group sedimentation and relaxation during Witteberg Group sedimentation. Incipient rifts trended sub-parallel to the current South- African coastline and formed the axis for southward directed downwarp of the continental lithosphere. The mechanism for, and extent of rifting has not been addressed by Tankard, et al. (1982) as it is assumed that post-Gondwana rifting has removed all evidence for the aulacogen.

1.1.3.4 Episutural Sag Basin Model (Tankard, et al., (2009))

Tankard, et al. (2009) noted a paucity in brittle failure and fault assisted subsidence in generating accommodation space within the Cape Basin due to the absence of rapid lithological, thickness and facies changes. This model suggests that subsidence within the Cape Basin is attributed to intra-crustal detachment and mantle flow extension along as a result of low angle subduction at the Famatinian orogenic front. This caused tilting of the Gondwanan plate, and transgression into the Cape Basin. Lithospheric deflection and recovery followed after subduction ceased, over a long time period. The authors suggested that flat plate subduction is responsible for the initiation of the Cape

Basin following the Mitroveca et. al. (1989) school of thought. This model suggests that during tectonically active stages at convergent zones, flat plate subduction and mantle flow beneath the subduction zone can cause the lithospheric crust to tilt and allow for a tilt induced transgression up to 1000 km from the subduction zone. During tectonic quiescent stages, it may take up to 25 Ma for the lithospheric crust to deflect. Details pertaining to specific phases of sedimentation within the Cape Supergroup have been discussed in 1.1.2, pg.9.

1.2 Structural Features Associated with the Cape Fold Belt

A change to a convergent tectonic regime with the subduction of the palaeo-Pacific Plate beneath the Gondwana Plate during the Late Carboniferous to Early Permian resulted in the Cape Supergroup being deformed as a series of thrusts and folds along its entire extent to form the CFB (alternatively, Cape Thrust Belt) and its successor, the Karoo Foreland Basin (de Beer, 1995, Catuneanu, et al., 1998, Bumby and Guirard, 2005, Booth, 2011). The CFB is an intraplate fold belt and part of the larger Gondwanide Orogen (aka. “The Gondwanides”) present along the former southern margin of Gondwana (Johnston, 2000, Vaughn and Pankhurst, 2008). Johnston (2000) suggested that far field convergent forces related to subduction may have been responsible for the creation of the CFB due to its apparent far distance (in excess of 1500 km) from the southern convergent margin of Gondwana during the formation of the Gondwanides. Lock (1980) suggested a flat plate subduction model that assumes that shortening was not taken up along the Gondwanan margin, but rather within the Gondwanan Plate. Other models for the mechanism by which the CFB formed at a distance from the convergent margin of southern Gondwana are beyond the scope of this study, but are discussed and referenced by Johnston (2000). Metamorphism associated with tectonism of the CFB has been demonstrated by Frimmel, et al (2001) to not exceed lowermost greenschist facies (no greater than 300º C).

The CFB is a notable region of the Gondwanides in that it is anomalous with respect to other regions of the Orogen. These include: An apparent distance in excess of

1500 km from the southern convergent margin of Gondwana at the time. The lack of intrusive igneous bodies, low metamorphic grade. Presence of two regions of fold interference and convergence around central points, i.e. the northeast-southwest striking Syntaxial Domain around Ceres and the northwest-southeast striking Port Elizabeth Antitaxis where the CFB appears to notably bend in different directions and differ in deformational intensity. These anomalies have been discussed at length by: de Beer (1995), Johnston (2000), Booth (2011). A distinct schism, in terms of deformation is notable in the CFB, around the Syntaxial Domain and the Port Elizabeth Antitaxis. With respect to the Syntaxial Domian, the CFB may be divided into a western and eastern limb (Fig.1.9). Here, the western limb of the CFB is deformed as a series of primary open synclines and anticlines, monoclines and upright folds with rough northwest-southeast trending axial fold traces and minor folds and thrusts (de Beer, 1995, Johnston, 2000, Booth, 2011). The eastern limb of the CFB shows a greater degree of deformation with the generation of northward verging recumbent primary folds with east-west trending axial fold traces and a higher degree of secondary folds and thrusts (de Beer, 1995, Johnston, 2000, Booth, 2011).

The coincidence of sedimentary strata striking at the same attitude as that of the fold axial traces, has led some workers to believe that the overall shape of the CFB mimics the South African coastline whilst the Cape Basin was active (Johnston, 2000, Booth, 2011). According to this model, the Syntaxial Domain may represent an outboard topographic high region of the Namaqua Natal Belt along episutural plate junctions during extension and the formation of the Cape Basin and was an active feature influencing sedimentation of the Cape Supergroup (Cloetingh, et al., 1992, Johnston, 2000, Booth, 2011). This model for the morphology of the Cape Basin appears to favour the deformation and fold interference model about the Syntaxial Domain of de Beer (1995).

de Beer (1995) proposes that the western limb of the CFB records deformation by an eastward directed palaeostress, whilst the eastern limb records a northward directed palaeostress of greater magnitude than that which affected the western limb. Deformation was simultaneous and interfered at the Syntaxial Domain forming the Ceres Arc, Stettyns and Hex Rivier Anticline (de Beer, 1995) (Fig.1.9). An alternate

theory for the Syntaxial Domain, Port Elizabeth Antitaxis and shape of the CFB was proposed by Johnston (2000). This theory proposes a dextral transpressional tectonic regime, based on similar dextral transform structural features found in the western domain of the CFB and those in South American and Antarctic regions of the Gondwanides. Tankard, et al. (2009) support the transpressional model of Johnston (2000), however, favouring a sinistral movement with compression. They refer to the presence of flower structures, en echelon folds and faults, fault bends, strike slip faults and vertically displaced blocks, all of which are typically seen in transpressional tectonic settings, for the CFB (Fig.1.10). Left lateral movement along the Cango and Worcester Faults are the key criterion for sinistral transpression as presented by Tankard, et al. (2009). According to Johnston (2000), the Syntaxial Domain and Port Elizabeth Antitaxis are oroclinal bends that formed by dextral directed transpression. Booth (2011) has suggested that the Port Elizabeth Antitaxis may be wholly unrelated to the creation of the CFB and may have formed with the creation of the more recent Agulhas Falklands Fracture Zone. Arguments for and against both models are discussed at length within Johnston (2000), Tankard, et al. (2009) and Booth (2011).

From teleseismic and structural data across the CFB, northward dominant dipping lystric faults, as well as associated southward dipping antithetic faults, and grains are present in rocks of the basement forming Namaqua-Natal terrane (Fig. 1.10) (Tankard, et al., 2009, Booth, 2011). These faults have been attributed to the formation of earlier Neoproterozoic rift margins (Tankard, et al., 2009, Booth, 2011). An opposite sense of inferred fault dip is present in the Cape Supergroup in these teleseismic images and from surface thrust faults (Fig. 1.10) (Tankard, et al., 2009, Booth, 2011). The difference in tectonic fabric has been interpreted by workers to represent a decollement surface that formed with convergence during the formation of the CFB (Tankard, et al., 2009, Booth, 2011). Tankard et al. (2009) and Booth (2011) considered that the CFB represents a combination of thick and thin-skinned tectonic thrust and folds. Booth (2011) commented that the features seen in the southern branch of the CFB, i.e. primary and secondary folds, intense faults and northward verging thrust planes and folding) represent thick-skinned tectonics, whilst the western branch of the CFB with primary folds and minor secondary folds and faulting represent a thin-skinned tectonic regime.

1.3 Geology of the Bokkeveld Group

1.3.1 Palaeoenvironmental Setting of the Bokkeveld Group

The Bokkeveld Group consists of five to six conformable upward coarsening megacycles (Table 1) (Theron, 1972, Johnson, 1976, Tankard and Barwis, 1982, Cooper, 1986, Theron and Loock, 1988, Tankard, et al., 2009). The cyclical upward coarsening nature among lithologies within the Bokkeveld Group records a history of alternate transgressive and regressive eustatic regimes active during deposition (Table 1.1) (Theron, 1972, Johnson, 1976, Tankard and Barwis, 1982, Cooper, 1986, Theron and Loock, 1988, Tankard, et al., 2009). Facies architecture of individual upward coarsening megacycles within the Bokkeveld Group are assumed to be those expected in marginal marine deltaic environments (Theron, 1972). Facies analysis of a typical prograding Bokkeveld succession reveals a broad two fold partitioning of facies of basal fine grained argillaceous beds, representative of shelf to prodelta slope environments, which grade upward into increasingly coarser grained arenaceous delta plain environment top set beds deposited in a range of associated sub-environments (Theron, 1972, Tankard and Barwis, 1982, Theron and Johnson, 1991).

Transgressive deposits, i.e. delta slope and prodelta deposits are thought to have been initiated with basin deepening, via inferred extensional subsidence (mantle extension) and normal faulting sub parallel to the palaeo-depoaxis (Tankard and Barwis, 1982, Tankard, et al., 2009). Regressive deposits, i.e. delta plain deposits, are thought to have been deposited during times of tectonic quiescence after tectonic subsidence had occurred (Tankard and Barwis, 1982, Theron and Loock, 1988). Here, it is assumed that a basinal sea-level fall is expected with progradation, thus, a normal regressive sequence is expected vertically from prodelta to delta plain environments (Theron, 1972, Tankard and Barwis, 1982, Theron and Loock, 1988).

The presence of five upward coarsening megacycles within the Bokkeveld Group thus indicated that the Cape Basin, during Bokkeveld Group sedimentation, was periodically tectonically unstable and shallow (Theron, 1972, Johnson, 1976, Tankard and Barwis, 1982, Cooper, 1986, Theron and Loock, 1988). From global sea level curve data, sedimentation of the Bokkeveld Group was also influenced by a global first order eustatic high for the Devonian Period (Cooper, 1986, House and Gradstein, 2005). Tankard and Barwis (1982) assumed that these deltas were strongly influenced by mixed wave and tidal processes at the time of sedimentation. This assumption is based on the relatively thick, quartz rich and mature nature of delta front, delta plain arenites that are laterally continuous across the strike of the Bokkeveld Group. Delta front and delta plain arenites in section have been noted to have erosive contacts with underlying strata (Tankard and Barwis, 1982). This is reasoned by Tankard and Barwis (1982) to represent basin wide shore zone reworking, dominantly by longshore drift, of sediment and coalescence of lateral lobate and arcuate delta lobes during times of relative tectonic quiescence. Other lines of evidence for a wave dominated deltaic model for the Bokkeveld Group are the presence of architectural features related to beach environments. These include wash over sheet sequences and tidal channel fill features in addition to occasional arenite lenses in prodelta facies attributed to storm activity. Sedimentary structures indicative of tidal processes have been noted (Theron ,1972, Tankard and Barwis 1982) and include tidal inlet and channel deposits, as well as tidal scour and ripple marks in addition to flaser and lenticular bedding. Prograding deltas of the western Clanwilliam sub-basin are thought to have been dominated by these processes (Tankard and Barwis, 1982, Tankard, et al., 2009). Prograding deltas of the Aghulas sub-basin are thought to have been dominated by river processes where delta lobe switching was common especially during sedimentation of the Adolphspoort Formation (Theron and Loock, 1988). Tankard, et al. (1982) envisions that delta plain fluvial systems of the Bokkeveld Group were possibly braided and well developed.

Theron (1972) envisioned four large depocentres (each comprising of several arcuate deltas) to have existed within the Bokkeveld Basin at the time of

sedimentation (Fig. 1.11). The shallow Clanwilliam sub-basin housed the Hantam and Saldanha Deltas, whilst the deeper Agulhas sub-basin housed the Nuweveld and Winterberg Deltas (Theron, 1972). Palaeocurrent directions, in tandem with isopach maps, fossil distributions, southward decrease in grain size and wedging out of sandstone units have been used by Theron (1972) to indicate a dominantly southerly and westerly directed palaeoslope and a constantly active north and eastern provenance area (Bushman, Kamdeboo and Winterberg Mountainlands) (Figs. 1.11 – 1.13). Progradation of these deltas thus had a general southward to south and south-westward trend (Theron, 1972).

There was a further western provenance area (Atlantic Mountainlands) that was active during Bokkeveld Group sedimentation (Saldhana Delta) (Theron, 1972, Fourie, et al., 2011). The Saldhana Deltas area assumed to have prograded in an opposite sense to their counterparts. Progradation was directed in a westward to west-southwestward fashion (Theron, 1972). At certain times, i.e. during Hexrivier and Boplaas sedimentation (~? Middle to Late Eifelian) this western source area became more dominant (Fig. 1.12) (Theron, 1972, Theron and Loock, 1988).

Detrital zircon data presented by Fourie (2010) venerate the geographically likely provenance areas suggested by Theron (1972). Fourie (2010) suggest three areas that were most likely to supply clastic material to the Bokkeveld Group at time of sedimentation in different parts of the Cape Basin (Fig. 1.14). The Clanwilliam and Aghulas Sub-basins were supplied with sediment derived mainly from a Namaqua – Natal Belt source area (Mesoproterozoic aged) with little imputus from a Pan – African Brasiliano Belt (Neoproterozoic aged) source, whilst a third and apparently distinct “southern basin” around the syntaxial domain was majorly supplied sediment from Pan-African-Brasiliano Belt source (the Dom Feliciano Belt) and from the Mesoproterozoic Rio de la Plata Craton and Sierra Pamapas Terrane in South America.

1.3.2 Allocyclic vs. Autocyclic Models for Megacycle Cyclicity

The initiation of each coarsening upward megacycle is associated with an apparent increase in accommodation space and reduction in sediment supply in the Cape Basin at time of deposition. The mechanism by which an increase in accommodation space and was initiated and its apparent cyclicity during Bokkeveld sedimentation is not yet fully understood. Currently two models have been proposed. Broadly, these are a ‘tectonic’ allocyclic model (Theron, 1972, Tankard, et al., 1982, Theron and Loock, 1988, Tankard, et al., 2009); and a ‘eustatic’ autocyclic model (Cooper, 1986).

The ‘tectonic’ allocyclic model includes ideas for Bokkeveld accommodation space generation and cyclicity presented by Theron (1972), Tankard and Barwis (1982) and Tankard et al. (2009). Theron (1972) suggested that alternation and generation of each upward coarsening megacycle and associated delta progradation was due to periodical shelf instability of a shallow Cape Basin, or oscillatory behaviour akin to that at the margins of intracratonic basins. This initial model was modified by Tankard and Barwis (1982) and Tankard, et al. (2009) to suggest that accommodation space and sea-level rise was created in the Cape Basin via inferred extensional subsidence and lithospheric sag by intra-crustal detachment and mantle extension with little to no normal faulting sub parallel to the palaeo-depoaxis of the basin. Transgressive deposits, i.e. delta slope and prodelta deposits are thought to have been laid down during these periods (Tankard and Barwis, 1982, Tankard, et al., 2009). Regressive deposits, i.e. delta plain deposits, are thought to have been deposited during times of tectonic quiescence after tectonic subsidence had occurred (Tankard and Barwis, 1982, Theron and Loock, 1988). Here, it is assumed that a basinal sea-level fall is expected with an increase in sediment supply with progradation, thus, a normal regressive sequence is expected vertically from prodelta to delta plain environments (Theron, 1972, Tankard and Barwis, 1982, Theron and Loock, 1988).

The ‘eustatic’ autocyclic model for Bokkeveld accommodation space and generation has been suggested by Cooper (1982 and 1986). Conodont fossils have been demonstrated to show some congruence with global transgressive-regressive cycles (T-R cycles) from the Devonian Period (Cooper, 1986, Fig. 14.2, pg. 209 – 211 House and Gradstein, 2004). Cooper (1986) noticed an apparent match of conodont zones within 3rd order T-R cycles from the Devonian Period of Euramerica of the time that may be of correlative use with lithofacies data and Malvinokaffric Realm invertebrate occurrences from the Bokkeveld Group. The scheme by which Cooper (1986) correlated Devonian sequences with conodont biozone and T-R cycle data from Euramerica is summarised in Figure 1.15. Based on the main occurrence of Malvinokaffric Fauna, mainly within benthic, deep water “distal offshore” environments (Hiller and Theron, 1988), Cooper (1986) favoured two large 2nd order flooding events within the Cape Basin during Bokkeveld Group sedimentation. The en masse occurrence of Malvinokaffric Realm fauna within typical distal offshore “black shales” occurs only at the base of the Gydo and the Waboomberg Formations (i.e. at the base of both the Ceres and Bidouw Subgroups respectively). This data, according to Cooper (1986) is congruous with a change in T-R cyclicity in Euramerica which is marked by a large and abrupt deeping event (Fig. 1.15). As such, Cooper (1986) suggested that both Gondwana and Euramerica were affected by global eustatic rises and falls during the Devonian Period.

1.4 Lithostratigraphy of the Bokkeveld Group in the Western Cape Province

The Bokkeveld Group conformably overlies the Nardouw Subgroup (Rietvlei / Baviaanskloof Formations) of the Table Mountain Group (Theron, 1972, Theron and Johnson, 1991, Johnson, et al., 1999). This contact is generally sharp, but may be gradational over a few metres in places (Theron, 1972, Theron and Johnson, 1991, Johnson, et al., 1999). The base of the Bokkeveld Group is considered to be represented by the first argillaceous bed overlying the last sandstone of the Nardouw Subgroup (Rietvlei / Baviaanskloof Formations) (Theron, 1972).

The Bokkeveld Group comprises 3 subgroups, which in turn house 14 formations (Table 1) (Theron, 1972, Theron and Johnson, 1991). The Ceres and Bidouw Subgroups occur only west of 21°E, whereas east of 21°E the Ceres and the assumed equivalent of the Bidouw Subgroup, the Traka Subgroups occur (Table 1.1) (Theron, 1972, Theron and Johnson, 1991). The entire succession thickens south and eastward (Theron, 1972, Johnson, 1976). In the Western Cape Province, the succession thickens from approximately 700 m in the Clanwilliam district to 2185 m near Villiersdorp (Theron and Johnson, 1991). In the Eastern Cape, these sections thicken eastward attaining a maximum thickness of ~ 3,500 m northwest of Port Elizabeth (Theron and Johnson, 1991). The average reported maximum thickness is approximately 3,000 m (Johnson, 1976).

The Ceres Subgroup contains 3 upward coarsening megacycles each consisting of basal mudrock to siltstone dominant lithologies that grade into increasingly sandstone rich lithologies (Table 1) (Theron, 1972, Theron and Johnson, 1991). These alternating mudrock/siltstone and sandstone lithologies form the basis for formation subdivision of the Ceres Subgroup and are summarised in Table 1.1.

Only the Bidouw Subgroup will be discussed since it crops out in the chosen study area. For comment on the Traka Subgroup in the Eastern Cape, refer to Theron (1972) and Johnson (1976). Contact with the underlying Ceres Subgroup is conformable and gradational over a few metres (Theron, 1972, Theron and Johnson,

1991). The base of the Bidouw Subgroup is formally accepted to be at the first laterally persistent mudstone, marking the start of a dominantly arenaceous unit (Theron, 1972, Theron and Johnson, 1991). In the Western Cape Province, the Bidouw Subgroup comprises 3 upward coarsening megacycles each consisting of basal mudrock beds which coarsen up into siltstone and sandstone beds (Theron, 1972, Theron and Johnson, 1991). The last upward coarsening megacycle of the Bokkeveld Group continues into the Weltevrede Subgroup (Wagen Drift/ Weltevrede Formations) of the Witteberg Group is conformable and gradational over a few metres (Theron, 1972, Johnson, 1976, Johnson, et al., 2006). As is the case in the Ceres Subgroup, the alternating mudrock/siltstone and sandstone lithologies form the basis for formation subdivision for the Bidouw Subgroup and are summarised in Table 1.

Formations of both the Ceres and Bidouw Subgroups grade into a progressively more mudrock rich facies southwards, where these units are collectively referred to as the Bokkeveld Formation beyond 34°S (Theron, 1972, Cooper, 1986).

The Gydo, Gamka, Voorstehoek, Hex River, Tra-Tra and Boplaas Formations have been formally described and accepted by the South African Committee for Stratigraphy as lithostratigraphic entities. Details pertaining to their lithological character have been presented in Theron (1999), Theron et al. (1995a), Theron (2003), Theron et al. (1995b), Basson et al. (1995) for the Gydo, Gamka, Voorstehoek, Hex River and Boplaas Formations respectively. All formations are conformable and gradational with each other Theron (1972).

1.5 Palaeontology of the Bokkeveld Group

The rocks of the Bokkeveld Group contain a rich fossil assemblage of some of the earliest embryophytes known from the Phanerozoic of Africa, these are found often in association with marine metazoans of which include a plethora of invertebrates, in addition to acanthodian, chondrichthyian, placoderm and osteichthyian gnathostomes (Seward, 1932, Plumstead, 1969, Cooper, 1982, Boucot, et al., 1983, Oosthuizen, 1984, Boucot, et al., 1986, Anderson, et al., 1999a,b, Boucot, 1999, Jell and Theron, 1999, Maisey and Anderson, 2001).

Embryophyte fossils are allochthonous to the Cape Basin and were probably brought into the Bokkeveld basin via storm events from a terrestrial source (Plumstead, 1969). These fossils are rare in the Ceres Subgroup, becoming more ubiquitous up sequence into the Bidouw Subgroup (Plumstead, 1969, Anderson, et al., 1999a). Recovered embryophytes from the Bokkeveld Group include rhyniophytes, lycopods, or proto- lycopods in addition to debated equisetophytes (Plumstead, 1969).

Reported invertebrate fossils from the Bokkeveld Group are of typical Malvinokaffric Realm affinity and similar to those reported from Devonian rocks in Antarctica, South America and the Falkland Islands (Goncalves de Melo and Boucot, 1990, Becker, et al., 1994, Boucot, 1999, Isaacson, 2007).

Gnathostome fossils have been reported from the Gydo, Tra-Tra, Klipbokkop and Adolphspoort Formations and show a high degree of endemism (Anderson, et al., 1999a, Anderson, et al., 1999b, Maisey and Anderson, 2001, Elliot, et al., 2010).

Oosthuizen (1984) recognised a distinct difference in size and faunal composition of invertebrate fossils from the Bokkeveld Group. In the western Bokkeveld Group, brachiopods and echinoderms are more common than they are in the east. In the eastern Bokkeveld Group, conularians, anthozoans and hyolithids are more common than they are in the west. Trilobites and bivalves from the eastern side of the basin tend to be larger than those in the west.

1.5.1 The Malvinokaffric Realm

The Malvinokaffric Realm, (“Austral Province”) was a biogeographic realm that persisted from the Silurian Period into the Middle Devonian of southwestern Gondwana where it is thought to have terminated at the end of the Eifelian Age (House, 1971, Boucot, 1988, Boucot, 1990, Goncalves de Melo and Boucot, 1990, Isaacson, 2007). It is possible that the Malvinokaffric Realm may have had a longer temporal range into the Ordovician Period (Boucot, 1999).

The Malvinokaffric Realm was one of three large biogeographic entities that persisted during the Early to Middle Devonian, the other two of which are the Old World Realm and the Eastern Americas Realm (“Appalachian Province”) (House, 1971, Boucot, 1988, Blodgett, et al., 1990, Boucot, 1999) (Fig. 1.16) The delineations of these palaeo - biogeographic realms are based on similarities in terms of endemic fossil fauna, in particular, those of brachiopods (House, 1969. Boucot, 1988). In addition, trilobite, rugose coral and gastropod biogeographic realms further, correlate with these brachiopod biogeographic realms, leading to the suggestion of global endemism of marine fauna during the Palaeozoic (Boucot, 1988, Eldredge and Ormiston, 1979, Pedder and Oliver, 1990, Blodgett, et al., 1990, Boucot, pers comm.). Only the Malvinokaffric Realm will be discussed since it is within the scope of this project. For details pertaining to the Old World and Eastern Americas Realm, refer to House (1971) and Boucot (1988). The fauna of the Malvinokaffric Realm are a distinct and highly endemic cold water invertebrate assemblage (Boucot, 1988, Adrain and Edgecombe, 1996, Boucot, pers. comm., Isaacson, 2007) with some authors suggesting the inclusion of highly endemic gnathostomes (Elliot, et al, 2010, Boucot, pers.comm.). The Malvinokaffric Realm of the Early to Middle Devonian extended from southern Africa (modern day South Africa), up north to Ghana, Guinea and south eastern Senegal (Early Devonian Epoch only), west into South America (modern day Bolivia, Chile, Brazil, Peru, Uruguay and Argentina) and southwards to the Falkland Islands and Antarctica (Fig. 1.16) (Boucot, 1988, Boucot, et al., 2001, Isaacson, 2007). The boundaries of the Malvinokaffric

Realm appears to have migrated southwards throughout its existence in the Palaeozoic (Boucot, 1988, Boucot, 1999). From worldwide palaeogeographic data and lithological indicators for palaeoclimatological conditions for the Devonian Period (i.e. “warm” region indicators: marine evaporites and non- marine coal, bauxite, kaolinite, calcrete, glendonite, and manganese, nickel laterite beds and “cool” region indicators: tillites and dropstones), the Malvinokaffric Realm has been envisioned to have occupied a high latitude region with cool water conditions near the palaeo-South Pole (60 to 90º S) (Boucot, 1988, Scotese and Barrett, 1990) (Fig. 1.16). Scotese and Barrett (1990) estimate that the palaeo-South Pole was situated variably in south- central Argentina (present day ~ 30ºS; 0º E to 35º S; 5º E) between Late Silurian to Late Devonian times.

Cooper (1982) questioned the inclusion of Antarctica within the Malvinokaffric Realm. He cites the presence of notable Malvinokaffric brachiopod, trilobite, bivalve and gastropod taxa in addition to Tentaculites (von Schlotheim), within the Horlick Formation of Antarctica that are the same as those found in the Bokkeveld Group and South America. However, Cooper (1982) cites the absence of several notable Malvinokaffric Realm brachiopod species, namely Australocoelia, Scaphiocoelia and Notiochonetes as well as the entire absence of calmoniid trilobites within the Horlick Formation. Based on close similarities between the trilobites Burmeisteria antarctica (Saul) from the Horlick Formation and Burmeisteria expansa (Hector) from New Zealand, which he cites as being distinctly different to that of Burmeisteria herscheli (Murchison) of the Bokkeveld Group, a closer link is suggested between Antarctica with Tasmania and New Zealand, and a less direct marine link with South Africa is more likely. The Devonian of New Zealand is within its own biogeographic realm, the New Zealand Region (Old World Realm) (Boucot, 1988).

Typical Malvinokaffric Realm faunal elements have also been found outside their palaeogeographic extent. Boucot (1988) mentions the presence of Malvinokaffric brachiopods; Austalocoelia in Tasmania and Victoria, Australia and Tanerhynchia and Pleurothyrella in New Zealand. Boucot, pers comm. He

28 suspects that the concomitant association of Malvinokaffric and non- Malvinokaffric fauna in New Zealand and Australia and possibly parts of Antarctica may possibly be located in what he has called a “boundary region mixing” where mixing of the Malvinokaffric Realm, Tasman and New Zealand Region (Old World Realm) faunules may have been possible.

1.5.2 The Faunal Assemblage and Attributes of the Malvinokaffric Realm

As mentioned earlier, the Malvinokaffric Realm occupied a high latitude position with cool water conditions near the palaeo-South Pole (60 to 90º S) during Silurian to early Late Devonian times (Boucot, 1988, Scotese and Barrett, 1990). The faunal assemblage of the Malvinokaffric Realm is unique bearing few similarities to the faunal assemblages of the contemporaneous East Americas and Old World Realms (Eldredge and Ormiston, 1979, Cooper, 1982, Boucot, 1988, Pedder and Oliver, 1990, Blodgett, et al., 1990, Boucot, 1999, Isaacson, 2007, Boucot, pers comm.). The main faunal assemblage and attributes of the Malvinokaffric Realm with respect to the East Americas and Old World Realms have been summarised by Boucot (1988) as follows: 1) Low number of taxa at superfamily, genus and species level. 2) Fewer species present in lower trophic (tier) communities than are characteristic in carbonate rich regions of the East Americas and Old World Realms. 3) An abundance of hyolithiids and conulariids. 4) A near total absence of rugose corals and bryozoans (limited to a single species each) and is summarised in Table 1.2.

1.5.3 Previous Attempts at Establishing a Biostratigraphy for the Bokkeveld Group

Biostratigraphically useful taxa for the Devonian Period worldwide are missing from Malvinokaffric Realm dominions. Most notable are ammonoids, graptolites, rugose corals and conodonts Boucot (1988). In light of this, attempts have been made to establish a biostratigraphic scheme for the Bokkeveld Group. Three informal biostratigrpahic schemes for the Bokkeveld Group were presented by Theron (1972), Cooper (1982) and Hiller (1995)

29 based on the biostratigraphic ranges of various flora and fauna. Theron (1972) presented a scheme based on the ranges of specific fossil taxa throughout the Bokkeveld Group (Fig. 1.17). Cooper (1982) has presented a biostratigraphic scheme based soley on the ranges of specific trilobite (Fig. 1.18). Hiller (1995) has presented the ranges of chonotacean brachiopods from the Bokkeveld Group and has mentioned their biostratigraphic use (Fig 1.19).

Chapter 2: Aims and Methods

2.1 Aims

1. Reassess the lithostratigraphy of the Ceres Subgroup in the Clanwilliam Sub – basin in the Western Cape Province. 2. Create a lithofacies model for the Ceres Subgroup in the Western Cape Province 3. Discuss the lithological relationships among sequences and parasequences in the Ceres Subgroup in an attempt to understand dynamics of the Cape Basin during the time of Emsian to Eifelian Age Bokkeveld Group sediment deposition via sequence stratigraphy. 4. Locate possible controls on fossil assemblages, temporal vs. spatial controls. 5. Attempt to create a biostratigraphic and ecostratigraphic framework for the Ceres Subgroup based on recovered fossil taxa. 6. Describe any new fossil taxa that may be discovered. 7. Discuss the palaeo - environments and palaeo - ecology of Ceres Subgroup taxa.

2.2 Materials and Methods

This study is restricted to the Ceres Subgroup in the Cederberg Range, Western Cape, Province South Africa. The Cederberg Range was selected as a study area due to the low degree of tectonic deformation and metamorphism present in this part of the CFB. Three study sections were identified along the Cederberg Range after a short fieldtrip was undertaken in the area. Study sections are present along a 42 Km north-south transect of the Cederberg Range and are located within the Cederberg Wilderness Area at the farms Grootrivierhoogte (GPS: 32°38'55"S; 19°24'43"E), Keurbosfontein (GPS: 32°27'32"S; 19°18'39"E) and the town of Wupperthal (GPS: 32°18'16"S; 19°14'24.88"E) (Fig 2.1). Study sections were selected due to their low tectonic deformation, accessibility and relatively continuous vertical exposure of the Ceres Subgroup.

Fieldwork was undertaken over three months in total in May – June and October - November 2012. Stratigraphic sections were measured at each locality using a Jacob’s staff and Abney level. The sedimentology of each section was described in terms of their sedimentary textures, architectural features, sedimentary structures, colour (according to the Munsell colour system), ichnofossils and flow current indicators were taken from various sedimentary structures. Photographs and photo mosaics of select attributes in outcrop were also taken. The stratigraphic height at which fossils were found was also recorded. Fossils found at the study sections were identified and described in situ, those in danger of being destroyed, or of descriptive value were removed, catalogued with field catalogue numbers and their Global Positioning System (GPS) co-ordinates recorded. Removed fossil specimens were compared and identified with the aid of Bokkeveld Group collections housed at the Council for Geoscience, Bellville, Iziko Museum, Cape Town, as well as from published literature on fossils from the Bokkeveld Group and international Malvinokaffric Realm correlates. Fossils in need of preparation to reveal descriptive features were done so by preparation staff at the Evolutionary Sciences Institute before being photographed. Removed fossils are to be housed in the collections of the Iziko Museum, Cape Town and are awaiting accession numbers from the museum.

Identification of formational boundaries of the Gydo, Gamka, Voorstehoek, Hex River, Tra-Tra and Boplaas Formations were made based on lithological lower boundaries presented by Theron (1999), Theron et al. (1995a), Theron (2003), Theron et al. (1995b), Basson et al. (1995) for the Gydo, Gamka, Voorstehoek, Hex River and Boplaas Formations respectively. The lower boundaries of the Gydo and Voorstehoek Formations are taken as the first laterally continuous horizon above which mudstone is more abundant than sandstone. The lower boundaries of the Gamka, Hex River and Boplaas Formations are taken as the first laterally continuous horizon above which sandstone is more abundant than mudstone. Since no formal lithostratigraphic description of the Tra-Tra exists, its lower boundary has been taken as the first laterally continuous horizon above which mudstone is more abundant than sandstone.

Chapter 3: Geology of Study Area

3.1 General Geology and Geography of Study Area

The study area is located in the Koue Bokkeveld region of the Cederberg and follows the boundary of the west lying Fynbos and east lying Succulent Karoo biomes. The division between biomes is noticeably drawn at the contact between the arenaceous Nardouw Subgroup (Table Mountain Group) and the argillaceous Bokkeveld Group.

Strata of the Bokkeveld Group within the study area are exposed within the eastern limb of an extensive open antiformal monocline with an apparent axial trace trending north-south along the entire study area (Fig. 3.1). Strata of the Bokkeveld dip eastwards and have an average orientation of 165°/ 28°/E (given as azimuth/dip angle/dip direction). The differences in weathering profiles between argillaceous and arenaceous rich formations of the Bokkeveld Group give it a distinctive hogsback topography. Shearing, as indicated by the development of slickenslide lineations (slickenlines) and associated quartz slickenfibre mineral growths is present in the bedding planes between individual beds (Fig. 3.2). This, however, is not ubiquitous. Slickenfibre growth directions indicate a right lateral (eastward) shear sense and displacement along beds in the dip direction of the fold limb. Shearing is most evident at the interface between shale and sandstone rich beds and appears to be a reflection in competency contrast between the lithologies. Kaolinitisation and sericitisation along shear planes may be locally present at outcrop, but is not ubiquitous.

The orientation of strata and accompanying deformation varies slightly throughout the study area. At Grootrivierhoogte strata have an average orientation of 186°/16°/E with little to no deviation in terms of dip and strike in the section. In the Hex River and Boplaas Formations at Grootrivierhoogte, folding may severely distort bedding such that internal features are unrecognisable, or completely absent (Fig. 3.3 A). This is accompanied by jointing and quartz “crack seals” within joints (Fig.3.3

B and C). Crack seals are oriented east-west in the inferred major compression (δ1) direction.

At Keurbosfontein the attitude of strata is highly variable. The base of the Keurbosfontein section follows a large negative weathering physiogeographic feature informally known as Moordenaarsgat (Fig. 3.4). Moordenaarsgat is a gorge which trends parallel to the Nardouw Subgroup – Bokkeveld Group contact and may be traced out to Wupperthal. Strata in close proximity to Moordenaarsgat, namely the Gydo, Gamka, Voorstehoek and Hex River Formations have an average orientation of 166°/40°/E. In the Gamka Formation at Keurbosfontein, strata may be folded and steepen up to a dip 60° (Fig. 3.5 A). This coincides with right lateral strike slip faulting and jointing (Fig. 3.5 B and C). A displacement of 3.0 m has been recorded along this strike slip fault. Strike slip faults are oriented roughly east-west and may be relays or jogs between folds. In the area around Moordenaarsgat, the attitude of the strata gradually changes to 168°/20°/E in the Hex River Formation and stays constant in the Tra-Tra and Boplaas Formations with no deformation.

At Wupperthal the attitude of strata varies from 162°/20°/E in the Gydo Formation to146°/30°/E in the Gamka Formation, 156°/ 29°/E in the Voorstehoek and Hex River Formations and 162°/20°/E in the Tra-Tra and Boplaas Formations. Accompanying deformation features are minor and restricted to slickenslides and slickenfibre mineral growths between bedding planes.

3.2 Introduction to Sedimentary Facies Analysis

Sedimentary facies refer to aerially restricted sedimentary bodies which are united by internal characteristics that separate them from other sedimentary bodies in stratigraphy (Boggs, 2006, Nichols, 2009). Internal characteristics which may be used to discriminate between sedimentary rocks in field observations include: mineralogy, texture, sedimentary structures, fossil content, and/or other organic contents (Boggs, 2006, Nichols, 2009). Facies may be discriminated by using particular criteria. Thus, a sedimentary facies identified upon the basis of key lithological characteristics would be a lithofacies, upon fossil content, a biofacies and so on (Boggs, 2006, Nichols, 2009). Facies may in turn be grouped into broader

associations of facies which are unified by similar characteristics linked by similar processes which aided their formation (Nichols, 2009).

Facies associations ideally should be conformable and obey Walther’s Law and thus be vertically distributed in stratigraphy as they are laterally related in space (Boggs, 2006, Nichols, 2009). Recognising how facies associations succeed one another in stratigraphy may be used as a first proxy for palaeoenvironmental reconstructions at time of deposition (Boggs, 2006, Nichols, 2009).

3.3 Current lithofacies model for the Bokkeveld Group (Theron, 1972, Tankard and Barwis, 1982, Theron and Loock, 1988)

Two lithofacies models for the Bokkeveld Group have been presented by Theron (1972), Tankard and Barwis (1982) and Theron and Loock (1988). Both models are strongly hinged upon an essentially deltaic palaeo-environment of deposition. These are a storm/wave dominated delta model and a mixed wave/tidal dominated delta model. Both models strongly suggest a high degree of uniformity in each of the 5 to 6 upward coarsening megacycles of the Bokkeveld Group depending, on the position in the Basin.

3.3.1 Storm/wave dominated Delta Model (Theron, 1972, Theron and Loock, 1988)

A storm/wave deltaic model for the Bokkeveld Group was first proposed by Theron (1972). This model was proposed on the recognition of the five to six upward coarsening megacycles which typify the Group. Each upward coarsening megacycle was reasoned by Theron (1972) to represent an event of progradation of a series of deltas into the Cape Basin after a period of high eustatic sea level. Theron (1972) based the storm/wave deltaic model on the recognition of barrier island type deposits.

Although no formal lithofacies model was proposed by Theron (1972) for the entire Bokkeveld Group, a broad palaeo-environmental model was presented

35 based on the different lithologies present in the Group. The lithologies and their corresponding palaeo-environments were later formalised by Theron and Loock (1988) into a “southern distal”, “medial” and “northern proximal” zone facies. The distal zone facies, dominated by mudstone lithologies is thought to represent prodelta and offshore shelf environments, these grade northwards into increasingly more silt and greywacke rich medial facies representative of delta front and distributary bay environments. The proximal zone facies gradationally overlies the medial zone facies and is representative of more mature lithic and quartz arenites with mudstone intraclasts and thin gritstone and conglomerate stringers. Minor intercalating siltstone and sandy shales may occur within the proximal facies. The proximal zone facies is representative of mixed deltaic and shallow marine environments encompassing beaches, distributary mouth bars, point bars, barrier lagoons and delta channels. Expected backshore and subaerial delta plain environments (encompassing backshore dunes, back swamps etc.) are thought to have been removed by transgressive erosion with initiation of overlying upward coarsening megacycles as well as shoreline reworking and lateral amalgamation by strong storm/wave activity (Theron, 1972 and Theron and Loock, 1988). Theron (1972) and Theron and Loock (1982) favour and illustrate sediment sourcing from the north, west and east for the Bokkeveld Group in the Clanwilliam Sub-basin at various times with deltas prograding into the Clanwilliam Sub-basin south, east and west respectively.

3.3.2 Mixed wave/tidal dominated Delta Model (Tankard and Barwis, 1982)

Tankard and Barwis (1982) presented a more formal lithofacies scheme for the Bokkeveld Group, albeit restricted to outcrop in the Cederberg area within the Clanwilliam Sub-basin. This model is strongly hinged upon Theron’s (1972) deltaic palaeo-environmental scheme for the Bokkeveld Group and assumes Theron’s (1972) interpretation of sandstone units to be laterally amalgamated by wave activity. Tankard and Barwis (1982) proposed a mixed wave and tidal dominated deltaic facies model on the basis of several tidal indictors mentioned in the proceeding text. They favour sediment sourcing

from the north and propose that delta’s of the Bokkeveld Group prograded southwards. This facies model has been summarised in Figure 3.6 and will be briefly discussed. Five major facies associations have been recognised by Tankard and Barwis (1982), these are: Shelf – Prodelta, Distributary Mouth Bar, Tidal Flat, Interdistributary Bay and a Beach-Shoreface Complex.

Shelf – Prodelta Facies Association: This facies association is characterised by grey to black finely laminated, ripple cross laminated and massive shales and siltstones which may be locally calcareous in their lower parts. Centimetre scale graded beds and fine grained ripple cross laminated greywacke lenses become more common up sequence as they grade into the overlying Distributary Mouth Bar Association. Soft sediment mass movement features, including flow rolls, slumps and disturbed bedding, may be present within this Association.

Casts and molds of typical Malvinokaffric Realm invertebrates are common and include trilobites, brachiopods, bivalves and crinoids. Coquinites occur within the Association as 10 to 15 cm thick lenses with randomly oriented shell debris. Large (up to 2.0 m) siderite nodules are interspersed within the Facies Association.

Tankard and Barwis (1982) considered thin finely laminated shales and siltstones to have deposited via slow suspension settling of silts and clays in a low energy, offshore environment. They proposed that massive mudstones represent periods of pervasive bioturbation. Greywacke lenses and coquinites represent periods of occasional storm activity. Siderite nodules are explained by Tankard and Barwis to be early diagenetic features. This Facies Association may reach up to 100 m in thickness.

Distributary Mouth Bar Facies Association: This Facies Association is characterised as being a 100 m thick thinly bedded heterolithic sandstone and mudstone unit which coarsens upward into very fine grained greywackes and eventually to fine to medium grained lithic arenites. The lower 20 to 40 m of the Association is dominated by climbing ripple cross laminated to lenticular

and wavy bedded and low angle cross bedded siltstone. Convolute, ball and pillow and flow roll structures may be present within the Facies Association. Numerous 30 to 100 cm thick sandstone lensoids are disconformably interspersed within the Facies Association. Towards the top of the Facies Association, greywackes coarsen upward to lithic arenites. This is accompanied by an increase of average bed thickness (from 10 to 50 cm) as well as an increase in average cross bed size (from 10 to 30 cm). Mudstone intraclasts are present on top of bedding planes, or interspersed among individual foresets of cross beds. A change from linear crested oscillation ripples to current ripples concomitantly occurs over this interval.

Brachiopod and mollusc casts and molds are present in the Facies Association, though being not as numerous as in the Shelf – Prodelta Facies Association. Coalified plant debris is also present on top of bedding planes. Bioturbation increases upward in the Facies Association and is common in sandstone beds. Representative ichnofossils are straight, branching and U- shaped traces which resemble Arenicolites. Tankard and Barwis (1982) reason that the ichnofossil assemblage resembles that of the Skolithos and Glossifungites ichnofacies.

Tankard and Barwis (1982) considered these features to be representative of distributary mouth bar environments. Heterolithic mudstone and sandstone beds have been likened to alternating suspension and traction transport of sediment by distributaries with fluctuating discharge under possible tidal activity. Cross beds are thought to represent distributary mouth bars which themselves may be cut by 3.0 to 5.0 m deep side filled distributary channels.

Tidal Flat Facies Association: Individual 2.0 to 10.0 m thick upward fining beds consisting of basal quartz arenites, overlain by mudstones and shales characterise this facies association. Heterolithic flaser and lenticular bedding and bimodal and bipolar mud draped trough and planar cross bedding is common within the Facies Association. Ladderback, asymmetric and symmetric ripple cross laminae are common at the tops of individual beds. Lenticular beds of1.4 to 2.0 m thick side filled quartz arenite and sandy

38 mudstone disconformably cut into underlying strata. Runzel and rill marks as well as sand infilled desiccation cracks and mud cracks are common features on tops of bedding planes. In phase climbing ripple cross lamination and soft sediment deformation are rare.

Body fossils are rare and largely restricted to linguloid brachiopods. Skolithos and Zoophycus ichnofossils indicative of bioturbation are common and present through the unit as a pervasive feature.

Tankard and Barwis (1982) favour a tidal environment of deposition for this Facies Association based on the assumption that the bimodal and bipolar foreset orientation of cross beds and flaser bedding represent alternating ebb and flow conditions. Runzel and rill marks, as well as mud and desiccation cracks are favoured to have been formed by emergence of tidal flats during ebb tide.

Interdistributary Bay Facies Association: The Interdistributary Bay Lithofacies Association is characterised as being a pervasively bioturbated lower grey-brown mudstone and an upper sandstone where it coarsens upward into the overlying Beach-Shoreface Complex Lithofacies Association. The lower parts of the Facies Association are disconfomably intercalated by erosionally based lensoidal lithic greywackes and arenites. Soft sedimentary deformation features such as load casts and slump structures are common throughout the Facies Association. Uppermost sandstones of the Facies Association are characteristically ripple cross laminated to ladderback rippled.

Zoophycus and other vertical burrows are typical of the ichnofauna present in the Facies Association. Articulated plant fossils are common on bedding planes and may be found with rare rhizolith horizons and carbonaceous mudstones. Siderite nodules may be present.

Tankard and Barwis (1982) reason that the lithofacies are indicative of fine grained overbank material that would have been deposited lateral to distributaries and distributary mouth bars in relatively calm and deeper water

settings. These deposits coarsen upward moving into distributary mouth bar and channel environments. Intercalating sharp based sandstones are thought to represent crevasse splay deposits with subaqueous delta construction. Rhizoliths are consideredto be those of semi-aquatic plants, found in more proximal interdistributary bay regions. The Facies Association may be up to 17.0 m thick in outcrop.

Beach-Shoreface Complex Facies Association: This Facies Association is representative of several palaeo-environments which formed in response to sediment deposited in the delta front being reworked by storm/wave activity. Each palaeo-environment shall be discussed separately.

Washover: Tankard and Barwis (1982) state that washover palaeoenvironments are rare in the Bokkeveld Group and are characterised as being planar parallel laminated, fine grained sandstones which may have shallow relief scours and washed out ripple laminae. Individual laminae vary from 2 to 5 mm thick. Washover successions tend to be thinner and muddier in their proximal expression and thicker and well sorted distally.

Tankard and Barwis (1982) typify these successions as being formed by storm washover processes in microtidal coasts derived from foreshore and foredune erosion which interfinger distally with lagoonal mud’s. These sequences may reach up to 18.0 m thick.

Tidal Inlets and Tidal Channels: These successions are the most common sequences in the Bokkeveld Group according to Tankard and Barwis (1982). Tidal inlet and channel successions disconformably overly both washover and distributary mouth bar successions and are in turn overlian by the tidal flat facies association. Sediment deposited withintidal inlets and channels are manifested as well sorted fine to medium grain quartz arenites with lenticular geometry with large lateral accretion sets. Individual lenticles range from 0.30 to 8.0 m thick (average of 1.2 m thick) and an average width to depth ratio of 15:1.

Individual lenticles are erosively based and truncate each other and stack on top of one another throughout the succession. Lenticles have quartz pebble and mudstone intraclast channel lags at their bases and fine upwards. Internal sedimentary structures within lenticles include small scale trough and planar cross bedding, intercalating plane beds and ripple cross lamination. Foresets of individual cross beds are clay draped, often with mudstone intraclasts. Toesets of individual cross beds may be herringbone and backflow ripple cross laminated and/or flaser bedded. Asymmetric current ripples and oscillation ripples are present on the topmost bedding planes of individual lenticular beds. Moving up individual tidal inlet and channel successions, Tankard and Barwis (1982) noted an overall increase in bed size and bed thickness.

Associated Skolithos – Monocraterion ichnofossils are common. Here individual Skolithos traces are topped by Monocraterion. These commonly only occur where cross beds are thinnest (~ 10.0 cm).

With comparison to cited modern examples as well as on the basis of sandstone maturity, bimodal and clay draped cross beds and the abundance of asscoiated Skolithos – Monocraterion ichnofossils, Tankard and Barwis (1982) suggested a tidal inlet for thicker parts of the succession which were fed by ebb dominant tidal channels.

3.4 Lithofacies Descriptions

As a result of fieldwork for this project, a total of 12 lithofacies have been recognised within the Bokkeveld Group in the study area and are described in terms of their grain-size, sorting, geometry, bounding surfaces, colour and bioturbation. These are summarised in Table 3.1. The vertical distributions of the described lithofacies at all three measured study sections have been illustrated in Appendix 1 - 4. Reported colour codes are referenced from the Munsell colour system. Quoted thicknesses of lithofacies are given as average bed thicknesses and entire vertical thickness in outcrop.

3.4.1 Lithofacies 1 (Clayshale Lithofacies)

Description: Lithofacies 1 is characterised as being a lower flow regime plane bed, planar horizontal, laminated bedded clayshale (sensu Potter et. al., 1980). Clayshales, as defined by Potter et al. (1980) are mudstones with 66- 100% clay content and are finely laminated or massive in places.

Fresh cuttings of Lithofacies 1 are typically Dark Grey to Dusky Blue Green (10 G 3/1), or Dark Blueish Grey (10 B 2.5/1) in colour but may vary to Olive- Grey (7.5 GY 4/2), Olive (5Y 5/3), Pale Greenish Grey (10 GY 4/2) in weathered exposures. The lithofacies is fissile and crumbles easily.

Lithofacies 1 is the most fossil rich lithofacies within the study area and fossils are commonly preserved as casts or moulds (Fig. 3.7 A, B). Nodular preservation of fossils is also present (Fig. 3.7 C). Fossiliferous nodules may singly be interspersed throughout the lithofacies, or occur within discrete nodule beds (Fig. 3.7 D). All three preservation styles may co-occur with one another within the lithofacies (Fig. 3.8).

Monospecific coquinites may also be present in Lithofacies 1. Coquinites are bioclast supported within a clay rich matrix (Fig. 3.9). Individual bioclasts show very little, or, no signs of damage. Monospecific coquinites of Lithofacies 1 tend to contain bioclasts of the same size, possibly indicating a degree of sorting, possibly as a result of their hydrodynamic properties.

Large carbonate nodules are interspersed within Lithofacies 1 (Fig. 3.10). These nodules (measured along their long axes) range in size from 0.3 m up to 2.0 m but commonly are between 0.6 to 0.9 m. Internally, nodules are massive coarse grained and may have calcite precipitated septarian cracks which permeate throughout the nodules (Fig. 3.10.B). This is the usual association. However, some carbonate nodules show evidence for being internally laminated (Fig. 3.10C). Carbonate nodules tend to be stratabound with numerous nodules occurring at the same stratigraphic height (Fig. 3.10 D).

Within the study area, Lithofacies 1 is present only within the Gydo Formation at all three measured sections (Appendix 1 - 4). Within the Gydo Formation at Grootrivierhoogte, Lithofacies 1 attains a maximum thickness in outcrop of 49.5 m and a minimum of 30.0 m. At Keurbosfontein, Lithofacies 1 attains a maximum thickness in outcrop of 84.0 m. At Wupperthal, Lithofacies 1 attains a maximum thickness in outcrop of 48.0 m.

Process Interpretation: Clayshales are interpreted to form in low energy water environments via slow suspension settling of clay minerals as defined laminae (Potter, et al., 1980, Collinson and Thompson, 1989a, Blatt, 2006). The localised massive nature of Lithofacies 1 may be explained by either bioturbation, or continuous rapid suspension sedimentation (Potter, et al., 1980)

3.4.2 Lithofacies 2 (Siltstone Lithofacies)

Description: Potter et al. (1980) characterises indurated shales with 0-32% clay mineral content as siltstones. Lithofacies 2 consists of two associated sub-lithofacies that may easily be distinguished in the field. These are namely Sub-lithofacies 2.1: Plane -laminated siltstone and Sub-lithofacies 2.2: Ripple bedded siltstone. Both subfacies are unified by possessing a mineralogy of <33 % clay sized grains and as such are dominated by silt sized grains with minor quantities of very fine to fine grained sand.

Lithofacies 2 tends to coarsen upward from Lithofacies 1 with a decrease in clay content. Lithofacies 2 may also grade into Lithofacies 3 with an increase in very fine, to fine gained quartz sand content.

As with Lithofacies 1, nodular, cast and mould style preservation of fossils is present within Lithofacies 2.

3.4.2.1 Sub-lithofacies 2.1 (Plane -laminated Siltstone Sub- lithofacies)

Description: Sub-lithofacies 2.1 is characterised as being a lower flow regime, plane-laminated bedded siltstone (sensu Potter et al., 1980). Individual lamina are fine and millimetre scale (approximately 1-3 mm thick per lamina). White mica may be present as an accessory mineral in Sub-Lithofacies 2.1. The sub-lithofacies is fissile and crumbles easily along lamina. Occasional slumps and pillow structures and accompanying soft sediment deformation may be present (Fig. 3.11).

Colour varies from Dark Yellowish Orange (10 YR 6/6), Strong Brown (5 YR 4/6), Medium Light Grey (10 Y 6/2), Dark Blue Green (2.5 Y 3/2), Pale Olive (5 Y 5/4) and Medium Blueish Grey (10 BG 5/1).

Within the study area, Sub-lithofacies 2.1 is present within the Gydo and Voorstehoek Formations at Grootrivierhoogte and in the Gydo Formation at Wupperthal. At Grootrivierhoogte, Sub-lithofacies 2.1 attains a maximum thickness of 15.0 m in the Gydo Formation and a maximum thickness of 10.5 m and a minimum thickness of 6.0 m within the Voorstehoek Formation. Sub-lithofacies 2.1 attains a maximum thickness in outcrop of 6.0 m within the Gydo Formation at Wupperthal.

Process Interpretation: Laminated siltstones are interpreted to form in low energy environments via slow suspension settling of clay to fine silt minerals (Potter, et al., 1980, Collinson and Thompson, 1989a, Blatt, 2006).

3.4.2.2 Sub-lithofacies 2.2 (Ripple bedded Siltstone Sub- lithofacies)

Description: Sub-lithofacies 2.2 is a ripple bedded siltstone. Individual ripple beds consist of a flat base overlain by a sinuous form crest

interpreted to be a ripple crest (Fig. 3.12 A and B). Internal ripple laminae are not visible. Individual ripple beds are discrete and are 10 to 30 mm thick with thicknesses of 10 to 20mm being most common. Sub-lithofacies 2.2 differs from Sub-lithofacies 2.1 in that it is harder, is not as fissile, and tends to break apart as irregular blocks. Red oxide nodules are almost always present (Fig. 3.12 C ) and may be fossiliferous (3.12 D).

Occasional intraclast and bioclastic conglomerate based sandstone filled scours and troughs are present within Sub-lithofacies 2.2 and may represent episodic storm event beds (Fig. 3.13).

Sinuous form and un-networked surface ichnofossils are present within the Sub-lithofacies and are restricted to the tops of bedding planes (Fig. 3.14) These resemble Nereites sp., or those associated with the Nerieites ichnofacies (Miller, 2009, Seilacher, 2007). Coalified fragmentary plant fossils may also be present on bedding planes as carbonaceous films (Fig. 3.15).

Colouration of fresh outcrop of Sub-lithofacies 2.2 ranges from Dark Blueish Grey (5 PB 3/2, 10 B 2.5/1, 5 PB 2.5/1) to Dark Greenish Grey (5 GY 4/1, 2.5 Y 3/2) to Dusky Blue Green (10 G 3/1) and Olive (5 Y 5/6). Sub-lithofacies 2.2 usually retains its colouration even when weathered, but may alter to 10 YR 6/4.

Sub-Lithofacies 2.2 is present at all three measured sections in the study area and is present in the Gydo and Voorstehoek Formations (Appendix 1 – 4). At Grootrivierhoogte, Sub-lithofacies 2.2 ranges in thickness in outcrop from a maximum of 26.0 m to a minimum of 13.9 m in the Gydo Formation and a maximum of 6.0 m in the Voorstehoek Formation. At Keurbosfontein, Sub-lithofacies 2.2 reaches a maximum thickness in outcrop of 19.5 m in the Gydo Formation. In the Voorstehoek Formation at Keurbosfontein, Sub-lithofacies 2.2 ranges from a minimum thickness of 6.0 to a maximum of 37.5 m in outcrop.

At Wupperthal, Sub-lithofacies 2.2 ranges in thickness in outcrop from a minimum of 1.1 m to a maximum of 2.3 m in the Gydo Formation. In the Voorstehoek Formation, Sub-lithofacies 2.2 ranges in thickness in outcrop from a minimum of 4.5 m to a maximum thickness of 24.0 m.

Process Interpretation: Ripples are bedforms which form via traction transport, deposition and migration of silt to coarse sand sized grains (ɸ 4 – 0.5) under lower flow regime conditions (Fr<1) with minor water surface disturbance or when water waves are out of phase with the deposited bedforms (Boggs, 2006).

3.4.3 Lithofacies 3 (Heterolithic Silty Sandstone and Mudstone Lithofacies)

Description: Lithofacies 3 is a heterolithic lithofacies consisting of alternate millimetre scale silty sandstone (silt to fine grained sand) and mudstone (silt and clay) laminae. Silty sandstone laminae variably may be continuous to discontinuous ripple laminated or planar laminated, or be present as discontinuous “starved ripples”. Mudstone laminae tend to occur as gradational drapes over silty sandstone laminae. Silty sandstone laminae tend to have sharp bases when overlying mudstone lamina. Internal geometry may variably be wavy, parallel to discontinuous, wavy or parallel.

Three variants of Lithofacies 3 are evident in the study area based on the relative proportions of very fine grained quartz sand and mud content as well as their internal sedimentary structure. As such, Lithofacies 3 may be representative of a continuum of lenticular and flaser laminated end members with wavy lamination being intermediate between the two. The characteristics of lenticular, wavy and flaser lamination have been summarised by Reineck and Singh (1975). Lenticular heterolithic bedding is characterised by incomplete, horizontally and vertically isolated sand lenses (or lenticular bodies) within an overall muddy body. Individual sand lenticles form on a muddy substratum and are completely overlain by mud. Flaser heterolithic bedding is characterised by horizontally and vertically continuous sand bodies

with thin mud laminae. Here, mud laminae are restricted to complete ripple trough fills and thin ripple crest coverings. Wavy heterolithic bedding is intermediate to flaser and lenticular bedding and is characterised by horizontally continuous sand and mud laminae. Sand laminae are not vertically continuous nor are they connected.

Flaser laminated heterolithic beds tend to be predominantly continuous ripple cross laminated. Ripple cross laminae silty sandstone flasers may either be symmetric wave ripples (Fig. 3.16 A), or asymmetric current ripples (Fig. 3.16 B). Wavy laminated heterolithic beds tend to be vertically discontinuous and horizontally continuous to discontinuous planar horizontal laminated and ripple laminated (Fig. 3.16 C). Sediment starved ripples are restricted to lenticular heterolithic beds and may be present in wavy laminated heterolithic beds (Fig. 3.16 C and 3.17). Sediment starved ripples are disconnected ripples which are deposited due to a dearth in sediment supply from flow (Reineck and Singh, 1975, Potter, et al., 1980).

Occasional hummocky and swaley cross stratified beds may occur within Lithofacies 3. Where hummocky and swaley cross stratification is present, Lithofacies 3 is sufficiently reworked such that that primary heterolithic bedding and associated sedimentary structures are destroyed, and appears as a sandy siltstone (Fig. 3.18). This reworking is most prevalent in the Gydo, Gamka, Voorstehoek and Hex River Formations at all three study sections, usually when intercalated with Lithofacies 4; see Lithofacies Association B.

When in association and lateral to lenticles of Lithofacies 6, Lithofacies 3 may be reworked as a ripple drift laminated sandy siltstone (Fig. 3.19). These climbing ripple cross lamina eventually laterally grade into rippled or planar laminated equivalents of Lithofacies 3.

Bioturbation may be evident in Lithofacies 3 and to such an extent that it may destroy the original sedimentary structure of the lithofacies. Bioturbation is greatest in Lithofacies 3 within the Tra-Tra and Boplaas Formations (Fig. 3.20

A), whilst minor in other formations at all three study sections (Fig. 3.20 B) (Appendix 1 - 4).

Coquinites within Lithofacies 3 differ from those typically found in Lithofacies 1. Bioclasts within Lithofacies 3 tend to contain a mixed fauna assemblage with a great degree of damage and no evidence of internal sorting since different size classes of bioclasts are present. Sand supported coquinites are found only in Lithofacies 3 when in Lithofacies Association B. Figure 3.21 is a typical example of such coquinas present in Lithofacies 3.

Upper and lower contacts of Lithofacies 3 are gradational with Lithofacies 1 and 2. Contacts between Lithofacies 3 and 4 are sharp and erosional. The contact between Lithofacies 3 and 11 is sharp. With respect to Lithofacies 11, this contact is marked by an erosional and tangential lower surface and a wavy upper surface (see Lithofacies Association E).

Colouration of Lithofacies 3 is varied and may be Moderate Red (5 R 5/4), Pale Yellow Brown (10 YR 6/4), Greyish Orange (10 YR 7/4), Bluish Grey (7.5 GY 5/2), Dark Greenish Grey (10 GY 4/2, 5 GY 4/1), Pale Olive (5 Y 6/4, 5 Y 8/4), Bluish White (10 Y 8/1), Olive Grey (7.5 GY 4/2), Light Bluish Grey (5 B 8/1, 10 BG 5/1, 10 B 8/1), Dark Bluish Grey (10 B 2.5/1, 5 PB 4/1), Dark Grey, Light Brownish Grey (5 YR 6/2), Pale Greenish Yellow (5 Y 8/2), Very Light Grey (5 Y 8/1), Olive Brown, Dark Greenish Grey, Greenish Grey (5 G 5/2), Brown (10 YR 4/6), Dark Yellow Orange (10 YR 6/6) or Pale Yellow (5 Y 7/4).

Lithofacies 3 is present at all 3 measured sections within the study area and is intimately associated within Lithofacies Association B, E, F and G thus making accurate measurements of Lithofacies 3 troublesome (Appendix 1 - 4). Thicknesses given here are of Lithofacies 3 as a standalone unit in outcrop. At Grootrivierhoogte, Lithofacies 3 attains a maximum thickness in outcrop of 16.5 m in the Gydo Formation, 21.0 m in the Voorstehoek Formation, 35.8 m in the Tra-Tra Formation and 3.0 m in the Boplaas Formation. A minimum thickness in outcrop of 7.5 m was recorded for Lithofacies 3 in the Voorstehoek Formation at Grootrivierhoogte. At Keurbosfontein, Lithofacies 3

attains a maximum thickness in outcrop of 10.0 m in the Voorstehoek Formation, 35.5 m in the Tra-Tra Formation and 19.5 m in the Boplaas Formation. At Wupperthal, Lithofacies 3 attains a maximum thickness in outcrop of 16.5 m in the Tra-Tra Formation and 13.5 m in the Boplaas Formation.

Process Interpretation: Pulsatory or episodic transport of sediment followed by suspension deposition with waning energy, or against a background of continuous suspension deposition in generally calm water conditions (Reineck and Singh, 1975, Potter, et al., 1980, Collinson and Thompson, 1989a). Sand and coarsest silt are deposited as ripple beds and mud settles out after by suspension deposition (Reineck and Singh, 1975, Potter, et al., 1980, Collinson and Thompson, 1989a). Episodic transport of sediment may be attributed to seasonal or climatic factors, or gradually increasing/decreasing flow episodes such as storm or flood events (Reineck and Singh, 1975, Collinson and Thompson, 1989a). Lenticular and flaser heterolithic beds form in environments which favour deposition and preservation of mud and sand respectively (Reineck and Singh, 1975).

3.4.4 Lithofacies 4 (Hummocky and Swaley Cross Stratified Sandstone Lithofacies)

Description: Lithofacies 4 is characterised as being a poorly sorted silt to fine grained hummocky cross stratified (HCS), and/or swaley cross stratified (SCS) sandstone. HCS beds are typified as possessing internal laminae which are convex up whilst SCS beds are concave up and mimic the hummock and swale topography of their underlying scoured beds (Fig. 3.22 A and B and Fig. 3.23). Internal laminae are between 1 and 5 mm and may rarely thicken up to 10 mm. Thickness of individual internal laminae may change within one bed with laminae pinching out against scoured planes. HCS and SCS beds have wavelengths (measured from hummock to hummock and swale to swale) of up to 5.0 m. Both HCS and SCS beds have internal laminae which are angled between 5° and 10° with respect to the bedding plane (Fig. 3.22 C). The dip angle of individual internal laminae

49 diminishes moving up individual beds of Lithofacies 4 until it becomes superficially planar. HCS and SCS beds always overly low angled planar to curved scour planes. In places, scour planes may contain erosional lag detritus consisting of mudstone intraclasts and shelly material (Fig. 3.24). Mudstone intraclasts may be present within the unit with variable size ranges (Fig. 3.22 D).

Variably, top bedding planes of HCS and SCS beds may possess symmetric, asymmetric and combined flow “ladderback” ripple laminae (Fig. 3.25 A - C). Combined flow “ladderback” ripples are most common. Top bedding planes of HCS and SCS beds may also posses parting lineations (Fig. 3.25 D).

Fossils within Lithofacies 4 are scarce and restricted to cast and mold type preservation with a lower species diversity compared with preceding Lithofacies (Fig. 3.26). Where present on lower bedding planes, fossils are disarticulated and may be highly damaged and contain an array of cast and mold fossils as fossil shelly middens (Fig. 3.27).

Lithofacies 4 is commonly Dark Grey but may be Greyish Blue (5 PB 5/1).

This Lithofacies is present within all formations in the study area, except for the Tra-Tra and Boplaas Formations (Appendix 1 - 4). The Lithofacies may either be part of Lithofacies Association C where it intercalates among Lithofacies 3, or part of Lithofacies Association D where it is amalgamated and may be overlain by Lithofacies 5.

Lithofacies 4 is present within Lithofacies Association B and C. In Lithofacies Association B, Lithofacies 4 is lensoidal in geometry and non-amalgamated with individual lensoids reaching 0.2 – 0.4 m thick and extending laterally up to 9.0 m in outcrop. In Lithofacies Association C, Lithofacies 4 is amalgamated making estimations of individual bed thicknesses troublesome. Thicknesses quoted here are of Lithofacies 4 as amalgamated units in outcrop in Lithofacies Association C. At Grootrivierhoogte, Lithofacies 4 attains a maximum thickness of 12.5 m and a minimum of 9.0 m in the Gamka

Formation. In the Hex River Formation, Lithofacies 4 attains a maximum thickness of 4.0 m. At Keurbosfontein, Lithofacies 4 attains a maximum thickness of 30.0 m and a minimum of 9.0 m in the Gamka Formation. In the Hex River Formation, Lithofacies 4 attains a maximum thickness of 16.2 m. At Wupperthal, Lithofacies 4 reaches a maximum thickness of 9.1 m in the Hex River Formation and 4.2 m in the Hex River Formation. A minimum thickness of 5.8 m was measured for Lithofacies 4 in the Gamka Formation.

Process Interpretation: HCS and SCS are “scour and drape” bedforms which often co-occur and are thought to form as a product of oscillatory flow (produced by strong wave action) on shorefaces, or by offshore unidirectional flow, or (as is most favoured) by combined oscillatory and offshore unidirectional flows between fair weather base and storm wave base (Dott and Bourgeois, 1982, Duke, et al., 1991, Cheel and Leckie, 1993, Leckie and Krystinik, 1989, Dumas and Arnott, 2006, Cummings, et al., 2009, Quin, 2011). A strong correlation exists between hummocky cross stratification only forming in silt to fine grained sand sediment and rare in coarser grained sediment (Leckie, 1988, Cummings, et al., 2009).

HCS and SCS share several key characteristics that are parsimonious with field observations in this study. The defining features of HCS and SCS have been outlined by Dott and Bourgeois (1982), Leckie and Krystinik (1989), Cheel and Leckie (1993), Dumas and Arnott (2006) and Quin (2011), and include: 1) HCS and SCS are large scale waveform sedimentary structures forming on erosional/scoured surfaces characterised by negative relief swales and positive relief hummocks; 2) Internal laminae are erosionally bounded and are low angled; 3) Internal laminae are approximately parallel to their accompanying lower bounding surfaces; 4) High degree in variance in dip directions of individual internal laminae and their associated scour surfaces. Thus internal laminae have no preferred orientation; 5) High degree of variance of individual lamina thickness. Individual lamina may thicken or thin laterally; 6) Upwards diminishment of dip angle of internal laminae, such that topmost laminae superficially appear planar; 7) Both bedforms appear to form as a result of oscillatory, or combined flows; 8) HCS and SCS may often

51 occur with otherwise parallel laminated beds, or planar laminated beds with some localised low hummocks and swales; 9) Wave ripple lamination, combined flow ripple lamination and sole marks, are often present within HCS and SCS beds. In idealised HCS beds, current ripple lamination and wave ripple lamination overlie hummocks, whilst sole marks are present at the base of HCS beds. Figure 3.28 is an idealised depiction of these features in a HCS bed.

The differences between HCS and SCS have been outlined by Dott and Bourgeois (1982), Leckie and Walker (1982), Leckie and Krystinik (1989), Cheel and Leckie (1993), Dumas and Arnott (2006) and Quin (2011). 1) SCS tend to be medium to coarse grained sandstone with pebbles whereas HCS often forms in silt to fine grained sandstones, being rare in coarse grained sandstones; 2) HCS internal laminae are usually < 10°, but may be up to 15° SCS internal laminae tend to be lower angled than HCS and never exceed 10°; 3) HCS have widths ranging from 1 to 5 m and amplitudes of up to 0.5 m. SCS beds are smaller than HCS and may be 0.5 to 2 m wide with smaller amplitudes than those expected for HCS beds; 4) SCS often truncates low angle to horizontal parallel lamination beds; 5) Geometry. HCS and SCS both form on a primary scoured surface and mimic the hummock and swale topography. Thus HCS beds are convex up and SCS beds are concave up.

It is possible that SCS beds may be lateral equivalents of higher angled cross stratified beds, but must not be confused with planar cross beds (Cheel and Leckie, 1993). SCS beds differ from trough cross beds in that internal laminae are inclined at an angle not associated with the angle of repose and that SCS bed deposits are associated with swales (Cheel and Leckie, 1993). SCS beds may be separate from HCS beds entirely and tend to vertically overly HCS beds in an idealised prograding sequence, but this may not always be present (Leckie and Walker, 1982, Cheel and Leckie, 1993, Dumas and Arnott, 2006). Dumas and Arnott (2006) favour sole SCS to form in more proximal environments.

A hypothetical continuum of isotropic to anisotropic HCS beds exist based on the symmetry of HCS beds based on the relative amount of oscillatory versus unidirectional flow action supplied within the basin at time of sedimentation (Duke, et al., 1991, Cheel and Leckie, 1993, Dumas and Arnott, 2006, Quin, 2011) (Fig. 3.29). Isotropic HCS beds tend to be symmetrical and appear to be accretionary bedforms with no sense of migration, or as single “scour and drape” bedforms. Isotropic HCS beds are favoured to form under pure oscillatory and oscillatory dominant combined-flow conditions (Duke, et al., 1991, Cheel and Leckie, 1993, Dumas and Arnott, 2006, Quin, 2011). Anisotropic HCS beds tend to be asymmetric and appear to have laminae oriented in a preferential dip direction which appear to migrate as large scale ripples (Cheel and Leckie, 1993, Dumas and Arnott, 2006, Quin, 2011). Anisotropic HCS beds have been demonstrated to form under combined flow by Duke et al. (1991), Dumas and Arnott (2006) and Quin (2011). Dumas and Arnott (2006) favour anisotropic HCS to be the most common HCS type to be encountered in the rock record. Anisotropic HCS beds are usually smaller than Isotropic HCS beds (usually > 5 cm thick) (Cheel and Leckie, 1993). Neither hypothetical bed could be identified with confidence in the study area.

Dumas and Arnott (2006) have calculated experimentally that oscillatory waves with long periods (8 – 10 sec) and velocities of 50 cm.sec-1 and oscillatory dominant combined flows with velocities of ~ 10 cm.sec-1 are required to form migrating anisotropic HCS beds.

3.4.4.1 Palaeocurrent Indicators Associated with Lithofacies 4

Leckie and Krystinik (1989) have illustrated several sedimentary structures associated with HCS and SCS bedding which may be used as palaeocurrent and palaeoshoreline indicators. These have been summarised in Figure 3.28. Only a few reliable wave ripple crests and parting lineations were obtained from HCS beds in Lithofacies Association B and C of the Gamka Formation and Lithofacies Association B of the Voorstehoek Formation at Grootrivierhoogte and are illustrated in Figures 3.30 – 3.32. Wave ripple crests from

Lithofacies Association B of the Gamka Formation show a north northeast to south-southwest trending ripple crest, whilst being more variable in the Voorstehoek Formation, with wave ripple crest trends oriented north-south, east-west and northeast-southwest at time of deposition. Parting lineation trends from Lithofacies Association B too show a large degree of variation among each other. Parting lineation trends at time of deposition show north-south, north-northeast – south south-west and east – west trends.

3.4.5 Lithofacies 5 (Coarse Grained Ripple Sandstone Lithofacies)

Description: Lithofacies 5 is a coarse grained ripple sandstone sensu Leckie (1988) and, in profile, has a relatively planar base overlain by a near symmetric and rounded ripple crest (Fig. 3.33 A), or a wavy base mimicking the underlying symmetric ripple crest bed and overlain by a near symmetric ripple crest (Fig. 3.33 B). Internal lamination within Lithofacies 5 is rarely preserved. Where internal laminae are preserved, laminae may be arranged in an overlapping chevron like fashion at ripple crests and troughs (Fig. 3.34 A). Alternatively laminae may be arranged as tangential cross laminae with an apparent preferred orientation (Fig. 3.34 B). Mudstone intraclasts and coalified plant debris may be present within the lithofacies and are often imbricate along internal lamina foresets (Fig. 3.34 C and D). This gives the appearance of intraclasts and plant debris tangentially feathering out from ripple crests. Individual beds range from 10 – 230 mm in thickness. Partings between individual beds are sharp.

Within the study area, coarse grained ripples of Lithofacies 5 have a variable crest to crest spacing (wavelength) between 125 and 280 mm with an average of 167 mm. Amplitudes of Lithofacies 5 (trough to ripple crest) range from 10 to 40 mm with 23 mm being most common. Ripple symmetry indices (stoss/lee slope) range from 1 to 2.1 with most averaging near symmetric values from 1.15 to 1.2. Ripple indices (wavelength/amplitude) range from 3.75 to 19 with an average of 7.4 to 7.8. At Grootrivierhoogte, wavelength values of Lithofacies 5 may extend from 2600 mm to 3500 mm with an

54 average are 2900 mm. This occurs without an appreciable change in amplitude. Amplitudes vary from 135 to 200 mm with an average of 168 mm. This change in thickness occurs with a change in bedform geometry such that the lithofacies has a superficial “whale back” hummock and swale type topography (Fig. 3.35).

In places, 3 dimensional exposures of Lithofacies 5 reveal that the lithofacies may possess another set of ripples giving it a combined flow ripple “ladderback” appearance. Crests of two sets of ripples are usually orthogonally arranged with respect to each other. Ripples may either be asymmetric current ripples (Fig. 3.35) or near symmetric and ladderback in form (Fig. 3.36). Characteristic features present on bedding planes of superficial “whale back” hummock and swale type topography Lithofacies 5 include mudstone intraclasts and red oxide nodules. These are preferentially concentrated in topographic lows in bedding plane exposures (Fig. 3.37 A). Red oxide nodules may weather out in profile as negative cavities; often leaving iron oxide rinds and/or staining halos (Fig. 3.37 B). It is possible that these nodules may have originally been pyritic and have later oxidised. Sandstone concretions are common on bedding planes and form as positive relief features (Fig. 3.37 C). The concretions are crudely round and may be found both within topographic highs and lows in Lithofacies 5. Prod marks, interpreted to be formed by plant debris are present often on the surfaces of sandstone concretions (Fig. 3.37 D), but may be found elsewhere on the bedding plane (Fig. 3.37 E). Conical and meandering rill marks are also present on the bedding plane of Lithofacies 5. Conical rill marks are confined to the leading edge of concretions and ripple crests, whilst meandering rill marks are present as scours directed into the bedding plane moving away from ripple crests/concretions (Fig. 3.37 F). Irregular blocks that appear to be angled into the bedding plane of Lithofacies 5 are common features which may be mistaken for sandstone concretions (Fig. 3.38). Some of these blocks appear to possess crude internal bedding, but it is hard to discern for certain. These have been interpreted to be sandstone intrabreccia’s Figure 3.39 displays all of these features co-occurring on a single bedding plane.

Body fossils within Lithofacies 5 are rare and are restricted to fragmentary coalified plant thalli and prod marks (Fig. 3.34 D and 3.37 E and D) with no metazoan fossils present. Ichnofossils within Lithofacies 5 are plentiful and resemble Skolithos sp. Within individual beds, Skolithos sp. penetrate both upper and lower bedding planes and are straight edged, parallel and perpendicular to bedding (Fig. 3.40 A). No backfill features are present within individual Skolithos sp. structures. Upper and lower bedding planes reveal “dimpled” surfaces marking individual Skolithos sp. structures. Upper bedding planes are negatively dimpled (Fig. 3.40 B), whilst lower bedding planes are positively dimpled (Fig. 3.40 C).

Three variants of Lithofacies 5 have been recognised in the study area. These variants are based on ripple stacking across partings of individual beds, the degree of amalgamation and the presence of finer grained clastic material at partings. All three variants may co-exist and grade into each other.

Across partings between individual beds of Lithofacies 5, ripple crests may stack on top of one another, usually in phase, to out of phase where overlying ripple crests lie on top of underlying ripple troughs (Fig. 3.33 A and B). It has been noted that flat based variants of Lithofacies 5 tend to stack in phase, to slightly out of phase (Fig. 3.33 A), whilst wavy based variants always are out of phase (Fig. 3.33 B). Both variants show no degree of amalgamation as individual bedding planes and partings are clearly defined. This variant of Lithofacies 5 is most common and is present at all three sections of the Gamka Formation within the study area.

The second variant of Lithofacies 5 is an amalgamated variant. Here individual partings and bedding planes may be undetectable. Individual beds cut into each other making the Lithofacies strongly amalgamated in appearance (Figs 3.41A and B, Fig 3.42). “Amalgamated sets” vary from 600 to 800 mm in thickness but may exceed 2.0 m. This variant of Lithofacies 5 is present at all three sections.

The third variant of Lithofacies 5 contains silt to fine grained sandstone intercalations between individual coarse grained ripple beds. Coarse grained ripple beds of this variant of Lithofacies 5 may superficially to be mistaken for sediment starved sandstone ripple flasers within a lenticular laminated heterolithic mudstone (Fig. 3.43 A). These are horizontally continuous and not discontinuous as would be typical for true sediment starved flasers within lenticular laminated heterolithic beds (Fig. 3.43 B). This variant has been recognised only in the Gamka Formation at Wupperthal where it has gradational contacts with the second variant of Lithofacies 5.

Lithofacies 5 usually is found in close association with Lithofacies 4 in the Gamka Formation (as Lithofacies Association C) with which it has sharp planar under and overlying contacts and is marked with an abrupt change in colour and sudden increase in coarse grained quartz sand content, see Lithofacies Association D. Where Lithofacies 7 overlies Lithofacies 5, the latter may intercalate with planar cross beds of Lithofacies 7, usually over a short interval of up to 1 m (Fig. 3.44).

The Lithofacies is Dark Yellowish Orange (10 YR 6/6) in outcrop with fresh cuttings being Greyish Orange Pink (5 YR 7/2).

Lithofacies 5 is a major component of the Gamka Formation at all three sections (Appendix 1 - 4). At Grootrivierhoogte, Lithofacies 5 attains a maximum thickness of 45.0 m. At Keurbosfontein, it attains a maximum thickness of 36.0 m and a minimum thickness of12.0 m. At Wupperthal, Lithofacies 5 attains a maximum thickness of 18.0 m and a minimum thickness of 5.0 m. Lithofacies 5 may be even thicker at Wupperthal, but due to scree obscuring outcrop, this remains unknown.

Process Interpretation: Symmetric coarse grained ripples senso stricto have been described by Leckie (1988). In profile, symmetric coarse grained ripples are almost symmetric to slightly asymmetric, as evidenced by internal laminae often being set up in a preferred dip orientation. Leckie (1988) considers these sedimentary structures to form under combined flow conditions. This is

57 supported by reported ripple symmetry indices ranging from 0.8 to 1.5 suggesting that these ripples formed under oscillatory flow conditions with a component of unidirectional flow (Leckie, 1988).

It has been suggested and demonstrated experimentally that wave conditions creating coarse grained ripples are almost identical to those which produce hummocky cross stratification (Leckie, 1988, Cummings et al., 2009). Here oscillatory flow acting upon coarse grained sediment will produce coarse grained ripples and oscillatory flow acting upon finer grained sediment would produce hummocky cross stratification (Leckie, 1988, Cummings, et al., 2009). The presence of mudstone intraclasts and coalified plant debris may be indicative of traction transport of rip up intraclasts and plant fragments during storm events.

3.4.5.1 Palaeocurrent Indicators Associated with Lithofacies 5

Leckie (1988) has outlined the use of coarse grained ripple crest trends as possible approximations of palaeo - shorelines. Reliable coarse grained ripple crest trends have been measured from Lithofacies 5 of the Gamka Formation at Grootrivierhoogte and Wupperthal. Crest trends for both study localities are presented in Figures 3.45 and 3.46. Ripple crest trends of Lithofacies 5 are difficult to distinguish from superimposed ladderback ripples; no distinction between the two was made when measured. At Grootrivierhoogte a dominant NE-SW trend and minor WNW-ESE and NNE-SSW trends are apparent. At Wupperthal, two equally weighted trends are present. These are a dominant NE-SW and minor NW-SE trend.

Current ripples were found on the top bedding plane of a “whale back” hummock and swale type bed within Lithofacies 5 in the Gamka Formation (Fig.3.47). These ripples indicate a southwestward directed palaeoflow direction.

3.4.6 Lithofacies 6 (Lenticular Wacke Lithofacies)

Description: Lithofacies 6 is characterised as lenticular wacke. Lithofacies 6 is poorly sorted and comprises silt to fine grained sand quartz with accessory mica. This lithofacies is present only in the Boplaas Formation. It is present at Keurbosfontein and Wupperthal and is suspected to be at Grootrivierhoogte, but is missing due to tectonism (Appendix 1 - 4).

Colouration of this lithofacies is variable with fresh exposures being Dark Grey to Medium Light Blueish Grey (10 Y 6/2) and Blueish Grey (5 PB 5/1). Weathered exposures of Lithofacies 6 may be Greyish Orange (10 YR 7/4), or Dark Greenish Grey (10 Y 6/1) as is most common.

At Keurbosfontein, Lithofacies 6 reaches a maximum thickness of 12.0 m and at Wupperthal reaches a maximum thickness of 15.0 m.

Lithofacies 6 is conformable and gradational with Lithofacies 3 at its lower contact over 2 – 3 m and is marked by a gradual increase in silt to fine grained sand quartz lenticles intercalating among Lithofacies 3. The Upper contact of Lithofacies 6 with Lithofacies 11 is also conformable and gradational over several metres and is marked by an increase in medium to coarse grained sand quartz and better sorting.

Internally, Lithofacies 6 consists of trough cross laminae at the base which become more planar until they are eventually planar horizontal towards the top of the lithofacies (Fig. 3.48). Alternatively, asymmetric climbing ripple cross lamination may overly trough cross laminae, often climbing in a preferential direction (Fig. 3.49), or by symmetric ripples (Fig. 3.50). Internally, Lithofacies 6 shows evidence for reactivation. This is demonstrated by sharp truncation surfaces which show evidence for erosion and incision of pre-existing internal sedimentary structures (Figs. 3.49 and 3.51). Lithofacies 6 co-occurs with Lithofacies 3. Contact between the two Lithofacies is sharp.

Lenticles of Lithofacies 6 have a depth measurements ranging from 0.15 m to 0.9 m, with widths from 1.5 m to in excess of 10.0 m. These compute to a range of width to depth ratios from as low as 3.3 to as high as 16.8 and an average of 11.4.

A broad two fold division of Lithofacies 6 may be recognised. This is based on the degree of amalgamation of lenticular bodies, overall amount of fine grained sand (represented by more numerous lenticles of Lithofacies 6) and weathering profiles. These are described as a: Lower sub-unit and Upper sub-unit. Lower and Upper sub-units are conformable and marked by an overall upward increase in sand content and decrease in the amount of intercalated Lithofacies 3.

Lower Single-story sub-unit: This sub-unit is present in the lower 3.0 to 9.0 m of Lithofacies 6 at Keurbosfontein and Wupperthal respectively (Appendix 1, 3 and 4). Lenticles of Lithofacies 6 co-occur within Lithofacies 3. Contact between Lithofacies 3 and 6 is marked by a sharp, curved and erosional lower surface and a sharp upper planar (Fig. 3.51), or rippled surface (Figs. 3.49 and 3.50), such that Lithofacies 3 totally surrounds Lithofacies 6. Numerous lenticles of Lithofacies 6 intercalate within Lithofacies 3 and hardly ever truncate each other (Fig. 3.52). Lenticles increase in frequency and size moving up the stratigraphic succession into the Upper sub-unit. Concomitantly, intercalated Lithofacies 3 decreases in thickness. The Lower sub-unit, owing to the presence of the more fissile Lithofacies 3, tends to weather negatively with respect to the overlying Upper sub-unit (Fig. 3.52). Lateral to Lithofacies 6, Lithofacies 3 may be climbing ripple cross laminated with a preferred orientation of foresets away from Lithofacies 6 (Fig.3.19). Here Lithofacies 3 eventually grades into a more planar parallel equivalent.

Upper Multi-story sub-unit: This sub-unit, present in the upper 6.0 to 9.0 m of Lithofacies 6, occurs at Wupperthal and Keurbosfontein respectively. The Upper sub-unit of Lithofacies 6 differs from the Lower sub-unit in that it has a more amalgamated geometry where individual lenticles of Lithofacies 6 may sharply truncate, or stack on top of each other (Fig. 3.53). Partings of

60 individual lenticles may be outlined by thin beds of Lithofacies 3. Lenticles are more numerous with respect to the Lower sub-unit. The tendency of the Upper sub-unit to be more amalgamated than the Lower sub-unit, as well as the greater amount of lenticles gives the Upper sub-unit a more positive weathering profile (Fig. 3.52).

Process Interpretation: The lenticular geometry and associated internal sedimentary structures of Lithofacies 6 is reminiscent of those expected for channels where migrating sediment laden fluids vertically incise substratum (in this case, Lithofacies 3) (Collinson and Thompson, 1989b, Collinson, 1996). Trough cross beds along channel margins may represent lateral accretion foresets along channel banks, inferred to be deposited perpendicular to flow and representative of channel migration or point bar deposition (Collinson and Thompson, 1989b, Collinson, 1996). The occurrence of planar parallel laminae, almost always overlying trough cross laminae, implies upper flow regime transport of sediment within the channel (Collinson and Thompson, 1989b, Collinson, 1996). Alternatively, due to the presence of mica within the lithofacies, flow conditions may have been gentle since mica’s inhibit ripple formation (Collinson, 1996). Sharp truncation surfaces within channels are interpreted to be indicative of periods of renewed erosion and incision (Collinson and Thompson, 1989b).

The isolated nature of individual lenticles of Lithofacies 6 within the Lower sub-unit and its non amalgamated geometry is reminiscent of single-story sandstones (Ribbon sandbodies, sensu Collinson, 1996). The Lower sub-unit is interpreted to be sedimentary infill of isolated channels. The lateral and vertical association of Lithofacies 3 may represent migration of Lithofacies 6 where Lithofacies 3 may be representative of aggradational/overbank deposits. This is demonstrated in Figure 3.19 where ripple climbing cross lamina with foresets dipping away from Lithofacies 6 and may indicative of overbank deposition, possibly levees which eventually grade into a more planar laminated equivalent.

The amalgamated nature of Lithofacies 6 at the expense of intercalated Lithofacies 3 in the Upper sub-unit is indicative of multi-story sandstones (Collinson, 1996). The Upper sub-unit is interpreted to represent numerous prograding channel systems eroding into aggradational/overbank deposits of Lithofacies 3. The truncation of individual channels by other channels may indicate periods of channel switching/migration or by either periods of renewed erosion.

The Lower sub-unit may be a distal equivalent of the Upper sub-unit due to the higher proportion of co-occurring fines present within the sub—unit (i.e. Lithofacies 3).

3.4.7 Lithofacies 7 (Macroform Cross-bedded Sandstone Lithofacies)

Description: Lithofacies 7 is a relatively well sorted, mature medium to coarse grained cross bedded sandstone. Accessory mudstone intraclasts and quartz extraclasts may locally be present (Fig. 3.54). The majority of cross beds are present as large (1.0 to 2.0 m thick) and laterally extensive (may be traced out in excess of 5.0 m in section) tabular and trough cross beds with minor zig-zag cross beds and occasional soft sediment deformed sandstone beds. Lithofacies 7 is restricted to the Hex River Formation at all three study sections and is present in the Gamka Formation at Keurbosfontein and Wupperthal (Appendix 1 - 4).

Colouration of this lithofacies is relatively uniform throughout the study area and varies from Greyish Orange Pink (5 YR 7/2) to sandy Yellowish Grey (5 Y 8/2) to Pale Yellow (5 Y 8/3, 5 Y 7/4). Fresh exposure of Lithofacies 7 may be Dark Greenish Grey (10 GY 4/2) to Light Greenish Grey (5 GY 8/1).

In the Gamka Formation, Lithofacies 7 is sharp and gradational with Lithofacies 5 along both its upper and lower contacts. Here, the two Lithofacies often co-occur with one another over a few metres (Fig.3.43). In the Hex River Formation, Lithofacies 7 sharply overlies Lithofacies 4 (Fig. 3.55). Lithofacies 7 frequently alternates with Lithofacies 8 in section (see

Lithofacies Association D) in the Gamka and Hex River Formations at all three sections and in the Boplaas Formation at Grootrivierhoogte. The contact and bounding surface between the two Lithofacies is sharp and planar (Fig. 3.56).

The attributes of individual cross beds shall be discussed separately.

Tabular cross beds:

Description: Tabular cross beds of Lithofacies 7 have straight crested dunes with high angled planar cross beds (with respect to intersecting bedding planes) (Fig. 3.57). Cross beds occur as sets which all dip in a preferred orientation Intersection of individual planar cross beds with respect to intersecting bedding planes on average ranges from 20 to 24° but may reach in excess of 40°. Individual cross beds commonly range in thickness from less than 10 mm to up to 20 mm. Sets of tabular cross beds (composed of numerous stacked planar cross beds) are variable in size and range from 0.85 – 1.7 m thick in the Hex River Formation and 0.05 to 0.3 m thick in the Gamka Formation. Numerous tabular cross beds may sharply stack vertically on top of one another with sharp planar, or tangential bounding surfaces which may be erosional (Fig. 3.58).

Process Interpretation: Tabular cross beds form by the down current migration of two dimensional, straight crested dunes by stoss erosion and lee deposition (Rubin, 1987, Collinson and Thompson, 1989c, Boggs, 2006). Tabular cross bedding forms under low to moderate flow velocities in lower flow regime conditions (Fr<1) (Boggs, 2006). Tangentially bounded tabular cross beds may be indicative of stronger flow conditions and separated flow in the lee slope area with eddying forming as a consequence (Collinson and Thompson, 1989c). Planar bounded tabular cross beds may be indicative of weaker flow conditions ( Collinson and Thompson, 1989c).

Trough cross beds:

Description: Trough cross beds of Lithofacies 7 have trough crested dunes with curved cross beds which merge tangentially with lower bedding planes and bounding surfaces (Fig. 3.59). These cross beds are present only in the Hex River Formation at all three study sections. Average dip of individual cross beds is between 18 - 20° and rarely exceeds 26° which is lower the dip of tabular cross beds. Sets of trough cross beds tend to be approximately 1.0 m thick in the Hex River Formation. Trough cross beds have a wedge shaped, bar form geometry and individual bars range in size from approximately 1.5 - 7.0 m and possibly greater. Numerous trough cross bed sets may stack laterally or vertically with respect to one another (Fig. 3.60). In vertical aspect, bedding planes of trough cross beds are curved and erosional since they scour into underlying beds.

Process Interpretation: Trough cross beds form by the down current migration of three dimensional, sinuous to linguloid crested dunes by stoss erosion and lee deposition and form under under moderate to high flow velocities in lower flow regime conditions (Fr<1) (Rubin, 1987, Collinson and Thompson, 1989c, Boggs, 2006). The tangential and erosionally based lower bedding planes of trough cross beds, suggest eddying and separated flow at the lee slope sides of trough cross bed sets (c. Collinson and Thompson, 1989).

Zig-zag cross beds:

Description: This bedform has been illustrated and described in Rubin (1987) (Figs. 59 and 60, pp: 142 – 143) as “zig-zagging cross bedding” and form by alternate stacking of spur and scour pits caused by current reversals. This structure may be mistaken for herringbone cross stratification, but differs in the angle of reactivation surfaces. Herringbone cross stratification is characterised as having low angle, planar reactivation surfaces, where as zig- zag cross beds are higher angled. Zig-zag cross beds of Lithofacies 7 are characterised as mutually bimodal oriented planar cross beds which sharply

truncate each other along a high angled and erosionally surfaced plane, or reactivation surface (Fig. 3.61). Zig-zag cross beds occur at only a few places in the study area and are present in both the Gamka and Hex River Formations. Individual zig-zag bed sets may reach up to 0.4 m thick in the Hex River Formation and up to 0.3 m thick in the Gamka Formation. Cosets of zig-zag cross beds may have a thickness of 1.0 m and a lateral continuity of 1.5 to 2.0 m in outcrop. Where present these beds tend to sharply co- occur with tabular cross bedded units, or Lithofacies 8 (Fig. 3.61).

Process Interpretation: Bipolar current directions at alternate times may cause sets of planar cross beds to stack on top of and migrate in opposite directions to one another, thus forming zig-zag cross bedding (Rubin, 1987). Alternatively, this structure may be formed by alternate stacking of spurs and scour pits, or by vertical stacking of trough shaped sets of cross beds, or long term shifts in channel geometry of ebb and flood dominated channels (Rubin, 1987).

Soft sediment deformation in sandstones: Soft sediment deformation is present at a few places in Lithofacies 7 and includes slumped and massive sandstones, recumbent cross bedded sandstones. Slumped sandstone beds with distorted to no visible internal laminae/stratification (Fig. 3.62 A) and may be up to 3.0 m thick. Recumbent cross bedded sandstones are also a common soft sediment deformation feature in Lithofacies 7 and is most common in cross bedded units (Fig. 3.62 B). Here cross beds have exaggerated dip angles greater than 45°.

Process Interpretation: Slumped and massive sandstone beds may indicate rapid sedimentation from a grainflow or fluidised flow (Collinson and Thompson, 1989c, Boggs, 2005). Intense burrowing and reworking by plants and animals may also destroy any primary bedding, rendering the unit to appear massive (c. Collinson and Thompson, 1989, Boggs, 2005). Liquefaction of sediment, due to shock or dewatering of sediment may also produce slumped and massive sandstone beds (Collinson and Thompson, 1989c, Boggs, 2005).

Recumbent cross bedding is formed by shearing and liquefaction of water logged and loosely packed subaqueous dunes by water currents passing over their crests (Nichols, 2009).

3.4.7.1 Palaeocurrent Indicators Associated with Lithofacies 7

Palaeocurrent indicators from Lithofacies 7 are given as a cumulative for cross beds. Reliable palaeocurrent indicators at time of sedimentation from cross bed dip directions were obtained from the Gamka Formation (from Keurbosfontein and Wupperthal), the Hex River Formation (from all three study sections) and the Boplaas Formation (at Grootrivierhoogte). Figures 3.63 – 3.64 depict cross bed dip directions from the Gamka Formation at Keurbosfontein and Wupperthal respectively. Figures 3.65 – 3.67 depict cross bed dip direction orientations from the Hex River Formation from Grootrivierhoogte, Keurbosfontein and Wupperthal respectively. Figure 3.68 depicts cross bed dip directions from the Boplaas Formation at Grootrivierhoogte. Palaeocurrent data from associated asymmetric current ripples beds from Lithofacies 7 have also been recorded. Asymmetric current ripple data has been recorded in the Gamka Formation at Keurbosfontein and the Hex River Formation and Boplaas Formations at Grootrivierhoogte and are depicted in Figures 3.69, 3.70 and 3.71 respectively.

Overall palaeocurrent data from the Gamka Formation suggest a dominant south –west to south-southwest directed palaeoflow direction for Lithofacies 7. In the Hex River Formation a dominant southwest to south-southwest directed palaeoflow direction for Lithofacies 7 at Grootrivierhoogte and Keurbosfontein at time of sedimentation. At Wupperthal, a co-dominant west-southwest and south-southeast directed palaeoflow direction was recorded with less dominant, but still significant southwest to south-southwest directed palaeoflow directions being evident at time of sedimentation. At all three sections, variable

subordinate palaeoflow directions are directed west-northwest to northwest and north with fewest palaeoflow directions trending north- northeast, northeast and southeast for the Hex River Formation. In the Boplaas Formation at Grootrivierhoogte a solitary reliable tabular cross bed dip direction was recorded. This indicated a south-westward directed palaeoflow direction at time of sedimentation.

Asymmetric current ripple data from a solitary asymmetric ripple laminated bed suggest a south-southwest directed palaeoflow direction at time of sedimentation for Lithofacies 7 of the Hex River Formation. In the Boplaas Formation, a south-southeastward directed palaeoflow is suggested at time of sedimentation based on a solitary asymmetric ripple laminated bed associated with Lithofacies 7.

3.4.8 Lithofacies 8 (Planar Parallel-bedded Sandstone Lithofacies)

Description: Lithofacies 8 is a well sorted, mature medium to coarse grained planar parallel – bedded sandstone. Individual beds range from 200 to 400 mm in thickness and consist of internal planar laminae that are approximately 10 mm thick (Fig. 3.72). Thin ripple laminae may overly the tops of otherwise planar laminated beds of Lithofacies 8, or internally be entirely composed of them (Fig. 3.73 A). Parting lineations may be present on the top surfaces of bedding planes of Lithofacies 8 (Fig. 3.73 B).

Lithofacies 8 is variably present within the Boplaas, Hex River and Gamka Formations in the study area. In the Gamka and Hex River Formations at all three sections, Lithofacies 8 generally alternates with Lithofacies 7 and in the Boplaas Formation at Grootrivierhoogte (see Lithofacies Association D) (Fig. 3.74) (Appendix 1 - 4). In the Boplaas Formation, Lithofacies 8 may sharply under, or overly Lithofacies 11 and 12 (see Lithofacies Association H) at Keurbosfontein and Wupperthal (Appendix A 1, 3 and 4).

In the Gamka and Hex River Formations, fresh cuttings of Lithofacies 8 may be Dark Greenish Grey (5 GY 4/1), whilst weathered outcrop exposures may

be Greyish Orange (10 YR 7/4) to Pale Yellow (5 Y 7/4, 5 Y 8/3). In the Boplaas Formation, fresh cuttings of Lithofacies 8 vary from Pale Green (5 G 8/2) to Light Greenish Grey (7.5 GY 7/4).

Body fossils in the Lithofacies 8 are restricted to fragmentary plant compression fossils. These fossils are randomly arranged on the top bedding planes of Lithofacies 8 and are most common in the Boplaas Formation at Grootrivierhoogte and Keurbosfontein (Fig. 3.74). In the Hex River Formation, fragmentary plant compression fossils were found only at Wupperthal. Bioturbation is rare, but is common when in close association with fragmented plant fossil compressions (Fig. 3.75). These give the sandstones a “mottled” appearance since the original internal laminae are broken. Ichnofossils may also be present in Lithofacies 8 but only penetrate the top bedding planes of Lithofacies 8 (Fig. 3.76 A). These ichnofossils, which are variable in size, resemble plant rootlets and are simple, curved, and taper towards a point with no evidence for bifurcation. Upper bedding planes have a dimpled negative topography in areas where these ichnofossils have penetrated the bedding plane (Fig. 3.76 B).

At Grootrivierhoogte Lithofacies 8 is present only in the Boplaas and Hex River Formations (Appendix 1,2). Here Lithofacies 8 attains a maximum thickness of 4.5 m thick in the Hex River Formation. Lithofacies 8 in the Boplaas Formation attains a minimum thickness of 4.5 m and a maximum of 9.0 m.

At Keurbosfontein Lithofacies 8 is present within the Gamka, Hex River and Boplaas Formations (Appendix 1,2). Lithofacies 8 attains a maximum thickness of 5.0 m. Within the Hex River Formation, Lithofacies 8 attains a minimum thickness of 13.0 m and a maximum of 45.0 m thick.

At Wupperthal, Lithofacies 8 is present only in the Hex River and Gamka Formations (Appendix 1 and 4). Within the Gamka Formation, Lithofacies 8 attains a maximum thickness of 4.0 m. In the Hex River Formation Lithofacies 8 attains a minimum thickness of 0.8 m and a maximum of 5.0 m.

Process Interpretation: Planar parallel - bedding in medium to coarse grained sandstones forms under high flow velocity conditions in upper flow regime conditions (Fr>1) where sediment is cannibalised from pre formed ripples and dunes (Collinson and Thompson, 1989c, Boggs, 2006). The presence of thin ripples and parallel laminae may indicate slight mica content to the original sediment composition and not lower flow regime conditions (Collinson and Thompson, 1989c). Alternatively, laminae may reflect slight differences in grain size, or slight differences in flow velocity close to the bedding plane (Collinson and Thompson, 1989c).

3.4.8.1 Palaeocurrent Indicators Associated with Lithofacies 8

Palaeocurrent indicators at time of sedimentation for Lithofacies 8 comes from parting lineation trend data and lee slope orientation of asymmetric current ripples.

Parting lineations in association with Lithofacies 8 have been recorded in the Gamka Formation at Wupperthal (Fig. 3.77), the Hex River Formation at Keurbosfontein and Wupperthal (Figs. 3.78 and 3.79) and the Boplaas Formation at Grootrivierhoogte (Fig. 3.80). Reliable asymmetric current ripple palaeocurrent data has been recorded from the Hex River Formation at Keurbosfontein only (Fig. 3.81).

A north-northeast to south-southwest trending palaeoflow at time of deposition was recorded from a parting lineation in the Gamka Formation at Wupperthal. In the Hex River Formation, a north- northeast to south-southwest trending palaeoflow at time of deposition and an east-northeast to west-southwest trending palaeoflow at time of deposition were recorded at Grootrivierhoogte and Wupperthal respectively. A northwest to southeast palaeoflow trend at time of deposition was recorded in the Boplaas Formation at Grootrivierhoogte.

Asymmetric current ripple data from the Hex River Formation at Keurbosfontein indicate a south-southwest directed palaeoflow at time of sedimentation.

3.4.9 Lithofacies 9 (Massive Bioturbated Sandstone Lithofacies)

Description: Lithofacies 9 is characterised as being a massive poorly sorted medium to coarse grained sandstone with pervasive bioturbation. Mud and white mica are present within the lithofacies as significant accessory minerals. Internal bedding is completely broken as a result of bioturbation (Fig. 3.82 A and B). Vague planar parallel and ripple beds may be visible, discrimination of either is dubious. Individual beds, representing inferred pulses of sedimentation when discernible are generally tabular in geometry with lumpy, amorphous bedding planes with sharp contacts and may stack on top of each other (Fig. 3.81 C and D). Individual beds have an average thickness ranging in size from 10 – 30 mm but may thicken to 150 – 300 mm.

Lithofacies 9 marks the base of the Boplaas Formation at all places where sections were measured. The base, sensu stricto, of the Boplaas Formation is taken to be the first laterally persistent sandstone above the Tra-Tra Formation (Basson, et al., 1995) and is present at all three localities (Appendix 1 - 4). The thickness of Lithofacies 9 at the base of the Boplaas Formation is relatively uniform within the study area. At Grootrivierhoogte, Lithofacies 9 attains a maximum thickness of 31.5 m. At Keurbosfontein, the lithofacies reaches a maximum thickness of 34.5 m. At Wupperthal, Lithofacies 9 reaches a maximum thickness of 33.1 m. Lower and upper contacts of Lithofacies 9 are conformable and gradational with Lithofacies 3. This is marked by an increase in coarse grained sand quartz and silt respectively. At Keurbosfontein, a solitary bed of Lithofacies 9 is present and overlies Lithofacies 6 and is 7.4 m thick. This contact is sharp, planar and conformable (Fig. 3.83). The upper contact of this bed is conformable and gradational with Lithofacies 3.

In outcrop weathered exposures of Lithofacies 9 range in colour from Pale Olive (5 Y 6/4), Olive (5 Y 5/3), Light Brownish Grey (5 YR 6/2) to Dark Greenish Grey (2.5 GY 4/1). Weathered exposures vary from Blueish White (10 Y 8/1), Olive Grey (7.5 GY 4/2), Light Blueish Grey (5 B 8/1, 10 B 8/1), Greyish Blue (5 PB 5/2), Blueish Grey (7.5 GY 5/2, 10 GY 7/2), Pale Yellow (5 Y 7/4), Light Brown (5 YR 6/4), Yellow Brown (10 YR 5/8), Pale Yellowish Brown (10 YR 8/6), Pale Red (10 R 6/2) to Greenish Grey (5 G 5/2, 5 Y 7/2).

Lithofacies 9 contains numerous coalified plant body fossils. Fresh examples have a characteristic silver sheen, whilst weathered examples typically are dull brown with oxidation. These may be either solitary or branched plant axes (Fig. 3.84). Coalified plant axes are usually present on bedding planes, but may be oriented discordant to bedding. At Grootrivierhoogte, molds of spiriferid brachiopod valves, lophophores and their supporting brachidia were found (Fig. 3.85 A, B). These spiriferids are considered to occur in their apparent life position based on two lines of evidence, 1) pedicle valves are most abundant (it is inferred that spiriferids rested on their pedicle valves on the substrate), 2) Valves occur in distinct clusters inferred to be communities (Fig. 3.85 B). Ichnofossils resembling Skolithos are also common in the lithofacies (Fig. 3.86). These ichnofossils, like those found in Lithofacies 8 are also straight edged ichnofossils which penetrate the top bedding planes of Lithofacies 9 leaving a characteristic dimpled negative topography on top bedding planes where ichnofossils penetrated the bedding plane in life (Fig. 3.86 A and B). Some ichnofossils, resembling rootlet ichnofossils appear to be originally organic rich. The presence of iron oxide staining around ichnofossils, suggest that they were originally pyritic and possibly formed in a dysaerobic, anoxic environment (Fig. 3.86 C). These ichnofossils also tend to be curved like those observed in Lithofacies 8. Casts of meandering surface trails on top bedding planes of Lithofacies 9 are common (Fig. 3.87).

Process Interpretation: Persistent and en masse bioturbation may entirely destroy all forms of bedding and sedimentary structure (Reineck and Singh, 1975, Droser and Bottjer, 1985). This may be indicative of periods of relative calm in the depositional basin. According to Droser and Bottjer’s (1986) field

71 ichnofabric indicators (an index of the degree of bioturbation based upon primary bedding destruction), the bioturbation index of Lithofacies 9 ranges from 5 -6 since some bedding and sedimentary structures may be visible.

3.4.10 Lithofacies 10 (Epsilon Cross-bedded Sandstone Bar Lithofacies)

Description: Lithofacies 10 is characterised as being a medium to coarse grained, moderately well sorted epsilon cross bedded sandstone, sensu Miall (1985). Epsilon cross bed sets in turn are comprised of numerous trough cross beds which taper and merge tangentially with lower bedding planes and bounding surfaces, giving the lithofacies a characteristic bar-form geometry (Fig. 3.88). Average dip of trough cross beds, and hence lateral accretion bars ranges from as low as 5° to as high as 20° with 10° to 15° being most common. This was measured in the dip direction towards Lithofacies 11.

Lithofacies 11 overlies Lithofacies 10 and is also lateral to it in the dip direction of trough cross beds. Lithofacies 10 is more planar away from Lithofacies 11 in the opposite sense of dip of internal trough cross beds where it tapers along a planar upper bedding plane (Fig. 3.88). Numerous bars may stack laterally and vertically on top of one another. Where stacked vertically, bedding planes of bars are curved and erosional since they scour into one another. Heterolithic sandstone and mudstone beds may be present within individual epsilon cross bedded units, i.e. within sets of trough cross beds (usually overlying individual cross beds as normal graded beds), or between individual bars, i.e. between cosets of epsilon cross beds, or at the foot of bars where they merge with Lithofacies 11. Heterolithic sandstone and mudstone beds are typically thin ripple laminated.

Lithofacies 10 is entirely restricted to the Boplaas Formation at Keurbosfontein and Wupperthal only and is found in close association with Lithofacies 11. In cuttings with three dimensional exposures, Lithofacies 11 has a right angled relationship to Lithofacies 10 with respect to the sense of apparent progradation of the two Lithofacies. Individual bar thickness is variable through the bar, owing to its geometry. Here quoted maximum

72 thicknesses are from bedding plane to bar crest. Thicknesses range from 0.17 m to as thick as 0.9 m and on average is 0.55 m. Width of bars, measured from centre of Lithofacies 11 to end of bar, may extend in excess of 10 m.

In outcrop fresh cuttings of Lithofacies 10 vary from Pale Green (5 G 5/2), Pale Olive (5 Y 5/4), Greenish Grey (5 G 5/2), Olive Grey (5 Y 3/2), Medium Blueish Grey (10 BG 5/1) to Pale Reddish Brown (10 R 5/4). Weathered exposures of Lithofacies 10 may be Light Grey (5 Y 7/1), Medium Orange Pink (10 R 7/4) or Yellowish Brown (10 YR 5/5).

Numerous fragmentary coalified plant fossils are present within the Lithofacies moving up the stratigraphic section at Keurbosfontein and Wupperthal (Fig. 3.89). Fragmentary coalified plant fossils are more ubiquitous at Wupperthal than they are at Keurbosfontein.

Process Interpretation: Epsilon cross bedding is synonymous with lateral accretion bars commonly form on the point channel margin (cite of active channel margin depositon) of either relatively straight channels with sinuous gutters, or in meandering channels (Allen, 1970, Miall, 1985, Fielding, et al., 1993, Collinson, 1996, Huggett, 2007). Here sediment is deposited as a series of lower flow regime trough cross beds at an angle towards to the main channel flow direction. Sediment is either derived from cannibalisation of the cut bank or from headwaters or deposited in local shallow water areas (riffles) (Allen, 1970, Miall, 1985, Fielding, et al., 1993, Collinson, 1996, Huggett, 2007). Lateral accretion bar growth and migration tracks the migration of channels away from the point bank (Allen, 1970, Miall, 1985, Fielding, et al., 1993, Collinson, 1996).

3.4.11 Lithofacies 11 (Lenticular and Tabular Cross-bedded Sandstone Bar Lithofacies)

Description: Lithofacies 11 is a medium to coarse grained lenticular, or bar- form tabular cross-bedded sandstone depending on the direction which the lithofacies is viewed from.

Colour of Lithofacies 11 is similar to those reported for Lithofacies 10 and ranges from Pale Green (5 G 5/2), Pale Olive (5 Y 5/4), Greenish Grey (5 G 5/2), Olive Grey (5 Y 3/2), Dark Blueish Grey (5 PB 2.5/1), Medium Blueish Grey (10 BG 5/1), Blueish Grey (5 B 7/1, 5 PB 4/1) to Pale Reddish Brown (10 5/4). Weathered exposures of Lithofacies 11 may be Light Grey (5 Y 7/1), Greyish Orange Pink (5 YR 7/2), Medium Orange Pink (10 R 7/4) or Yellowish Brown (10 YR 5/5).

As in Lithofacies 10, fragmentary coalified plant fossils are numerous in Lithofacies 11 at both Keurbosfontein and Wupperthal in the Boplaas Formation. At Wupperthal, fragmentary coalified plant fossils are more ubiquitous than at Keurbosfontein (Fig. 3.90).

Lithofacies 11 always has erosional contacts, individual beds may downcut and incise into each other (Fig.3.91, 3.92, 3.94 and 3.95), or into Lithofacies 10 (Figs. 3.88 and 3.98). Heterolithic sandstone and mudstone beds may overlie Lithofacies 11; this is usually marked by a gradational contact and is accompanied by an increase in mud and decrease in overall sand content.

Due to the difference in bed geometry between lenticular sandstone beds and tabular cross-bedded sandstone bars, the defining characteristics of each are discussed separately.

Tabular cross-bedded sandstone bars: In cuttings that are parallel to the inferred direction of bar migration and growth, Lithofacies 11 appears as a tabular cross-bedded sandstone bar. An idealised threefold sub-division of individual beds of these bars is possible based on their internal bedding

74 structures, all three of which may not always be present. These are a basal planar parallel laminated bed, overlain by a planar cross bed which is in turn overlain by a thin ripple bed. In this idealised representation of tabular cross- bedded sandstone bars, planar parallel laminated beds appear to be in the downcurrent direction of planar cross beds. The defining attributes of these sandstone bars is summarised in Figure 3.91 and 3.92. Thin ripple beds tend to be silt to fine grained sandstone in content and asymmetric with lee slopes dipping in the same direction as planar cross beds. Planar cross beds may have mudstone intraclasts present within foresets and are 10 to 30 mm thick per foreset. Figure 3.93 depicts these additional internal bedding structures in detail. Planar parallel laminated beds possess individual laminae that are 2 to 5 mm thick.

Numerous tabular cross-bedded sandstone bars may stack on top of each other in section (Fig. 3.94). Contact between individual bars is sharp and planar (Fig. 3.95). Here planar cross beds may entirely truncate thin ripple beds and overly underlying planar cross beds, or directly overly thin ripple beds (Fig. 3.91).

Tabular cross-bedded sandstone bars are present in the Boplaas Formation only at Keurbosfontein and Wupperthal in association with Lithofacies 10 (see Lithofacies Association H) and the lowermost Tra-Tra Formation, where it is in association with Lithofacies 3 (see Lithofacies Association E), at all three study sections. Average thickness of bars in the Boplaas Formation is variable based on whether all three recognised subdivisions are present in one bed. As such, thicknesses presented here are average ranges for respective elements. Thin ripple beds: 10 to 20 mm. Planar cross beds: 0.2 to 0.35 m is common, but may be as thick as 0.4 to 0.45 m. Average foreset thicknesses range from 10 to 30 mm. Planar parallel laminated beds range in thickness from 0.3 to 0.57 m. Average thicknesses of individual laminae are between 2 to 5 mm. Thus, expected thicknesses for complete bars in the Boplaas Formation range from a minimum thickness of 0.5 m to a maximum of 1.0 m and an average of 0.7 m. In the Tra-Tra Formation, tabular cross- bedded sandstone bars sharply intercalate among and truncate Lithofacies 3

(see Lithofacies Association E) (Fig. 7.4.11.7). Internal laminae are composed of numerous centimetre scale dunes. Average thickness for individual tabular cross-bedded sandstone bars in the Tra-Tra Formation ranges from 0.02 to 0.6 m.

Ichnofossils superficially resembling Diplocraterion sp. have been found in tabular cross-bedded sandstone bars in the Tra-Tra Formation at Grootrivierhoogte. These ichnofossils are inclined in a southward dip direction and are characterised as being relatively straight edged with a rounded terminus and internal backfill which penetrates both upper and lower bedding planes (Fig. 3.97). The top bedding plane is characterised as being dimpled where individual ichnofossils penetrated the bedding plane in life. Ichnofossils decrease in abundance moving up the stratigraphic section of Lithofacies Association E until being absent in the uppermost beds.

Lenticular sandstone: In cuttings that are perpendicular to the inferred direction of bar growth, Lithofacies 11 appears as a lenticular bed. Individual lenticular beds have a curved, lensoidal, channel form geometry. Internally, lenticular beds are trough cross bedded with foresets oriented roughly perpendicular to those of tabular cross-bedded sandstone bars (Fig. 3.98). When found in association with Lithofacies 10, lenticular beds are always lateral to and overlie Lithofacies 10 (Fig. 3.88). With reference to Figure 3.88, when looking down the inferred channel axis (hence perpendicular to internal trough cross laminae), lenticular beds have a characteristic “ribbon” shape (sensu Collinson, 1996) and consist of a central lenticular sandstone with lateral “wings” of sandstone. Width to thickness ratio’s of the main lenticular sandstone (i.e. excluding the lateral wings of sandstone) commonly are less than 15, and range from 4.0 to 10.63 and rarely may be as large as 19.3 and on average 8.57. Lenticular beds have only been found in association with tabular cross-bedded sandstone bars in the Boplaas Formation at Keurbosfontein and Wupperthal. Lenticular beds, as mentioned previously are found in close association with Lithofacies 10 in the Boplaas Formation at Keurbosfontein and Wupperthal as Lithofacies Association H.

Process Interpretation: As previously mentioned, tabular cross-bedded sandstone bars and lenticular sandstones are co-occurring features. These are both associated with channel fill deposits (Collinson and Thompson, 1989c, Collinson, 1996, Miall, 1985). Tabular cross-bedded sandstone bars are explained as migration of subaqueous dunes within channels, i.e. as downstream accretion bars. These are lower flow regime dunes. The three fold sub-division of tabular cross-bedded sandstone bars into basal planar parallel laminated beds, upwards into planar cross beds and eventually overlain by a thin ripple bed may be indicative of an idealised upward decrease in flow energy with emergence due to gradual channel fill and concomitant decreasing water depth (Collinson and Thompson, 1989c, Collinson, 1996). Here basal planar parallel laminated beds may be indicative of upper flow regime conditions which grade into lower flow regime planar cross bedded dunes and eventually thin ripple laminae. Trough cross beds present at the sides of lenticular sandstones may possibly represent side channel fill, or small scale lateral accretion foresets and may represent channel migration or point bar deposition (Collinson and Thompson, 1989c, Collinson, 1996, Miall, 1985). The lateral “wings” of sand may be representative of levee/overbank deposits (Collinson, 1996).

3.4.11.1 Palaeocurrent Indicators Associated with Lithofacies 11

Reliable palaeocurrent indicators at time of deposition for Lithofacies 11 come from lee slope dip directions of tabular cross-bedded sandstone bars and approximate gutter trends of lenticular sandstones. These palaeocurrent indicators were recorded for Lithofacies 11 in the Boplaas Formation at Keurbosfontein and Wupperthal and are respectively presented in Figures 3.99 and 3.100 for cumulative lee slope dip directions for tabular cross-bedded sandstones. Figures 3.101 and 3.102 illustrate gutter trend approximations from lenticular sandstone beds of Lithofacies 11 at Keurbosfontein and Wupperthal respectively. In the Tra-Tra Formation at Grootrivierhoogte, palaeocurrent indicators at time of deposition were obtained from lee slope dip directions from associated asymmetric current ripples (Fig.

3.103), furrow trends from apparent rib and furrow structures (Fig. 3.104), as well as a solitary parting lineation trend (Fig. 3.105).

Cumulative lee slope dip directions from tabular-cross bedded sandstones of Lithofacies 11 from the Boplaas Formation at Keurbosfontein and Wupperthal indicate a dominant southwest to south-southwestward and a less dominant south to south- southeastward palaeoflow direction at time of deposition, with subordinate west-southwestward and south-eastward directed palaeoflow directions at Keurbosfontein. Cumulative gutter trend approximations suggest a northeast southwest to north-northwest south-southeast trending palaeoflow for Lithofacies 11 at Keurbosfontein with less dominant north south, north-northeast south- southwest and least dominant northwest southeast trending palaeoflows present at time of deposition. At Wupperthal, a dominant north-northeast south south-west trending palaeoflow is evident with less dominant north-northeast south southwest and northwest southeast trending palaeoflow trends present at time of deposition.

Lee slope dip directions from asymmetric current ripples on the top bedding plane of a tabular cross-bedded sandstone bar of Lithofacies 11 from the Tra-Tra Formation indicated a dominant south- southwestward and less dominant southward directed palaeoflow directions. Approximate furrow trends from approximate rib and furrow sedimentary structures in a tabular cross-bedded sandstone bar of Lithofacies 11 from the Tra-Tra Formation at Grootrivierhoogte indicate a dominant northwest southeast to north-northwest south south-east trending palaeoflow at time of deposition. A solitary parting lineation trendfrom a tabular cross bedded sandstone bar of Lithofacies 11 suggest a north-west south-east trending palaeoflow at time of deposition.

3.4.12 Lithofacies 12 (Carbonaceous Mudstone Lithofacies)

Description: Lithofacies 12 is a carbonaceous mudstone. Carbonaceous mudstones are not ubiquitous in the study area and have been found only within Lithofacies Association H of the Boplaas Formation at Wupperthal and Lithofacies 5 of the Gamka Formation at Keurbosfontein.

The occurrence of carbonaceous mudstones coincides with a local abundance of coalified plant material and fine grained silt and clay.

Colour of Lithofacies 12 in outcrop is Black to Dark Grey and Light Grey (5 Y 7/1) to White (8/N) where sufficiently oxidised.

Within the Boplaas Formation, individual beds are laterally continuous and range in thickness from 0.02 to 0.45 m with sharp planar upper and lower contacts with Lithofacies Association H (Fig. 3.106). In the Gamka Formation, Lithofacies 12 is non strataform with amorphous contacts with Lithofacies 5 and reaches a maximum thickness of 0.3 m (Fig 3.107).

Process Interpretation: Carbonaceous mudstones form by the accumulation of organic matter in environments which favour a rapid burial of organic matter (Boggs, 2005). Burial and accumulation of organic material ideally should be at a rate greater than the rate of decomposition by bacterial and/or chemical processes such as oxidation (Boggs, 2005).

3.5 Lithofacies Associations

A total of 8 Lithofacies Associations are recognisable in the study area. Lithofacies Associations A – D are restricted to the Gydo – Gamka and Voorstehoek – Hex River Systems. Lithofacies Associations E – H are restricted to the Tra-Tra – Boplaas System. These are summarised in Table 3.2.

3.5.1 Lithofacies Association A

Description: Lithofacies Association A (L.A. A) is characterised by laterally extensive lower flow regime plane bed laminated to locally massive clayshales (Lithofacies 1) which coarsen upward into lower flow regime plane bed laminated to ripple bedded upper siltstones (Sub-lithofacies 2.1 and 2.2 respectively).

This lithofacies association is entirely restricted to the Gydo and Voorstehoek Formations at all three sections within the study area (Appendix 1 - 4). Within the Gydo Formation, both Lithofacies are expressed, whereas Lithofacies 2 is only expressed in the Voorstehoek Formation.

LA.A is laterally continuous throughout the outcrop area in the Gydo and Voorstehoek Formations. At Grootrivierhoogte, L.A. A reaches a maximum thickness of 134 m in the Gydo Formation. In the Voorstehoek Formation at Grootrivierhoogte, L.A. A attains a maximum thickness of 10.5 m and a minimum thickness of 6.0 m. At Keurbosfontein, L.A. A attains a maximum thickness of 109 m. L.A. A attains a maximum thickness of 37.5 m and minimum thickness 6.0 m in the Voorstehoek Formation at Keurbosfontein. At Wupperthal, L.A. A attains a maximum thickness of 52.5 m in the Gydo Formation and 12.0 m in the Voorstehoek Formation with a minimum thickness of 1.5 m.

3.5.2 Lithofacies Association B

Description: Lithofacies Association B (L.A. B) is characterised by lower and laterally extensive heterolithic silty sandstone and mudstone beds (Lithofacies 3) which are intercalated by silt to fine grained non-amalgamated HCS and SCS sandstone beds (Lithofacies 4) in its upper extent. HCS and SCS sandstones are present as discontinuous lensoidal beds with sharp upper and lower contacts (Fig. 3.108). Lensoids attain a maximum thickness between 0.2 and 0.4 m and may laterally extend to 9 m in outcrop.

L.A. B is present in the Gydo, Gamka and Voorstehoek Formations at all three sections in the study area (Appendix 1 - 4). In the Gydo Formation, L.A. B gradationally overlies Lithofacies Association A. In the Voorstehoek Formation, L.A. B both under and overlies L.A. A.

At Grootrivierhoogte L.A. B attains a maximum thickness of 36.0 m in the Gydo Formation, 3.0 m in the Gamka Formation and 39.0 m in the Voorstehoek Formation. Minimum thicknesses of L.A. B were 2.5 m for the Gamka Formation and 2.5 m for the Voorstehoek Formation. At Keurbosfontein L.A. B attains a maximum thickness of 7.5 m in the Gydo Formation, 10.5 m in the Gamka Formation and 35.1 m in the Voorstehoek Formation. A minimum thickness of 7.5 m for L.A. B in the Voorstehoek Formation at Keurbosfontein was recorded. At Wupperthal, L.A.B attains a maximum thickness of 10.7 m in the Gydo Formation, 6.0 m in the Gamka Formation and 21.0 m in the Voorstehoek Formation. Minimum thicknesses of 3.7 m and 9.5 m for Lithofacies Association B were recorded in the Gamka and Voorstehoek Formations respectively.

In the Gamka Formation, L.A. B both under and overlies L.A. C. This contact is conformable, sharp and erosional and is marked by a gradual decrease in Lithofacies 3 as well as an overall increase in the average size and lateral continuity of intercalating beds of Lithofacies 4 when moving from L.A.B to C.

3.5.3 Lithofacies Association C

Description: Lithofacies Association C (L.A. C) is characterised as being a lower and laterally extensive amalgamated silt to fine grained HCS and SCS sandstone (Lithofacies 4) and upper symmetric rippled coarse grained sandstone (Lithofacies 5). Lithofacies 3 may be present around individual beds of Lithofacies 4 but is rare.

This lithofacies association is restricted to the Gamka and Hex River Formations at all three sections within the study area (Appendix 1 - 4). In the

Gamka Formation, both Lithofacies 4 and 5 are expressed whereas only Lithofacies 4 is present in the Hex River Formation.

At Grootrivierhoogte L.A. C attains a maximum thickness of 49.9 m and 6.0 m for the Gamka and Hex River Formations respectively. A minimum thickness of 12.5 m was recorded for L.A. C in the Gamka Formation. At Keurbosfontein L.A. C attains a maximum thickness of 66.0 m in the Gamka Formation and 16.2 m in the Hex River Formation. A minimum thickness of 9.0 m was recorded for L.A. C of the Gamka Formation. At Wupperthal L.A. C attains a maximum thickness of 26.5m in the Gamka Formation and 4.2 m in the Hex River Formation. A minimum thickness of 15.0 m was recorded for L.A. C in the Gamka Formation.

Contact with the overlying Lithofacies Association D is sharp, planar and conformable.

3.5.4 Lithofacies Association D

Description: Lithofacies Association D (L.A. D) is characterised by alternating medium to coarse grained sandstone beds of Lithofacies 7 and 8. Contact between individual beds is conformable and sharp.

This lithofacies association is present in the Hex River Formations at all three sections, the Boplaas Formation at Grootrivierhoogte and the Gamka Formation at Keurbosfontein and Wupperthal (Appendix 1 - 4).

At Grootrivierhoogte, L.A. D attains a maximum thickness of 47.5 m in the Hex River Formation and 46.3 m in the Boplaas Formation. At Keurbosfontein L.A. D attains a maximum thickness of 10.4 m and 91.0 m in the Gamka and Hex River Formations respectively. A minimum thickness of 3.0 m was recorded for L.A. D of the Gamka Formation. At Wupperthal L.A. D attains a maximum thickness of 8.1 m in the Gamka Formation and 42.1 m in the Hex River Formation. A minimum thickness of 4.0 m was recorded for L.A. D of the Gamka Formation.

Contact with the overlying Lithofacies Association E is sharp, planar and apparently paraconformable and is marked by an abrupt change to a dominantly argillaceous lithofacies association.

3.5.5 Lithofacies Association E

Description: Lithofacies Association E (L.A. E) is characterised as being a laterally extensive heterolithic silty sandstone and mudstone (Lithofacies 3) which is intercalated by laterally continuous medium to coarse grained tabular cross-bedded sandstone bars of Lithofacies 11 (Fig.3.109). Tabular cross- bedded sandstone bars of Lithofacies 11 have sharp planar to wavy upper contacts and sharp planar lower contacts with Lithofacies 3.

L.A. E has been recognised to occur only within the Tra-Tra Formation and is present at all three sections within the study area and may be traced along the entire study outcrop area (Appendix 1 - 4).

At Grootrivierhoogte L.A. E attains a maximum thickness of 16.5 m and a minimum thickness of 9.0 m. At Keurbosfontein L.A. E attains a maximum thickness of 19.5 m. At Wupperthal L.A. E attains a maximum thickness of 10.3 m.

Contact with the overlying Lithofacies Association F is gradual and is characterised by a sharp decrease in the amount of intercalating bars of Lithofacies 11 until it is entirely composed of Lithofacies 3.

3.5.6 Lithofacies Association F

Description: Lithofacies Association F (L.A. F) is characterised as consisting of lower and laterally extensive heterolithic sandstone and mudstone beds (Lithofacies 3) which coarsens into a laterally extensive and upper massive and heavily bioturbated coarse grained sandstone (Lithofacies 9).

L.A. F traverses the recognised Tra-Tra/Boplaas contact with more argillaceous parts of the lithofacies association wholly within the Tra-Tra Formation, with the base of Lithofacies 9 marking the base of the Boplaas Formation and may be traced along the entire study outcrop area (Appendix 1 - 4).

At Grootrivierhoogte L.A. F attains a maximum thickness of 67.5 m. At Keurbosfontein L.A. F attains a maximum thickness of 70.0 m. At Wupperthal L.A. F attains a maximum thickness of 58.7 m.

Contact with the overlying Lithofacies Association G is gradational and is marked by an overall upward fining and return to Lithofacies 3.

3.5.7 Lithofacies Association G

Description: Lithofacies Association G (L.A. G) consists of lower and laterally extensive heterolithic sandstone and mudstone beds (Lithofacies 3) which are intercalated by silt to fine grained sandstone lenticles which in turn become more amalgamated in their upper extent (Lithofacies 6).

L.A. G is entirely restricted to the Boplaas Formation and is present at Keurbosfontein and Wupperthal (Appendix 1 - 4). It is suspected that L.A. G may have been present at Grootrivierhoogte, but has been destroyed with tectonism as a synformal fold axis is present at the stratigraphic height where the lithofacies association should be thus rendering rocks of the lithofacies association unrecognisable. If this is so, L.A.. G may be traced along the entire study outcrop area

At Keurbosfontein L.A. G attains a maximum thickness of 31.4 m. At Wupperthal L.A. G attains a maximum thickness of 28.5 m.

At Keurbosfontein L.A. G is overlain by Lithofacies 9, this contact is sharp, planar and conformable. The overlying contact relationship at Wupperthal is unknown due to scree obscuring outcrop.

3.5.8 Lithofacies Association H

Description: Lithofacies Association H (L.F. H) variably consists of coarse grained sandstones of Lithofacies 8, 10 and 11. Contact between individual lithofacies is sharp and conformable.

L.A. H is present only in the Boplaas Formation at Keurbosfontein and Wupperthal. At Keurbosfontein, L.A. H attains a maximum thickness of 64.8 m and 83.0 m at Wupperthal.

L.A. H is underlain by Lithofacies 9at Keurbosfontein. This contact is sharp, planar and conformable. Underlying contact relationships at Wupperthal is unknown due to scree obscuring outcrop.

Chapter 4 Systematic Palaeontology

Several fossils were discovered in the course of fieldwork. Typical Malvinokaffric Realm invertebrate taxa are most numerous in the Ceres Subgroup and appear to be associated with particular lithofacies. The recognisible fossils were described as far as possible as they are useful for geological, stratigraphic and palaeo – environmental interpretation. Fossils from the study area are preserved as either molds, casts or within nodules. Authigenic shell preservation is absent. Fossil specimens commonly are well preserved with identifiable features.

Kingdom PLANTAE Subkingdom EMBRYOPHYTA Phylum TRACHAEOPHYTA Class, Family, Genus, Species indet.

Description: Plant fossils in the study area are commonly preserved as fragmentary elements and appear to be plant thalli or axes. Fossil thalli or axes most commonly are unbranched, with branched and bifurcating forms being rare (Fig. 3.84 A and B). Since plant fossils are fragmentary, size of plant fossils is variable and commonly between 10 to 100 mm and never exceeding 300 mm. Surface ornamentation is absent owing to the non-preservation of these features. Preservation of plant fossils may either be as coalified films on bedding planes (Fig.3.15), prod casts (Fig. 3.37 D and E) or as entire coalified fossils (Figs. 3.84 and 3.90).

Localities: Fossil plants are present in Lithofacies Association B in the Voorstehoek Formation at Keurbosfontein, Lithofacies Association C of the Gamka Formation at Grootrivierhoogte and Keurbosfontein, Lithofacies Association D in the Hex River Formation at Wupperthal and the Boplaas Formation at Grootrivierhoogte and Lithofacies Association H in the Boplaas Formation at Keurbosfontein and Wupperthal. Numerous fossil plant fragments are also present within Lithofacies 9 (Lithofacies Association G) at all three study sections. Plant fossils are most abundant at the Wupperthal locality, followed by Keurbosfontein and lastly Grootrivierhoogte.

Kingdom METAZOA Subkingdom EUMETAZOA Buetschli, 1910 ? Phylum CNIDARIA Hatschek, 1888 ? Class SCYPHOZOA Götte, 1887 Order CONULARIIDA Miller and Gurley, 1896 Family CONULARIIDAE Walcott, 1886 Genus Conularia

Conularia africana (Sharpe, 1856)

Description: This description is from a partially complete internal mould of Conularia africana (specimen KBKGY 006). Conularia africana is characterised as having a four faced rectangular pyramidal geometry. Each face is equal in shape and surface ornamentation. Owing to its rectangular pyramidal geometry, alternate faces are equal in size. Surface ornamentation consists of numerous transversely oriented parallel projecting ribs (Fig. 4.1 A). These ribs are curved towards the inferred oral end and shallowly depressed along a line in the mesial plane (the “mesial line”) of individual faces. Margins of individual faces are rounded and recessed and run longitudinally across specimens as longitudinal bands. Faces meet along these longitudinal bands. In transverse section, Conularia sp. indet. has a rough rectangular shape with corners rounded off and depressed (Fig. 4.1 B). Depressed regions are synonymous with longitudinal bands and are depressed towards the oral-aboral axis (if constructed as bisectors from corners). These attributes in transverse section give Conularia africana a radial tetrameric symmetry.

Localities: Conularia sp. indet is entirely restricted to Lithofacies Association A of the Gydo Formation and is present at all three study localities.

Superphylum LOPHOTROCHOZOA Halanych, et al., 1995 Phylum BRACHIOPODA Duméril, 1806 Subphylum LINGULIFORMEA Williams et al., 1996 Class LINGULATA Gorjansky and Popov, 1985 Order LINGULIDA Waagen 1885 Family DISCINIDAE Gray, 1840 Genus Orbiculoidea d’Orbigny, 1849

Orbiculoidea baini (Morris and Sharpe, 1846)

Description: This description is from a complete internal mold of Orbiculoidea baini (specimen OW 11/11/12/6.0). Both pedicle and brachial valves of Orbiculoidea baini are impunctate and have a circular outline with concentric ringed surface ornamentation representing alternate growth lines and no plication (Fig. 4.2 A). When viewed anteriorly towards the gape of Orbiculoidea baini is rectimarginate. The brachial valve is dorso-posteriorly convex when viewed laterally and along the commissural plane (Fig. 4.2 B). This gives the brachial valve a conical to triangular shape in lateral profile. The pedicle valve is flat to weakly ventrally concave when viewed laterally and along the commissural plane (Fig 4.2 B). The pedicle foramen is large, oval and depressed it extends along the posterior surface of the palintrope to the posterior margin (Fig. 4.2 C).

Localities: Orbiculoidea baini is entirely restricted to Lithofacies Association A of the Gydo Formation at all three study sections.

Class RHYNCHONELLATA Williams et al. 1996 Order RHYNCHONELLIDA Kuhn1949 Family LEPTOCOELIIDAE Boucot and Gill, 1956 Genus Australocoelia Boucot and Gill, 1956

Australocoelia cf. tourteloti (Boucot and Gill, 1956)

Description: This description is from internal valve moulds of Australocoelia cf. tourteloti. Definitive features of Australocoelia and Australocoelia tourteloti have

88 been outlined by Boucot and Gill (1956) and will be used in this description. Species of this genus are typified by being biconvex (pedicle valve being most convex), angularly plicate (valve folds) with a sub erect pedicle beak. Hinge line is short and curved anteriorly. The gape is typified as being sulcate. Interiorly, brachial valves possess a mound like cardinal process on top of which a bulbous and terminally cleft linear median ridge (the notothyrial mound). The interior of the pedicle valve possesses a deeply impressed pedicle callist with narrow muscle attachment loci. Hinge teeth are triangular in shape, stout with the apex directed anteriorly.

Australocoelia tourteloti is characterised by being sub-circular in geometry and having 5 to 6 angular plications per side with an additional plication in the pedicle valve sulcus and an additional pair on the brachial valve fold. Anterior margins of both valves are crenulated (Boucot and Gill, 1956) (Figure 4.3 A and B).

Localities: Australocoelia cf. tourteloti is present in Lithofacies Association A – C of the Gydo and Gamka Formations at all three study sections.

Order SPIRIFERIDA Waagen, 1883 Family HYSTEROLITIDAE Termier and Termier, 1949 Genus Australospirifer Caster, 1939

Australospirifer sp. indet.

Description: This description is from external brachial valve casts and internal pedicle valve moulds of Australospirifer sp. indet. Definitive features of Australospirifer have been outlined by Reed (1925) and Isaacson (1977) and will be used in this description. Pedicle and brachial valves of Australospirifer sp. indet. are convex, plicate (up to 18 plicae per valve) and have a trapezoid outline with straight margins and rounded cardinal angles. Gape is pauciplicate. Valves have a triangular outline. Externally brachial valves have a strong and smooth mesial sulcus, narrow hinge line and no dorsal inter-area. Internally, pedicle valves possess stout hinge teeth which are triangular in shape with the apex directed anteriorly, large muscle attachment loci and a broad ventral inter-area and a strophic hinge line (Fig.4.4). Australospirifer sp. indet. specimens from Lithofacies 9 of the Boplaas

Formation preserve fine detail preservation of lophophores and supporting brachidium (Fig. 3.85 B). Valves of Australospirifer sp. indet. are also complete and in their life position. Individual specimens are oriented gape up, or resting on their brachial valves, and appear to be epifaunal to semi-infaunal (Fig. 3.85 A).

Localities: Australospirifer sp. indet. is present in Lithofacies Association A – C of the Gydo and Gamka Formations at all three study sections and in Lithofacies 9 from the Boplaas Formation at Grootrivierhoogte.

Class STROPHOMENATA Williams et al., 1996 Order ORTHETIDA Waagen, 1884 Family SCHUCHERTELLIDAE Williams, 1953 Genus Schuchertella Girty, 1904

Schuchertella sp. indet.

Description: This description is entirely from in situ external brachial valve molds tentatively identified as belonging to Schuchertella sp. indet. Valves have a sub- circular outline with smooth margins and are impunctate. Surface ornamentation is costellate with numerous fine ribs which radiate anteriorly from the umbo. Radial growth lines may be visible. Hinge line is slightly astrophic. Detail regarding internal features of brachial valves and external and internal features of pedicle valves is unknown (Fig. 4.5).

Localities: Schuchertella sp. indet. is present in Lithofacies Association A of the Gydo Formation at all three study sections and the Voorstehoek Formation at Grootrivierhoogte.

Order PRODUCTIDA Sarytcheva and Sokolskaya, 1959 Family CHONOSTROPHIDAE Muir-Wood, 1962 Genus Notiochonetes Muir-Wood, 1962

Notiochonetes cf skottsbergi Clarke, 1913

Description: This description is from an internal brachial valve cast and internal pedicle valve mold. Definitive features of Notiochonetes skottsbergi have been outlined by Hiller (1987) p. external features are summarised as: “Medium to large, subelliptical, planconvex to slightly concavoconvex shells; maximum width usually anterior to hinge line; hinge line straight, spines not observed, apparent spine openings seen along hinge line in one specimen; valves two thirds as long as wide; anterior commisure rectimarginate; cardinal angles obtuse; pedicle valve gently convex in lateral profile; anterior profile gently convex but slightly arched medianly; ventral interarea catacline to aspacline; brachial valve flat or very gently concave in lateral profile; dorsal interara short, anacline; chilidial plates disjunct; ornament of rounded costellate separated by narrower V-shaped interspaces; ribs increase by branching and intercalation to 8-9 per 5 mm at 15 mm growth stage.”

Internal features of pedicle and brachial valve interiors’ of Notiochonetes cf. skottsbergi are summarised as follows from Hiller (1987): Pedicle valve interior - Characteristically have large elongate teeth which diverge at 120-130 with no dental plates. A large deeply impressed posteriorly flabellate muscle field occupies ¾ of the pedicle valve inner surface. A thick medium myophragm divides the muscle field and extends to ¾ of the valve length with smaller radiating ridges dividing the muscle field further either side of the medium myophragm. Adductor muscle scars are small and posteriorly situated on either side of the medium myophragm and extending to no more than 1/3 of the valve length. The muscle field is bounded posteriorly by thick ridges. Brachial valve interior – Ventrally projecting cardinal process is tri- or quadralobed. Tooth sockets are deep and bounded by inner and outer ridges (socket ridges). Inner ridges are unfused to the cardinal process and diverge at 110 – 120. Anderidia are well developed and extend from the anterior to the cardinal process to less than 1/3 of the valve length. Anderidia diverge at about 20. The medium septum is posteriorly thickened and extends from in front of the cardinal process to 4/5 the valve length. Posterior adductor muscle scars are triangular and bounded by anderidia and inner tooth socket ridges. Posterior muscle scars in some specimens are bisected by a low and short septum. Anterior muscle scars are smaller and more elongate with respect to posterior muscle scars and are situated between the medium septum and anderidia (Fig. 8.6).

Localities: Notiochonetes cf. skottsbergi is present in Lithofacies Association A of the Gydo Formation at all three study sections.

Phylum MOLLUSCA Linnaeus, 1758 Class BIVALVIA Linnaeus, 1758 Order MODIOMORPHIDA Newell, 1969 Family MODIOMORPHIDAE Miller, 1877 Genus Sphenotomorpha Williams and Berger, 1917

? Sphenotomorpha sp. indet.

Description: The identification of bivalve valves belonging to Sphenotomorpha sp. indet. is dubious as the genus is known only from individual external valve molds. Laterally, valves are elliptical in outline. Umbo and beak is situated in the anterior 1/4 of the valve. Beak apex is directed dorso-anteriorly. Feint and concentric growth lines radiate ventrally towards the valve margin from the umbo (Figs. 3.8 and 4.7).

Localities: ? Sphenotomorpha sp. indet. has been found in Lithofacies Association A of the Gydo Formation and Lithofacies Association C of the Gamka Formation at Grootrivierhoogte.

Order PROTOBRANCHIA Pelseneer, 1889 Family MALLETIIDAE Adams and Adams, 1858 Genus Nuculites Conrad, 1841

Nuculites sp. indet.

Description: This description is from complete internal valve casts of specimen (OW 001) from the Gydo Formation at Keurbosfontein. In lateral aspect, Nuculites sp. indet is sub-ovate in outline. Feint and concentric growth lines radiate ventrally from the dorso-anteriorly situated umbo and beak. The umbo and beak are situated at approximately the anterior 1/3rd of the valve length. Beaks are triangular in outline and projected above the hinge line and occupy approximately 5/8th of the valve length. A prominent and mesial directed umbonal ridge is present. The umbonal

92 ridge is situated towards the anterior surface of the umbo. In lateral aspect, the umbonal ridge extends ventrally from the beak along approximately 3/4 of the valve width and ends at the approximation of the pallial line. In dorsal aspect, Nuculites sp. indet. is biconvex and bilaterally symmetric. Beaks are incurved mesially toward the hinge line. Apex of beak is ventrally directed. Feint impressions of hinge teeth are present on the hinge line. Hinge line is relatively flat and extends the entire length of the valve (Fig. 4.8).

Localities: Nuculites sp. indet. is common in Lithofacies Association A of the Gydo Formation at all three study sections.

Genus Palaeoneilo Hall and Whitfield, 1869

Palaeoneilo sp. indet.

Description: This description is from complete internal valve casts of specimen (KBK ) from the Gydo Formation at Grootrivierhoogte. In lateral aspect, Palaeoneilo sp. indet. is rhomboido-ovate in outline. Feint and concentric growth lines radiate ventrally from the dorso-anteriorly situated umbo and beak. The umbo and beak are situated at the anterior margin of the valve. Beaks are triangular in outline and only slightly projected above the hinge line. Beaks occupy approximately half of the valve length. In dorsal aspect, Palaeoneilo sp. indet. is biconvex and bilaterally symmetric. Beaks are incuved mesially towards the hinge line. Apex of beak is dorso-anteriorly directed. Hinge line is curved and follows the contour of the valve along its entire length (Fig. 4.9).

Localities: Palaeoneilo sp. indet. is common in Lithofacies Association A of the Gydo Formation at all three study sections.

Class GASTROPODA Cuvier, 1797 Order EUOMPHALINA de Koninck, 1881 Family HOLOPEIDAE Wenz, 1938 Genus Holopea Hall, 1847

? Holopea sp. indet.

Description: The identification of conches belonging to Holopea sp. indet. is dubious as the genus is known only from relatively complete internal conch casts. Conches are smooth, turbinate and right laterally coiled about the umbilical axis (Fig. 8.10).

Localities:? Holopea sp. indet. is known only from Lithofacies Association A of the Gydo Formation at all three study sections.

Superphylum ECDYSOZOA Aguinaldo et al., 1997 Phylum ARTHRODOPDA Latreille, 1829 Subphylum TRILOBITOMORPHA Class TRILOBITA Walch, 1771 Order PHACOPIDA Salter, 1864 Family HOMALONOTIDAE Chapman, 1890 Genus Burmeisteria Salter, 1865

Burmeisteria sp. indet.

Description: Burmeisteria sp. indet. is known from specimens OW 11/11/12/4.5, OW11/11/12/3 and uncatalogued fragmentary internal casts and external molds of cephalons, thoraxes and pygidia attributed to the genus Burmeisteria(Salter) in accordance with criteria provided by Cooper (1982). The following description is based upon features manifested in specimens collected for this project.

Cooper (1982) characterised the cephalon of Burmeisteria to be moderately to highly dorsally convex with a subtrigonal to trapezoidal outline. Tubercles may be present on the cephalon with no sense of organisation and vary among individuals. Facial suture is proparian, separating the librigena and fixigena with posterior branches cutting lateral margins of the cephalon in front of the genal angles. Anterior branches are varied and may be transverse, concave or biconcave. Genae are strongly convex, rounded and bear prominent bosses ontop of which elevated, but poorly developed eyes are situated surrounded by shallow circum-ocular furrows. One specimen from the Gydo Formation at Grootrivierhoogte bears small genal

spines (Fig. 4.11 C), thus differing from what has been described by Cooper (1982). Eye bosses are located lateral to the glabella and are located at the lateral termination of 2p furrows with the glabella. The glabella of Burmeisteria is slightly elevated and is sub-trapezoidal in outline. The anterior margin of the glabella is abruptly truncate with rounded corners and possesses a gently convex preglabellar furrow. Glabellar furrows are present in some individuals (reasoned to be more pronounced in juveniles by Cooper (1982)). 3 furrows are present per side. Furrows are mesially directed from the lateral glabella margin. 3p furrows are weak, short and directed posteromedially sub-parallel to the larger and more pronounced 2p furrow. 1p furrows are mesially curved toward the occipital furrow. 1p furrows tend to separate weakly developed sub-trigonal pre-occipital lobes. Lateral border furrows are shallow whilst occipital furrows are prominent. Posterior border furrows are deep and shallow laterally. Cooper (1982) recognises a prominent and gently concave preglabellar field but with no anteriorly situated anterior border and border furrow. Specimens found in this project do appear to have a small anterior border and anterior border furrow (Figs. 4.11 A – C). The internal mold of Burmeisteria herscheli (Murchison) illustrated in Fig. 9 pg. 19 of Cooper (1982) also has an anterior border and accompanying anterior border furrow.

Thoraxes attributed to Burmiesteria are usually anteriorly broad and comprised of 13 segments. Individual segments are in turn comprised of (moving mesially-laterally) an axial ring, axial furrow and pleura. Spine bases are arranged along the inner lateral borders of axial rings. Axial furrows are indistinct but are pitted at the lateral ends of interpleural furrows. Pleural furrows are deep and diverge towards the main trilobite axis (Fig. 4.11 D).

Pygidia are convex, trilobed and triangular in outline. Axis of pygidium may comprise 16-17 granulose ornamented axial rings with posterior most axial rings being indistinct. Pygidial border is rounded and pointed (Fig 4.11 D).

Localities: Burmeisteria sp. indet. is ubiquitous in Lithofacies Association A of the Gydo Formation at all three study sections.

Genus Bainella Rennie, 1930

Bainella africana Salter, 1856

Description: This description is from a relatively complete specimen (specimen ZW 001) and consists of an internal cephalon mold and thorax and pygidium cast. The defining features of Bainella africana (Salter) have been discussed at length by Cooper (1982) and shall be used in this description of available material.

Cephalon: Convex, subtrigonal in outline. Surface is relatively smooth. Glabella is convex and pentagonal in outline. Facial suture is proparian as the posterior branches of the facial suture ending in front of the genal angle and separates the librigena and fixigena. Anterior branch continues anteriorly and terminates against the glabella. Genae are convex and sub-trigonal in outline with well developed eyes lateral to 2p furrow and raised upon palpebral lobes. Genal angles are angular to subrounded with posteriorly directed genal spines which are incurved towards the mesial plane. Glabella is furrowed, possessing 3 furrows per side, and distinctly lobed, 3 lobes per side. 3p and 2p glabellar furrows are weak and indistinct, whilst 1p furrows are well developed and deepen towards the mesial plane. 3p furrows are anterolaterally directed whilst 2p and 1p furrows are transverse. L2 and L3 lobes coalesce at the glabellar margin about the 2p furrow. Anterior border in this specimen is noticeably absent. The occipital lobe is present, and incomplete in this specimen. The occipital lobe usually possesses a long median spine. Occipital furrows are present and are deepest towards the mesial plane and situated behind L1 lobes. Posterior border furrows are deep, broad and shallow towards the mesial plane. Lateral borders are shallow and almost indistinct (Fig. 4.12 A and B).

Thorax: Broad consisting of 11 segments. Individual segments composed of (moving mesially-laterally) axial ring, axial furrow, pleura. Axial rings are dorsally convex and possess a prominent dorsally directed medial spine. Axial furrows are deep and well developed. Pleura are dorso-laterally convex and possess well developed pleural furrows and facets (grooves sensu Cooper (1982)) (Fig. 4.12 C and D).

Pygidium: Triangular in outline with posterior margin directed posteriorly. Pygidium is fused and consists of 7 individual segments. Individual segments consist of (moving mesially-laterally) axial ring, pleural furrow and pleural rib. Axial rings are dorsally convex and lack medial spines. Pleura are dorso -laterally convex. Pleural furrows are deep and well developed. Pleural furrows become deeper and better developed posteriorly. Pleural facets are indistinct and furrows do not extend to the pygidial border. Posterior 4-5 axial rings appear to be fused and raised above the ventral surface of the specimen, making the specimen superficially appear to have a tail (Fig. 4.12 C and D).

Localities: Bainella africana has been found in Lithofacies Associaiton A of the Gydo Formation at the farm Zonderwater and Grootrivierhoogte. It is not as common as Burmeisteria sp. indet.

Superphylum DEUTEROSTOMIA Grobben, 1908 Phylum ECHINODERMATA Klein, 1734 Subphylum CRINOZOA Class CRINOIDEA Miller, 1821

Indeterminate crinoid ossicles and columns

Description: Crinoid ossciles are known from internal and external casts and molds and are representative of scattered elements of disarticulated individuals. In cross section, ossicles are round with articulation facets which radiate around a circular central opening. Geometry of articulation facets are variable and broadly are either synstosial or symplexial.

Localities: Crinoid ossicles are common in Lithofacies Association A – C of the Gydo, Gamka Formations at all three study sections. Crinoid ossicles have also been found in Lithofacies Association C of the Voorstehoek Formation at Keurbosfontein and Wupperthal.

Phylum CHORDATA Haeckel, 1874 Subphylum VERTEBRATA Cuvier, 1812

Indeterminate vertebrate neurocranium

Description: This description is of a tentative vertebrate neurocranial cast (specimen KBKGY 009). Only the otico-occipital unit of the neurocranial cast appears to be present. Sutures between individual bones are poorly visible making identification among bones in the ptico-occipital region difficult. In light of this, three broad regions of the otico-occipital cast may be identified when viewed along the palatal surface. These are the (moving posteriorly-anteriorly): exoccipital, supraoccipital and basioccipital bones (Fig. 4.13 E and F).

In palatal aspect, a prominent anterior-posterior directed suture runs along the lingual plane. This suture unites left and right precursors of the exoccipital, supraoccipital and basioccipital bones, the parachordals and occipitals. Collectively these bones form larger entities, the basal plate (basioccipital) and the occipital arch (exoccipitals and supraoccipitals). Otic capsules are lateral to the basioccipital and have an ellipsoid outline. Exoccipitals are lateral to the supraoccipital and are cone shaped with apexes directed palatally. The articulation facet between the otico- occipital and ethmo-sphenoid units is present at the anterior end of the neurocranial cast and may represent a possible intra-cranial articulation facet (Fig. 4.13 E and F).

In lateral aspect, the neurocranial cast is noticeably cranially convex, with the base of the cast being planar. The basioccipital bone is the most basal bone, followed by the supraoccipital and the exoccipital. A process, the postotic process, is postero- palatally projected from the otic capsule, in the same plane as the basioccipital. The supraoccitpital is rounded and convex anteriorly. The exocciptals have prominent processes which project palatally as exoccipital condyles, one per side. Exoccipital condyles have a triangular outline with the apex palatally directed. The anterior margin of the exoccipital is noticeably posteriorly curved. A prominent palatal- cranial directed furrow is present between the basal plate and occipital arch this may represent the hyomandibular articulation facet. The articulation facet between the otico-occipital and ethmo-sphenoid units is curved and is convex posteriorly in the same plane as the postotic process and conxex anteriorly above it (Fig 4.13 C and D).

In cranial aspect, the neurocranial cast is convex lingually. The palatal-cranial directed furrow between the basal plate and the occipital arch is present towards the posterior third of the cast. The otico-occipital and ethmo-sphenoid articulation facet is concave anteriorly (intracranial surure) (Fig. 4.13 B).

In occipital and palato-occipital aspect, the anterior margin of the exoccipital is noticeably curved and rhomboidal in outline and may be representative of the atlantico-occipital facet. This facet extends from the inferred exocipital-supraoccipital junction to the apex of the exocipital condyles. Exoccipital condyles are cone shaped with the apex directed palatally. The foramen magnum is situated between the exoccipitals and is triangular in outline with the apex oriented palatally. The palatal-cranial directed furrow between the basal plate and occipital arch show a well defined delineation between these two areas. The occipital arch is notably square in outline, with the exoccipital condyles projecting palatally from palatal angles. The palatal surface of the basioccipital is cranially convex along the lingual plane. The cast overall is both cranially and palatally convex (Fig 4.13 A and E).

Localities: Neurocranial cast (KBKGY 009) represents a single specimen from Lithofacies Association A of the Gydo Formation at Grootrivierhoogte.

PROBLEMATICA

Superphylum Unknown Phylum Unknown Class TENTACULITA Order TENTACULITIDA Ljaschenko, 1955 Family TENTACULITIDAE Walcott, 1886 Genus Tentaculites von Schlotheim, 1820

Tentaculites sp. indet.

Description: This description is from complete internal conch casts and internal conch molds. Conches are conical in shape. Here conches are inferred to taper to a

point (the apex) adapically and are open (the aperture) abapically. Numerous transversely oriented and parallel ridges or folds make up the surface ornamentation of Tentaculites sp. indet (Fig. 4.14).

Localities: Tentaculites sp. indet. is common in Lithofacies Association A – C of the Gydo, Gamka. Tentaculites sp. indet. conches have also been found in Lithofacies Association C of the Voorstehoek Formation at Keurbosfontein and Wupperthal.

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Chapter 5 Discussion

5.1 Introduction to Marginal and Shallow Marine Environments

Marginal marine environments are located at the boundary between continental and marine realms (the coast) and extend until the shoreline (Reading and Collinson, 1996, Boggs, 2006, Huggett, 2007b, Nichols, 2009). The coast exists at the interface of fluvial, tidal, wave and current processes (Reading and Collinson, 1996, Bhattacharya and Giosan, 2003, Boggs, 2006, Nichols, 2009). The shoreline is defined as the physical margin between continental and marine realms (Reading and Collinson, 1996, Nichols, 2009). Here sediment may be sourced and reworked from and by any process, thus strongly influencing expected marginal marine environments and their morphology (Reading and Collinson, 1996, Bhattacharya and Giosan, 2003, Boggs, 2006, Nichols, 2009). Sediment supply to coasts is strongly from an original fluvio-deltaic source draining from the interior of continents, but may also be sourced from more distal shallow marine shorefaces, or by longshore drift along the shoreline (Reading and Collinson, 1996, Bhattacharya and Giosan, 2003, Boggs, 2006, Nichols, 2009). A continuum therefore exists within and among marginal marine environments and includes cheniers, deltas, lagoons, estuaries, beaches, strandplains, barrier islands, tidal flats and inlets (Reading and Collinson, 1996, Boggs, 2006, Nichols, 2009) (Fig. 5.1).

Shallow marine environments extend basinward from marginal marine environments and thus lie beyond the shoreline and extend from below the inter-tidal zone to the shoreface, shelf and slope environments (Johnson and Baldwin, 1996, Hampson and Storms, 2003, Boggs, 2006, Nichols, 2009). Shallow marine environments may either be located pericontinentally (i.e. on continental margins along continental margins in passive margin settings), or epicontinentally (within continents as broad partially enclosed shallow epeiric seaways) (Johnson and Baldwin, 1996, Boggs, 2006, Nichols, 2009). These environments characteristically are < 200 m deep with gentle basinward directed depositional gradients (~ 1 – 0.1°) (Johnson and Baldwin, 1996). Sediment within shallow marine environments is substantially derived from proximal marginal marine source areas and usually is relatively mature (Nichols,

101

2009). In shallow marine environments, oceanic processes such as tidal, wave and storm processes as well as oceanic currents strongly affect the morphology and characteristics of shallow marine environments (Johnson and Baldwin, 1996, Boggs, 2006, Nichols, 2009). As in marginal marine environments, a continuum too exists among these processes (Fig. 5.2).

The physiographic expression of shallow marine environments and associated large scale sub-environments are the foreshore, shoreface and shelf (Johnson and Baldwin, 1996, Reading and Collinson, 1996, Hampson and Storms, 2003, Boggs, 2006, Nichols, 2009). These environments and their characteristic attributes have been summarised in Figure 5.3. The foreshore corresponds with the intertidal zone and lies between mean high and low tide and is within the swash zone and dips seaward at 2 - 3° (Johnson and Baldwin, 1996, Reading and Collinson, 1996, Hampson and Storms, 2003, Boggs, 2006, Nichols, 2009). The shoreface, as is traditionally envisioned, exists in the sub-tidal zone between mean low tide and mean fair-weather base (Reading and Collinson, 1996, Hampson and Storms, 2003, Boggs, 2006, Nichols, 2009, Peters and Loss, 2012). Hampson and Storms (2003) have discussed at length the different criteria workers have used to define the base of the shoreface being at, or near, mean storm wave base as well as other subdivisions of the shoreface (Fig.5.4).

The lithofacies recognised within the Ceres Subgroup, strongly suggest that sedimentation within each upward coarsening megacycle of the Subgroup spans a continuum of shallow to marginal marine environments. Broadly two palaeo - environments have been recognised within the Ceres Subgroup based on comparisons with various lithofacies succession models from modern and ancient environments. These are a wave dominated and storm influenced shoreline (Gydo – Gamka System and Voorstehoek – Hex River System) and a wave and tide influenced estuary (Tra-Tra – Boplaas System). The lithofacies associations and defining attributes of each recognised palaeo - environment within the Ceres Subgroup shall be discussed.

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5.2 Wave and Storm Dominated Shoreline

The lithofacies successions of ancient wave and storm dominated shorelines are well recognised and have been studied by various workers (Leckie and Walker, 1982, Smith and Tavner – Smith, 1988, Kalkreuth and Leckie, 1989, Leckie and Krystinik, 1989, Van Wagoner, et al., 1990, Brenchley, et al., 1993, Kamola and Van Wagoner, 1995, Reading and Collinson, 1996, Johnson and Baldwin, 1996, George, 2000, Bhattacharya and Giosan, 2003, Hill, et al., 2003, Hampson and Storms, 2003, Charvin, et al., 2010, Hampson, 2010, Allen and Johnson, 2011, Green and Smith, 2012, Vakarelov, et al., 2012).

Wave and storm dominated shorelines of the Bokkeveld Group interpreted in this project are coarsening upward depositional systems consisting of a distal and lower wave and storm dominated shoreface environment and a proximal and upper beach environment. Five distinct sub-environments are recognisable in idealised wave and storm dominated shorefaces. These sub – environments are (moving from distal to proximal): Middle to Outer Shelf/Offshore (Os), Inner Shelf/ Offshore Transition Zone (OT), distal Lower Shoreface (dLSF), proximal Lower Shoreface (pLSF) and Upper Shoreface (USF). Beach environments are composed of (moving from distal to proximal) Foreshore (Fs) and Backshore (Bs) sub – environments. The intimate hydrodynamic processes which typify these environments and related sub-environments have been used by workers for the basis of lithofacies association sub-division. These have been used in this study and are summarised in Table 5.1.

Slug models for idealised wave and storm dominated shorelines for the Gydo – Gamka and Voorstehoek – Hex River Systems have been illustrated in Figures 5.5 and 5.6 respectively. These figures show the envisioned vertical and lateral distribution among the various sub - environments at each locality. Each system is regressive.

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5.2.1 Lithofacies Association A (Middle to Outer Shelf/Offshore)

Representative lithofacies of L.A. A indicate sedimentation of clay and silt sized grains by slow suspension deposition in a low energy environment (Lithofacies 1 and 2.1.) with occasional lower flow regime (Fr<1) transport of sediment as ripples (Lithofacies 2.2.). Such hydrodynamic conditions are typical in sub-tidal Os pelagic environments (Potter, et al., 1980, Johnson and Baldwin, 1996, Stow, et al., 1996). In an idealised upward shoaling succession, Lithofacies 2 may possibly represent a proximal equivalent of Lithofacies 1. The presence of surface trail ichnofossils associated with Nereites and Cruziana ichnofacies have been noted to occur within offshore environments (Table 5) (Miller, 2007, Seilacher, 2007). Only Nereites sp. and another unknown surface trail ichnofossil were found in L.A. A of the Gydo and Voorstehoek Formations (Fig. 3.14).

Intraclastic and bioclastic sandstone filled scours and troughs in Lithofacies 2.2 (Fig. 3.13) may represent exceptionally large and short lived storm activity in the basin. Storm wave orbital’s may have extended below mean storm wave base and have scoured into Lithofacies 2.2 and deposited clastic and bioclastic material from the shoreface in scours as return storm surge ebb deposits (Duke, 1990). Return flow storm surge ebb deposits generally are transported as density currents. Density current transport of clastic and bioclastic material may also explain the lack of apparent damage of bioclasts in bioclast supported coquinites in Lithofacies 1 (Fig. 3.9).

In the Gydo Formation, L.A. A is characterised by both Lithofacies 1 and 2, whereas only Lithofacies 2 is expressed in the Voorstehoek Formation. L.A. A overall is also thicker in the Gydo Formation than it is in the Voorstehoek Formation at all three study sections. It is possible that L.A. A of the Voorstehoek Formation represents a more proximal expression of the Os sub - environment, whilst L.A. A of the Gydo Formation represents a complete and more distal expression of the Os sub – environment.

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5.2.2 Lithofacies Association B (Lower Offshore Transition Zone to Upper distal Lower Shoreface)

The heterolithic sandstone and mudstone lithologies of Lithofacies 3 are indicative of alternate traction (sandstone) and suspension (mudstone) in response to alternate high and low energy hydrodynamic conditions. In idealised storm influenced shorefaces, this has been attributed to deposition by oscillatory and unidirectional currents during major storms near storm wave base (Kalkreuth and Leckie, 1988, Reading and Collinson, 1996, Charvin, et al., 2010, Vakerelov, et al., 2012). Here sandstones are deposited by storm related currents, and mudstones by suspension deposition with waning storm energy and background fair-weather conditions. The presence of symmetric sediment starved sandstone flaser ripples and reworked small scale HCS silty sandstones exemplify the oscillatory storm generated genesis of Lithofacies 3 (Figs. 3.17 and 3.18). Kalkreuth and Leckie (1988) and Vakerelov, et al. (2012) have noticed identical sedimentary structures to have formed at the same inferred palaeobathymery in sedimentary deposits of the Lower Cretaceous Western Interior Seaway and Upper Cretaceous Bearpaw and Horseshoe Canyon Formation interval respectively. All of these characteristics are typical for Offshore Transition Zone sub-environments.

In idealised upward shoaling successions, the dLSF overlies the OT sub - environment (Kalkreuth and Leckie, 1988, Reading and Collinson, 1996, Dumas and Arnott, 2006, Charvin, et al., 2010, Vakerelov, et al., 2012). The dLSF is characterised as possessing sharp based and non - amalgamated HCS silt to fine grained sandstone lensoids which intercalate among heterolithic sandstones and mudstones (Kalkreuth and Leckie, 1988, Van Wagoner, et al., 1990, Cheel and Leckie, 1993, Kamola and Van Wagoner, 1995, Hampson and Storms, 2003, Charvin, et al., 2010).

Here, HCS sandstones (in this study, Lithofacies 4) represent storm “scour and drape” deposits from periodic storm related oscillatory, unidirectional or combined flow currents at storm wave base during storm conditions (Kalkreuth and Leckie, 1988, Van Wagoner, et al., 1990, Cheel and Leckie,

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1993, Kamola and Van Wagoner, 1995, Hampson and Storms, 2003, Dumas and Arnott, 2006, Charvin, et al., 2010, Quin, 2011). Intercalated heterolithic sandstones and mudstones (in this study, Lithofacies 3) represent suspension deposition with waning storm flow energy and calmer background fair – weather wave base conditions (Kalkreuth and Leckie, 1988, Myrow and Southard, 1991, Cheel and Leckie, 1993, Hampson and Storms, 2003, Charvin, et al., 2010). The high degree of bioturbation of Lithofacies 3 (in places) may attest to calmer basinal conditions before and after periodic storm events (Fig. 3.20 B).

The taphonomy of bioclastic coquinites and fossils may attest to deposition in turbulent conditions at time of deposition in this sub - environment. Coquinites and fossiliferous horizons are always mixed assemblage, disarticulated from their life positions and with a high degree of damage (Figs. 3.21, 3.26 and 3.27). Fossils (when found in Lithofacies 4) are always present on the lower bedding planes with mudstone intraclasts as storm scour lags (Fig. 3.24). Plant fossils found in Lithofacies Association B are also disarticulated and damaged. These fossils are inferred to have been transported to the dLSF/OT sub – environments from terrestrial environments, possibly in close proximity to the palaeo - foreshore/backshore; as such these fossils are allochthonous to the dLSF/OT. Transport may have been assisted by storm backwash, or as detritus blown into the basin.

Since the hydrodynamic processes operating at both sub-environments is so similar, both have been included in this Lithofacies Association.

L.F. B is well developed in the Gydo, Gamka and Voorstehoek Formations.

5.2.3 Lithofacies Association C (proximal Lower Shoreface)

Amalgamated HCS and SCS sandstones have been proposed as indicative of pLSF sub - environments (Kalkreuth and Leckie, 1988, Van Wagoner, et al., 1990, Cheel and Leckie, 1993, Kamola and Van Wagoner, 1995, Hampson and Storms, 2003, Dumas and Arnott, 2006, Charvin, et al., 2010). In

106 idealised upward shoaling successions, amalgamated HCS and SCS sandstones overlie non – amalgamated HCS beds (Kalkreuth and Leckie, 1988, Van Wagoner, et al., 1990, Cheel and Leckie, 1993, Kamola and Van Wagoner, 1995, Hampson and Storms, 2003, Charvin, et al., 2010). Amalgamation of HCS and SCS beds has been explained by erosion and deposition from storm related oscillatory, unidirectional or combined flow currents at, or above storm wave base during storm conditions, but below mean fair-weather base (Kalkreuth and Leckie, 1988, Cheel and Leckie, 1993, Charvin, et al., 2010). Here, aggradation rates are high and storm beds are not reworked by fair-weather wave processes, thus encouraging amalgamation (Dumas and Arnott, 2006, Charvin, et al., 2010). Based on experimental data, Dumas and Arnott (2006) have illustrated a predictable occurrence of isotropic and anisotropic HCS and SCS bedforms within an idealised progradational shoreface under storm conditions (Fig. 5.7). Neither of which could be distinguished with confidence from field observations. Ichnofossils attributed to Skolithos and Cruziana ichnofacies (Table 5.1) have been found in association with HCS and SCS beds. No ichnofossils whatsoever were found in association with L.A. C, this may be due to high aggradation rates. Here, HCS and SCS deposits may be reworked, thus destroying original ichnofabrics and rendering the rocks as “barren” (Cheel and Leckie, 1993).

As in L.A. B, fossils present in L.A. C are always disarticulated from their life positions, are highly damaged and form a mixed assemblage. Fossils are nearly always present on the lower bedding planes of beds as storm scour lags.

As previously mentioned, coarse grained ripple sandstones (Lithofacies 5) have been demonstrated to form in similar hydrodynamic conditions as those expected for HCS and SCS and thus are expected to form in combined flow conditions (Leckie, 1988, Cummings, et al., 2009). Leckie (1988) has described coarse grained ripples occurring in OT, dLSF, pLSF and USF sub – environments (Fig. 5.8 and Table 5.1). Coarse grained ripples are most common in USF sub – environments, rare in pLSF sub – environments and

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very rare in OT sub – environments, or in association with transgressive surfaces (Leckie, 1988). Coarse grained ripples, therefore may overlie or intercalate with finer grained HCS an SCS beds (USF, pLSF and dLSF sub – environments), and/or with heterolithic sandstone beds (dLSF and OT sub – environments) or intercalated with shoreface associated bars (USF). The placement of Lithofacies 5 within L.A. C in this study is thus tentative and based entirely on similar hydrodynamic conditions to that expected for the formation of HCS and SCS.

In the Gamka Formation, the presence of Lithofacies 5 overlying Lithofacies 4 and in turn intercalating with the overlying Lithofacies 7 (L.A. D) strongly matches hypothesised distributions of coarse grained ripple sandstone beds in an idealised upward coarsening progradational sequence presented by Leckie (1988) (Fig. 5.8). Distributions of coarse grained ripple sandstones by Leckie (1988) are based on observations from numerous modern and ancient localities worldwide. Coarse grained ripple sandstones overlying amalgamated HCS and SCS sandstones have also been recognised in the Chungo Member (Waipabi Formation) of Alberta, Canada by Cheel and Leckie (1992).

The three and possibly four (including “whale back” structures resembling hummock and swale topographic features) recognised variants of Lithofacies 5 within the study area may represent subtle changes in erosion and deposition (aggradation) rates, energy conditions and sediment supply within the environment of deposition, or a complete change in environment of deposition. Given interpretation of variants is thus tentative and require further research.

The amalgamated variant of Lithofacies 5 (Fig. 3.41 A and B) bears a strong resemblance to the anisotropic hummocky cross stratified sandstone lithofacies described by Basilici et al. (2012) from the Cambrian – Ordovician Lagarto and Palmares Formations from north-eastern Brazil. Basilici et al. (2012) did not recognise their lithofacies to be coarse grained ripples, nor did they recognise the lithofacies to be amalgamated. Their interpretation is

108 based upon the asymmetric form of the ripples which bear a resemblance to anisotropic hummocky cross stratification and to have deposited in response to unidirectional dominant combined flow storm activity. Leckie (1988) has noted that coarse grained ripples are slightly asymmetric. It is possible that coarse grained ripples may be mistaken for anisotropic hummocky cross stratified beds due to similar geometry and hydrodynamic conditions for formation. Amalgamated coarse grained ripple beds have been described by Cheel and Leckie (1992) and too bear a strong resemblance to the amalgamated variant of Lithofacies 5 in this study and the anisotropic hummocky cross stratified sandstone lithofacies of Basilici et al. (2012). Cheel and Leckie (1992) have attributed amalgamation to high aggradation rates associated with aggradation due to lateral erosion and coalescence of individual beds.

Heterolithic mudstone and sandstone partings between individual coarse grained ripple beds (Fig.3.42) may be indicative of waning energy and suspension deposition after storm events.

It is possible that the superficial “whale back” hummock and swale type topographic features (Fig. 3.35, 3.37 – 3.39) in Lithofacies 5 of the Gamka Formation at Grootrivierhoogte may be HCS and SCS beds in coarse grained sediment under combined flow. HCS and SCS beds in coarse grained sandstones are rare and require larger energy to form (Cheel and Leckie, 1993, Dumas and Arnott, 2006). Evidence supporting this claim is met by the presence of large sandstone intrabreccias, often angled into the bedding plane, en masse collection of mudstone intraclasts in topographic lows as well as prod marks made by fragmentary plant material, rill marks and asymmetric current ripples on tops of bedding planes.

A two part time-wise sequence of events is envisioned and is based on sedimentary features generated first by an onshore directed oscillatory flow superimposed by a later related offshore directed unidirectional flow event and is summarised as follows: Onshore directed oscillatory flow features and deposits: It is suggested that an exceptionally large oscillation wave orbital

109 scoured a pre-existing surface, which was otherwise dominated by coarse grained ripple deposits, and deposited low angle coarse grained HCS and SCS beds. The velocity of this oscillation wave orbital was great enough to rip up and traction transport sandstone intrabreccias and mudstone intraclasts from the basin as it moved proximal (shoreward). With waning flow, sandstone intrabreccias and mudstone intraclasts deposited out from the flow. Offshore direct unidirectional flow features and deposits: Turbulent basinward directed backwash (unidirectional flow) entrained plant material from a more proximal source area to the deposit thus generating prod marks on bedding plane surfaces and sandstone concretions. Rill marks, indicative of emergence, may represent erosion of topographic high features with concomitant local fall of water depth with backwash. Backwash also generated asymmetric current ripples which prograded across the bedding plane of coarse grained HCS and SCS beds. Mudstone intraclasts were re- entrained within this flow and re-deposited into topographic lows (swales) where they collected. Skolithos sp. ichnofossils are well documented in HCS and SCS sandstones (Cheel and Leckie, 1993, Miller, 2007) and are indicative of environments characterised by relatively high levels of current or wave activity in clean well sorted loose substrates.

From field data, Lithofacies 5 in the Gamka Formation may therefore represent a “middle Shoreface” sub – environment and transitional zone between hydrodynamic and sediment sorting conditions between USF and pLSF sub – environments. This, however, remains to be proven.

Within the study area, Lithofacies Association C is thicker and better developed in the Gamka Formation than it is in the Hex River Formation at all three sections. This may indicate that Lithofacies Association C of the Gamka Formation is a more distal pLSF environment and Lithofacies Association C of the Hex River Formation, a more proximal pLSF environment.

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5.2.4 Lithofacies Association D (Upper Shoreface –Beach Complex)

Macroform cross bedded (Lithofacies 7) and planar parallel laminated (Lithofacies 8) medium to coarse grained sandstones have been associated with USF and beach related sub – environments in the Gamka and Hex River Formations (Reineck and Singh, 1975, Allen, 1980, Van Wagoner, et al., 1990, Reading and Collinson, 1996, Johnson and Baldwin, 1996, Hampson and Storms, 2003, Charvin, et al., 2010, Allen and Johnson, 2011, Vakerelov, et al., 2012). Lithofacies 7 and 8 demonstrably intercalate with one another (Fig. 3.56) in the study area making discrete recognition of either sub – environment tentative. As such, both Lithofacies have been included within this Lithofacies Association representative of a broad palaeo – environmental interpretation.

In USF sub – environments, tabular and trough cross – bedded bars and dunes are associated with shoreface progradation (Reineck and Singh, 1975, Leckie and Walker, 1982, Reading and Collinson, 1996, Johnson and Baldwin, 1996, Charvin, et al., 2010, Allen and Johnson, 2011). Bars and dunes may prograde normal to the shoreline, or as is most common, in the prevailing longshore drift direction (Reineck and Singh, 1975, Leckie and Walker, 1982, Reading and Collinson, 1996, Johnson and Baldwin, 1996, Charvin, et al., 2010, Allen and Johnson, 2011). Tabular and trough cross – bedded bars and dunes generally are subtidal features which form above mean fair weather wave base and are affected by fair-weather processes such as shoaling and oscillatory wave action as well as tidal processes (as is the case for sandwaves) (Allen, 1980, Reineck and Singh, 1975, Leckie and Walker, 1982, Reading and Collinson, 1996, Johnson and Baldwin, 1996, Charvin, et al., 2010, Hampson and Storms, 2003, Allen and Johnson, 2011, Vakerelov, et al., 2012). Planar parallel laminated beds in USF environments are rare and are thought to be confined to deeper parts of the USF sub – environment (Reineck and Singh, 1975).

Beaches, in part, are typified by medium to coarse grained planar parallel laminated sandstones. In Fs sub – environments, planar parallel laminated

111 sandstone beds dip seawards at a gradient of 2 - 3°being steepest at the beach face (Reineck and Singh, 1975, Reading and Collinson, 1996, Hampson and Storms, 2003). Fs sub – environments are within the intertidal zone and thus are partially subaerially and subaqueously exposed. As a result of this, Fs sub –environments are affected by breaker, surf, swash, backwash and tidal processes (Reineck and Singh, 1975, Reading and Collinson, 1996, Hampson and Storms, 2003). Bs environments are supratidal and gently dip landward behind the berm (Reineck and Singh, 1975, Reading and Collinson, 1996, Hampson and Storms, 2003). The berm marks the upper boundary of the Fs and mean high water mark. The Bs sub – environment is subject to wind activity reworking and depositing sediment (Reineck and Singh, 1975, Reading and Collinson, 1996, Hampson and Storms, 2003). These sub – environments rarely are affected by onshore directed flooding by exceptionally high tides, strong waves and currents and are deposited as washover fans (Reineck and Singh, 1975, Reading and Collinson, 1996, Hampson and Storms, 2003, Allen and Johnson, 2011). Thin ripple lamination may also form in backshore environments in response to winnowing of sediment due to wind action, or decaying flow with flooding events (Reineck and Singh, 1975). Planar parallel lamination in beaches may thus be developed by various processes making their genesis uncertain. Rootlet ichnofossils (Fig. 3.76) and fragmentary plant compression fossils (Fig. 3.74) within Lithofacies 8 does tentatively suggest a subaerial (at times) palaeo-environment akin to that of a beach. If rootlet ichnofossils within Lithofacies 8 are indeed those of plants, it is suggested that these plants grew in Bs sub-environments owing to calmer expected conditions to those of the Fs sub – environment.

In the Boplaas Formation at Grootrivierhoogte, this palaeo – environment possibly represents a strandplain and lateral equivalent of an inner estuary bayhead delta and should correctly be included with L.F. H, if demonstrable.

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5.3 Wave and Tide Influenced Estuary

Estuarine systems have been defined by Dalrymple et al. (1992), p.1132 as: “the seaward portion of a drowned valley system which received sediment from both fluvial and marine sources of which contains facies influenced by tide, wave and fluvial processes. The estuary is considered to extend from the landward limit of tidal facies at its head to the seaward limit of its coastal facies at its mouth.” Dalrymple, et al. (1992) have noted a broad tripartite sub – division of estuaries based on respective regions that are dominated by marine, fluvial or mixed processes. These are: An outer marine process dominated sub – environment (here net sediment transport is landward), a low energy central sub - environment (here sediment transport is convergent) and an inner river dominated, but marine influenced, sub - environment (here net sediment transport is basinward). This has been summarised in Figure 5.9. Estuaries fall into one of two broad types of estuary, depending on the type and amount of energy received at the outer distributary, these are a wave influenced and tide influenced estuary (Fig. 5.10).

Due to the unique distribution of processes which operate in estuarine systems, numerous intertidal, subtidal and intracoastal sub – environments may be present within an individual system making estuarine sequences complex (Reinson, 1997). Depositional environments thus may encompass tidal deltas, inlets, shoals, back- barrier beaches, spits, washover fans, swash bars, point bars, tidal flats, marshes, swamps, stream banks, lagoons, bays and fluvial sub – environments (Reinson, 1997). This palaeo-environment is restricted to the Tra-Tra – Boplaas System and records an overall inferred transgressive sequence.

5.3.1 Lithofacies Association E (Channelised Tidal Flat)

L.A. E has been interpreted to represent a channelised tidal flat (Lithofacies 3 and 11). Mudstone – draped foresets are present on the lee slopes of tabular cross – bedded sandstone bars (Lithofacies 11) (Fig. 3.96) and heterolithic sandstones and mudstones (Lithofacies 3) are flaser bedded (3.16 B). The heterolithic nature of Lithofacies 3 may be attributed to sedimentation in

113 response to rhythmitic energy changes in the depositional environment such as tidal or seasonal changes (Reineck and Singh, 1975, Bridge and Demicco, 2008). Mudstone drapes in both Lithofacies have been recognised to occur in tidal environments and are deposited as suspension depositional features in response to slack water (high tide) conditions (Reineck and Singh, 1975, Bridge and Demicco, 2008). Bioturbation is common in tidal flats which may entirely destroy stratification (Bridge and Demicco, 2008). Lithofacies 11 in L.A. E is interpreted to be tidal channels with their associated channel fill which have erosively cut into a tidal flat (Lithofacies 3).

This lithofacies association is heavily bioturbated with bioturbation becoming more frequent up the succession into L.A. F (Fig. 3.20 A). Diplocraterion sp. ichnofossils have only been found in Lithofacies 11 at Grootrivierhoogte (Fig.3.97). Diplocraterion sp. is a Skolithos ichnofacies ichnofossil and is associated with a wide array of environments characterised by relatively high levels of current or wave activity in clean well sorted loose substrates (Miller, 2007, Seilacher, 2007). Skolithos ichnofacies ichnofossils have been reported from tidal environments, often within tidal channels and inlets (Miller, 2007, Seilacher, 2007). This lithofacies association and interpreted palaeo – environment is similar to Unit 2 of the Albian Stage Paddy Member (Peace River Formation) of Leckie and Singh (1991).

According to the tripartite sub - division of estuaries by Dalrymple et al. (1992), this palaeo – environment is indicative of the outer estuary sub - environment since sedimentation is entirely due to tidal marine processes.

5.3.2 Lithofacies Association F (Lagoon to Washover Flat)

Owing to the en masse bioturbated nature of this entire Lithofacies Association (Lithofacies 3 and 9), detail of sedimentary structures which would otherwise aide in determining a likely palaeo-environment of deposition is absent making this palaeo – environmental interpretation tentative (Fig. 3.82).

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The heterolithic nature of the lower Lithofacies 3 is interpreted to represent sedimentation in a calm bay or lagoonal environment. Lagoons are generally calm water depositional environments dominated by suspension deposition of mud with periodic coarser grained material (Bridge and Demicco, 2008). The general calm water conditions expected in lagoons (Bridge and Demicco, 2008) may explain the highly bioturbated nature of Lithofacies 3 within this Lithofacies Association.

Lithofacies 9 shows an array of fossils associated with terrestrial (coalified plant fossil fragments and plant rootlet ichnofossils) and marine environments (in situ brachiopod molds, Skolithos sp. ichnofossils and meandering surface casts) (Figs.3.84 – 3.87). These fossils are variably associated with one another in Lithofacies 9. Visible sedimentary structures are restricted to vague planar parallel laminations with bedding planes being amorphous and lumpy and sharp partings. The sharp bedding plane contacts between individual beds are assumed to represent alternate pulses of deposition. The sedimentology of Lithofacies 9 indicates that it is a medium to coarse grained sandstone, but with significant mud and white mica as accessory minerals. From the poorly sorted nature of Lithofacies 9 it is assumed that individual beds were originally normal graded events which were later bioturbated.

The features seen in L.A. F within the study area closely allude to those expected for washover flats. Washover deposits commonly extend into lagoon environments when large wind and storm surge events cut through barriers and spill into lagoons, on the landward sides of aeolian dunes, beach ridges or tidal flats (Reinson, 1997, Bridge and Demicco, 2008). Washover fan deposits are generally a few centimetres to up to two metres thick per overwash event and form large sheet-like tabular bodies which extend over a few hundred metres in width oriented normal to the shoreline (Reinson, 1997, Bridge and Demicco, 2008). Washover fans fans may coalesce and be kilometres in width and create extensive washover flats (Reinson, 1997, Bridge and Demicco, 2008). Lithofacies 9 is easily traced throughout the entire study outcrop area as a prominent sandstone unit at the base of the Boplaas Formation with little pronounced change in thickness (Appendix 1 -

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4). Washover flats play a crucial role in creating environments encouraging marsh growth and barrier stabilisation behind lagoons. Such washover flats are intensively bioturbated by fauna and plant roots (Reinson, 1997, Bridge and Demicco, 2008).

With this information, it is suggested that marsh stabilisation by plants is seen as a viable candidate to explain plant rootlet ichnofossils and coalified plant fossils in Lithofacies 9. Episodic spill over events and mixed salinity waters in lagoons too may explain the presence of marine fauna and related ichnofossils. According to Reinson (1997), washover fans deposits are reworked upper shoreface sediment and may be explained by shoreface retreat with transgression. In the shoreface retreat model, it is thought that sediment is eroded from the USF and transported to the LSF to Os as storm beds, or to the lagoonal estuary as washover fan deposits when transgressed. Over time, the USF will migrate landward with transgression and erode barrier, washover flats and lagoonal facies leaving a planar ravinement surface (Fig. 5.11). It is possible that the USF – Beach Complex of the Hex River Formation, in part or wholly supplied sediment for the washover flat environment as it was being transgressed and may explain the absence of a “beach barrier” environment within the Tra-Tra Formation which would have been situated seaward of the lagoon.

According to the tripartite sub - division of estuaries by Dalrymple et al. (1992), this may be indicative of the central estuary sub - environment since sedimentation is within a low energy environment (lagoon) partially derived from a marine component (washover flat).

5.3.3 Lithofacies Association G (Estuary Bay to Subaqueous Bayhead Estuary)

Lowermost L.A. G is heterolithic sandstone and mudstone (Lithofacies 3) which is interpreted to represent sedimentation in calm water environments as the same criteria is met for Lithofacies 3 in L.A. F, but has been interpreted to

116 represent either a lagoon, or as most likely, an interdistributary bay environment owing to the presence of intercalating Lithofacies 6.

Towards the top of the Lithofacies Association numerous single story lenticular wackes (Lithofacies 6) are intercalated within Lithofacies 3 (Figs. 3.48 – 3.52). The lenticular geometry of Lithofacies 6 has been interpreted to represent channels. Lower most channels are single story which becomes more amalgamated up the succession as multi - story channel complexes (Figs. 3.52 and 3.53). Top bedding planes of Lithofacies 6 may either be flat, asymmetric climbing ripple cross laminated or symmetric ripple laminated (Fig. 3.48 – 3.51). Symmetric ripple lamination is usually indicative of wave activity and deposition from oscillatory flows (Tanner, 1967, Reineck and Singh, 1975). Although lagoons and estuaries are protected from marine environments, occasional breaking wave activity may influence lagoons and estuaries behind barriers, or be generated within lagoons (Dalrymple, et al., 1992, Reinson, 1997). Asymmetric climbing ripple cross lamination may indicate decaying flow velocity and current migration (Reineck and Singh, 1975). Lateral to Lithofacies 6, Lithofacies 3 may be asymmetric ripple drift laminated (Fig. 3.19) and have been interpreted to be crevasse splays (Reineck and Singh, 1975). The presence of symmetric ripples suggests deposition within a subaqueous environment; as such this component of L.A. G has been interpreted to represent estuary distributary channels which have prograded into a lagoon environment. This lithofacies association is similar to those expected for delta front environments

5.3.4 Lithofacies Association H (Subaerial Bayhead Estuary)

L.A. H has been interpreted to represent an inner estuarine environment dominated by fluvial sedimentation considered typical for these environments (Dalrymple, et al., 1992, Reinson, 1997). This lithofacies association consists of two lithofacies. Lithofacies 10 has been interpreted to be lateral accretion bars which form on the point channel margin of channels (Fig. 5.12) (Allen, 1970, Miall, 1985, Fielding, et al., 1993, Collinson, 1996, Hugget, 2007). Lithofacies 11 has been interpreted to be downstream accretion bars and the

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sedimentary fill of channels as the downstream migration of subaqueous dunes (Collinson and Thompson, 1989c, Collinson, 1996, Miall, 1985). Rahmani (1989) has noted a similar association of lithofacies within an interpreted middle to upper estaruine environment (Facies E and EE respectively) of the Upper Cretaceous transition between Bearpaw and Horseshoe Canyon Formations and Unit 1 C and D of the Albian Stage Paddy Member (Peace River Formation) of Leckie and Singh (1991) and has been noted in Dalrymple, et al. (1992) and Reinson (1997).

According to the tripartite sub - division of estuaries by Dalrymple et al. (1992), L.A. G and H are indicative of the inner estuary sub - environment since sedimentation is entirely due to fluvial processes with minor marine influence. Figure 5.13 is an idealised diorama of the palaeo – environment of the Tra –Tra – Boplaas System, depicting all associated sub – environments.

5.4 Palaeoflow directions, progradational sense and basin morphology for the Gydo-Gamka and Voorstehoek-Hex River Systems

Lithofacies successions associated within the Gydo – Gamka and Voorstehoek – Hex River Systems indicate a wave and storm dominated shoreline as the most likely palaeo – environment of deposition. Various sedimentary structures within lithofacies have been used to try determine the shoreline orientation of both systems as well as their progradational sense at time of sedimentation.

5.4.1 Shoreline orientation data from Lithofacies Association B and C

5.4.1.1 Shoreline orientation from wave ripple crest trends

Leckie (1988), Leckie and Krystinik (1989) and Duke, et al. (1991) have demonstrated the use of ripple crest trends of coarse grained ripples and wave ripples associated with HCS and SCS beds (Fig. 3.28). The

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crest trends of both ripple types have been demonstrated to have formed parallel to the palaeo-shoreline at time of sedimentation under combined oscillatory flow conditions. Coarse grained ripples and wave ripple crests in HCS and SCS beds form parallel to the shoreline trend since waves are assumed to approach the shoreline obliquely and refract perpendicularly. Ripple crest trends data presented in Figures 3.30, 3.32 and 3.45 – 3.47 from Lithofacies 4 and 5 respectively and have been combined in Figure 5.14. Ripple crest trends show a large degree of spread with the majority suggesting a NE - SW striking palaeo - shoreline with subordinate (but still significant) populations indicating possible NW – SE and ESE - WNW striking palaeo – shorelines for the Gamka Formation time sedimentation. These subordinate ripple crest trends may be interference ripples from another unrelated flow event.

5.4.1.2 Shoreline orientation from parting lineation trends

Parting lineations have been shown by Leckie and Krystinik (1989) to be oriented perpendicular to the palaeo – shoreline. Parting lineations in the Gamka Formation at Grootrivierhoogte show variable orientations (Fig. 3.31) with trends suggesting an E-W, ESE - WNW and N -S striking palaeo – shoreline at time sedimentation.

5.4.1.3 Shoreline orientation from asymmetric current ripples

Current ripples found in association with a “whale back” variant of Lithofacies 5 (Fig. 3.35) in the Gamka Formation at Grootrivierhoogte suggest a SW directed palaeoflow at the time of sedimentation and are interpreted to have prograded parallel to the palaeo – shoreline orientation by longshore drift.

The sum of shoreline orientation data from Lithofacies B and C therefore strongly suggest either a NE - SW or ESE – WNW striking palaeo – shoreline. The implication of these shoreline orientations are discussed on p. 123.

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5.4.2 Shoreline orientation data from Lithofacies Association D

5.4.2.1 Shoreline orientation from cross bed lee slope dip directions

Cross bed data from Lithofacies 7 of the Gamka and Hex River Formations at specific study sections is presented in Figures. 3.63 – 3.67 and has been combined in Figure 5.15. The majority of palaeo – flow vectors suggest a SSW, SW and S directed palaeoflow and hence a NNE to NE source for sediment.

5.4.2.2 Shoreline orientation from asymmetric current ripples

Asymmetric current ripples have been found in association with Lithofacies 7of the Gamka and Hex River Formations at various specific study sections (Figs. 3.69 – 3.70) and Lithofacies 8 of the Hex River Formation at Keurbosfontein (Fig. 3.81). Asymmetric current ripples indicate a variable palaeo – flow direction ranging from SW to SSE and are summarised in Figure 5.16.

5.4.2.3 Shoreline orientation from parting lineation trends

Parting lineation trends have been found in association with Lithofacies 8 of the Gamka and Hex River Formations at all three study sections (Figs.3.77 – 3.79). Cumulative parting lineation trends suggest a NNE – SSW trending palaeoflow trend and is summarised in Figure 5.17.

Data obtained from various shoreline indicators from Lithofacies Association D suggest two likely scenarios for shoreline orientation if the interpretation of L.A. D as a USF-Beach complex is correct. 1) If cross beds represent USF bars, these bars may have migrated in response to longshore drift parallel to the shoreline and flowed south to southwestwards at time of deposition. This would mean that the palaeo – shoreline orientation for the Gydo – Gamka and

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Voorstehoek – Hex River Systems would strike between NNE – SSW and NE - SW. Planar parallel laminated sandstone beds (Lithofacies 8) may represent (if the majority of trends are assumed to be derived from a USF sub – environment) and have deposited as bars parallel to the palaeo - shoreline. Asymmetric current ripples associated with Lithofacies 7 and 8 further support both notions and suggest that flow was directed SW to SSW. 2) If longshore drift currents were weak during sedimentation, sedimentary structures represented in Lithofacies 7 and 8 would be assumed to have prograded perpendicular to the shoreline at time of sedimentation. The palaeo – shoreline would thus have been oriented between ENE – WSW to NW - SE.

5.4.3 Combined approach to shoreline orientation and progradational sense of the Gydo – Gamka and Voorstehoek – Hex River Systems

Based on the sum of all evidence, two potential shoreline orientations appear to be possible for the Gydo – Gamka and Voorstehoek – Hex River Systems.

5.4.3.1 Northeast – southwest striking shoreline model (Fig. 5.18)

This model assumes that the most significant wave ripple and coarse grained ripple crest trends reflect the original shoreline strike orientation and is oriented NE - SW. Parting lineations, cross bed and asymmetric current ripple data reflect sedimentation by strong longshore drift currents redistributing sediment along the coastline causing the beach to migrate SW to SSW. The shoreline would thus prograde in the longshore drift direction and normal to itself. This model would suggest a N to NW source for sediment.

5.4.3.2 Northwest - southeast striking shoreline model (Fig. 5.19)

This model assumes that a less significant combined wave and coarse grained ripple crest trend reflects the original shoreline strike orientation and is oriented NW - SE. Parting lineations, cross bed and asymmetric current ripple data reflect sedimentation approximately

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perpendicular to the shoreline as and therefore assume longshore drift currents were weaker at time of deposition. Progradation of the shoreline will be perpendicular to itself and in a SW to SSW direction. This model suggests a N to NE source for sediment.

Both shoreline models infer a Namaqua – Natal Belt as an immediate source for sediment due to its proximity and major onlap relationship with the Namaqua – Natal Belt when the basin was active since outcrop of the Namaqua - Natal Belt is present to the north, north east and northwest of the Cape Supergroup. Shoreline orientation models presented in this project as well as assumed offshore directed sediment transport are supported by detrital zircon data by Fourie (2010) who present the notion of a Namaqua – Natal Belt source for sediment within the Clanwilliam Basin during Bokkeveld Group sedimentation. Panel maps presented by Theron (1972) representing thicknesses of individual formations of the Bokkeveld Group indicate that arenaceous formations thicken northwestwards whilst argillaceous formations thicken south eastwards for the Clanwilliam Sub-basin (Fig 1.13). This would suggest a NW source for sediment and onlap onto the Namaqua – Natal Belt. Alternatively, both shorelines are valid and a shallow embayment, in the sense of that proposed by Theron (1972) for the Clanwilliam Sub-basin is valid. In this model,s ediment would be supplied to the Cape Basin during Bokkkeveld Group sedimentaiton from the north, northeast and northwest and would still be supported by detrital zircon data of Fourie (2010) suggesting a major Namaqua – Natal Belt source.

5.5 Palaeoflow directions, progradational sense and basin morphology for the Tra-Tra – Boplaas System

Lithofacies successions of the Tra-Tra – Boplaas System suggest sedimentation within a wave and tide influenced estuary palaeo – environment. Delta and estuarine channels have been used to estimate shoreline orientation. Shoreline trends are assumed to be perpendicular to the average channel migration direction and assume minimal meandering. Lithofacies 11 have been interpreted to be channel sandstones and have

122 been found in interpreted channelised tidal flat environments (Lithofacies Association E) of the Tra-Tra Formation and sub-areal inner bayhead esturine channels (Lithofacies Association H) of the Boplaas Formation.

5.5.1 Shoreline orientation from lee slope dip direction of downstream accretion bars

Lee slope dip directions of downstream accretion bars from LF 11 of the Boplaas Formation have been recorded at Keurbosfontein and Wupperthal in LFA H (Figs. 3.99 and 3.100). Lee slope dip directions from both localities indicate a dominant SW and less significant S to SSE directed palaeoflow direction at time of sedimentation. This would suggest a NW – SE, or E - W striking shoreline at time of sedimentation.

5.5.2 Shoreline orientation from apparent gutter trends

Apparent gutter trends from Lithofacies 11 in LFA H at Keurbosfontein and Wupperthal (Figs. 3.101 and 3.102) suggest that channels trended NNE – SSW to N – S. This suggests an ENE – WSW to E - W oriented shoreline at time of sedimentation.

5.5.3 Shoreline orientation from rib and furrow structure trends

Trends of rib and furrow structures have been found in LF 11 within LFA E in the Tra-Tra Formation (Fig.3.104). These structures indicate that tidal channels trended NNW - SSE. This suggests a WSW - ENE trending shoreline at time of sedimentation.

5.5.4 Shoreline orientation from asymmetric current ripples

Asymmetric current ripples found in LF11 within LF E in the Tra-Tra Formation indicate a SSW to SW directed palaeoflow (Fig. 3.103 )

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within tidal channels. This suggests an E-W to ESE - WNW trending shoreline at time of sedimentation.

Combined shoreline orientation data from Lithofacies 11 indicate that the Tra- Tra – Boplaas System was possibly oriented E – W at time of sedimentation. This shoreline orientation thus assumes that sediment for the Tra-Tra Boplaas Systems was sourced from the north. This shoreline orientation model would assume a southward prograding Tra-Tra – Boplaas System and would still be supported by detrital zircon data by Fourie (2010) who present the notion of a Namaqua – Natal Belt source for sediment within the Clanwilliam Basin during Bokkeveld Group sedimentation.

5.6. Palaeontology

5.6.1 Preservation and Taphonomy

Most of the body fossils discovered in the study area in the course of this study are disarticulated and articulated fossil specimens were relatively rare. Complete fossils have only been found in Lithofacies Association A and Lithofacies 9 (L.A. F). The incomplete and disarticulated nature of trilobite fossils is most likely due to the fact that trilobites are ecdysozoans and undergo regular ecdysis shedding their exoskeletons during growth and development and that sclerites readily disarticulate from their life position after death (Clarkson, 1998, Brusca and Brusca, 2003). Brachiopod and bivalve fossils too may have disarticulated from their life positions soon after death. Trilobite sclerites, brachiopod and bivalve valves, gastropod and tentaculitid conches and conulariid polyps were probably distributed further within the basin with passing water currents attributed to storm, wave and unidirectional bottom flow conditions.

Body fossils are preserved as any one of three preservation regimes the study area, these are namely cast, mold and nodular preservation. Within L.A. A, all three preservation styles co-occur. In L.A. B, C and F only cast and mold

124 preservation styles are evident. Body fossils are completely absent in L.A. D, E, G and H.

Browning (2009) attributed nodular preservation of fossils, often in association with authigenic pyrite, to early diagenesis with fluctuating sea levels and concomitant seawater geochemistry. This diagenetic origin for nodular preservation also assumes an epeiric seaway model since fluctuating sea levels, even minor, would greatly affect normally oxic and anoxic environments resulting in alteration of local eH and pH conditions and encouraging nodular growth in addition to encouraging reworking of nodules, a trend noted by Theron (1972) and Browning (2009). Nodular preservation of fossils within the study area always occurs in close association with stratabound carbonate nodules and is abundantly present in Lithofacies Association A. The reason for this association together has been attributed to early diagenesis by Tankard and Barwis (1982) and Browning (2009).

Fossil elements found in Lithofacies Association A and Lithofacies 9 (L.A. F) are nearly always complete with little to no damage, whereas fossil elements in L.A. B and C are always highly damaged. This may be attributed to local energy conditions within the basin where these metazoans were fossilised. L.A. A has been demonstrated to represent a low energy Os sub – environment below mean storm wave base. The relatively calm and persistent conditions thus allowed for the high degree of preservation of fossils with minimal to no damage. Lithofacies 9 has been interpreted to be a washover flat environment which entered into a lagoon environment. Sudden back barrier flooding events are envisioned to have smothered life communities of Australospirifer sp. indet. preserving them in their life positions (Fig. 3.85 A) with fine detail preservation of brachiopod lophophores and brachidiums (Fig. 3.85 B). The demonstrable storm and wave dominated conditions which prevailed during sedimentation of Lithofacies Association B and C accounts for the fossils being present as bioclastic storm lags on lower bedding planes of HCS and SCS beds (Fig.3.24, 3.26 and 3.27) as well as sand supported coquinites (Fig. 3.21). Turbulent storm surge conditions in Lower Shoreface sub – environments are envisioned to have entrained

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metazoans from their respective habitats (either before or after death) leading to them being highly damaged by abrasion and attrition before coming to rest.

Plant fossils are present either in L.A. A, C, D, F and G. Plant fossils are always incomplete and restricted to fragmentary axes and thalli (Fig. 3.15, 3.79, 3.84, 3.89, 3.90). Fossil plant fragments are preserved either two dimensional, flat relief coalified compressions or as three dimensional coalified body fossils. Damage and disarticulation of plants before fossilisation appears to be a reflection of transport after death.

5.6.2 Palaeo - ecology and Biostratigraphy

The accepted palaeo – ecologic model of Hiller and Theron (1988) assumes that benthic marine metazoan communities were intimately associated with discrete sub – environments in a wave and storm influenced deltaic system. This study has demonstrated a storm and wave dominated shoreline for the depositional environment of the Gydo – Gamka and Voorstehoek – Hex River Systems, and a wave and tide influenced estuary for the the Tra-Tra – Boplaas System. Findings in this project suggest that marine species were entirely restricted to Os sub – environments.

All metazoan fossils found within the study area are those of recognised Malvinokaffirc Realm marine organisms and are restricted to the Gydo, Gamka and Voorstehoek Formations (Appendix 1 - 4). Metazoan fossils are most diverse in L.A. A, in particular Lithofacies 1 which is present only in the Gydo Formation (Appendix 1 - 4), and are also present, but not as abundant, in Lithofacies 2. In the Voorstehoek Formation within the study area, metazoan fossils strangely are rare in comparison with its lithofacies equivalent in the Gydo Formation and almost contradict the species richness portrayed by Oosthuizen (1984). This may in part be due to a preservation or non – intensive collection bias and needs to be addressed and/or the inferred proximal position of the Voorstehoek Formation in the study area.

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The findings of this project suggest a strong facies control on fossil occurrences and a major Os sub – environment dominant habitat in the Ceres Subgroup. This supports the benthic lifestyle mode for the Malvinokaffric Realm fauna as portrayed by Hiller and Theron (1988), Boucot (1999). The common disarticulated nature of invertebrate fossils suggests a high degree of reworking of elements after death. This may influence any attempt at biostratigraphy and species richness indicators since the elements are scattered throughout the sub – environment.

In L.A. B and C, the only fossil taxa present are the brachiopods Australospirifier sp. indet. and Australocoelia cf. tourteloti, and fragmentary crinoid ossciles. These fossils are always disarticulated with varying degrees of damage. The original habitat of these brachiopods and crinoids may be misleading as they are out of context. It is considered that fossils within L.A. B and C may have been transported from the Os environment by storm and wave action (as indicated by the disarticulated and high degree of damage of bioclasts) and thus do not represent their typical habitat occurrence. Fossils are totally absent from L.A. D, this may be attributed to non – intensive fossil collection, or simply that there were no biota in these habitats.

In L.A.F and H, plant fossils and plant rootlet ichnofossils are numerous and reflect their habitat in fresh water to mixed salinity associated inner and middle estuarine environments. Plant fossils and plant rootlet ichnofossils in L.A. D indicate a possible Bs habitat for plants.

5.7 Sequence Stratigraphy and Basinal Correlation

5.7.1 Introduction to Sequence Stratigraphy

Sequence stratigraphy has been used as a tool to provide high resolution correlation between and among lithostatigraphic units and allows for analysis of depositional systems (depositional environments) over time. This is achieved by compartmentalising depositional systems within parasequences,

127 sequences, the systems tracts to which they belong and the genetically related sequence stratigraphic surfaces which bind them (Van Wagoner, et al., 1990, Reading and Levell, 1996, Catuneanu, 2006, Embry, 2009). Depositional systems are strongly controlled by base level change. Sequence stratigraphy thus provides a time-wise sequence of basin evolution. Basinal events, basin conditions and depositional histories may be examined and may be interpreted in terms of base level change (Van Wagoner, et al., 1990, Reading and Levell, 1996, Catuneanu, 2006, Embry, 2009). Sequence stratigraphy therefore is a powerful tool for facies analysis.

Sequence stratigraphic surfaces (or sequence boundaries) mark significant changes in lithofacies and palaeo – environments and are a means of compartmentalising separate depositional sequences and occur where a change in depositional trend occurs, i.e. a change from a fining upward succession to a coarsening upward succession, or a shallowing upward succession trend to a deepening upward trend (Embry, 2009). Such changes in depositional trend may be expressed by sequence stratigraphy in terms of regression, transgression, or in terms of base level fall and rise over time (Embry, 2009). Similarities in sequence boundaries and the genetically related successions they envelope, may be used in basin wide correlations since it is assumed that the same event affected the entire basin and at the same time (Embry, 2009).

The understanding of the term ‘Sequence’ varies among authors. By its original definition, a ‘sequence’ is an unconformity bounded stratigraphic unit that is bound by sub-aerial unconformities (Vail, et al., 1977, Embry, 2009). It is cautioned that proximal basin edge depositional systems are more susceptible to base level changes, and thus sub-aerial unconformities are both expected and more restricted to these regions and decrease when traced distally along depositional dip to basin centres (Catuneanu, 2006, Embry, 2009). Distal basin centres are less susceptible to base level changes (Embry, 2009). Thus unconformities are uncommon in distal basin centres, if base level change is not drastic. Time equivalent conformities in basin centres are expected for their proximal basin edge correlates (Van Wagoner, et al.,

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1990, Catuneanu, 2006, Embry, 2009). The definition of a depositional sequence may therefore be defined as ‘genetically related strata bounded by surfaces of erosion or non-deposition and/or their correlative conformities been modified to include distal genetically related ‘correlative conformities’ (Catuneanu, 2006). System tracts, parasequences and parasequence sets are components of sequences (Van Wagoner, et al., 1990). Embry (2009) thus favours a ‘sequence’ to rather be defined as a ‘depositional sequence’ which shall be used in this project.

‘Parasequence’ has been defined by Van Wagoner, et al. (1990) p.8 as a “relatively conformable, genetically related succession of beds or bedsets by marine – flooding surfaces (F.S.) or their correlative surfaces.” This usually is placed at a sharp facies shift from arenaceous to argillaceous lithologies. Embry (2009) argues that marine - flooding surfaces are poor sequence stratigraphic boundaries since they are highly diachronous across the boundary. Embry (2009) favours that parasequences should be bound by surfaces with low diachroneity and favours maximum regressive surfaces as parasequence boundaries. The definition of a parasequence according to Embry (2009) p. 59 is thus “a small-scale sequence stratigraphic unit bound by maximum regressive surfaces (M.R.S.) and their correlative surfaces.” Embry (2009) has noted that M.R.S. often correlates with shoreline ravinement unconformities (S.R – U) as a bounding surface. It is argued that if a parasequence is bounded by upper and lower correlative S.R. – U. as M.R.S. then the parasequence represents a depositional sequence. ‘Parasequence sets’ are defined as “a succession of genetically related parasequences that form a distinctive stacking pattern, bounded, in many cases by major F.S. and correlative surfaces” (Van Wagoner, et al., 1990) p. 17. Parasequences referred to in this project shall follow the scheme presented by Embry (2009)

‘Systems tracts’ are defined as a linkage of contemporaneous depositional environments and is a component of a sequence which is bound by sequence stratigraphic surfaces (Van Wagoner, et al., 1990, Embry, 2009). Numerous systems tracts have been presented by various workers. Each scheme has

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defined criteria pertaining to their bounding sequence stratigraphic surfaces based on whether they are material or time based the arguments for, and against the different schemes is addressed in Embry (2009) and are summarised in Figure 5.20. This project follows the two systems tract method of Embry (1993) which recognises regressive systems tracts (R.S.T.) and transgressive systems tracts (T.S.T.). R.S.T. are defined as a sequence stratigraphic unit bounded by a lower M.F.S. and its correlative surfaces and an upper M.R.S (Embry, 1993). R.S.T. is equivalent to Van Wagoner et al. (1988) L.S.T. and H.S.T (Embry, 2009). T.S.T. are defined as sequence stratigraphic units bound by a lower M.R.S. and an upper M.F.S. and its correlative surfaces (Embry, 1993).

5.7.2 Sequence stratigraphy of the Ceres Subgroup

Sequence stratigraphic surfaces, systems tracts and 4th and 3rd order stacking patterns have also been indicated in Appendix 1 - 4 and in Figure 5.21. A total of 5 parasequences have been recognised within the Ceres Subgroup and have been referred to in their numerical order from the base of the Gydo Formation to the top of the Boplaas Formation. Interpreted palaeo - environments have a remarkable lateral continuity across the entire outcrop area and follow their respective bounding sequence stratigraphic surfaces and stacking patterns (Appendix 5 and 6). Only 3rd order stacking trends representing transgressive-regressive (T-R) cycles have been used in order to address sequence stratigraphic surfaces and identify sequences and parasequence..

In the Gydo – Gamka System, M.R.S. have been selected to occur only within (and probably towards the top of) L.A. D (at Keurbosfontein and Wupperthal only) and L.A. C (at Grootrivierhoogte only) since they represents their most regressive environments, a USF – Beach Complex and pLSF respectively. The base of the Gydo Formation is assumed to represent the M.F.S since L.A. A has been interpreted to represent an Os sub – environment. The Gydo – Gamka System is thus regressive overall and a typical R.ST. It is thought that an M.R.S. is present in the underlying Rietvlei Formation, if this is

130 demonstrable; the 1st parasequence (parasequence 1.1) would extend from the Rietvlei Formation to L.A. D and L.A. C of the Gamka Formation and would represent both the 1st T-R cycle and major flooding event of the Cape Basin during Bokkeveld Group time deposition. If an M.R.S. is present in the Rietvlei Formation and is demonstrated to represent an SR – U, this would mark a 2nd order sequence boundary and the absolute base of Bokkeveld Group related transgression within the Cape Basin.

The Gamka - Voorstehoek transition is represented by a T.S.T and shows a gradation from either USF – Beach Complex (at Keurbosfontein and Wupperthal) environments or dLSF (at Grootrivierhoogte) in the Gamka Formation to Os sub – environments in the Voorstehoek Formation. Three discrete parasequences (parasequences 1.2 – 1.4) and T-R cycles are present from the Gamka - Voorstehoek transition to the top of the Hex River Formation. These parasequences may be traced out across the entire outcrop area with a high degree of confidence (Appendix 5 and 6). Selected M.R.S. in the Voorstehoek is thought to occur within L.A. B since it represents a more regressive dLSF environment as opposed to L.A. A, an Os sub – environment. The sharp, planar and abrupt onlap of the argillaceous Tra-Tra Formation (L.A. E) on the coarse and arenaceous Hex River Formation (L.A. D) has been identified as possibly being a paraconformable contact and is diagnostic of an S.R – U. This is supported by the palaeo - environmental interpretation of L.A. E as forming in response to transgression (as well as the overlying L.A. F, G and H) and L.A. D representing a typical regressive environment.

The ramifications of this sequence stratigraphic surface suggest that the Hex River, Tra-Tra contact represents a large sequence boundary and a discrete 2nd order transgression into the Cape Basin marking the start of Tra-Tra – Boplaas time sedimentation. It is hypothesised that 1st parasequence above this 2nd order sequence boundary (parasequence 2.1) within the Bokkeveld Group is entirely associated within the Tra-Tra – Boplaas System and continues into the overlying Waboomberg and Wupperthal Formations (both in the Bidouw Subgroup) and represents a part of the 5th T-R cycle. Most of

131 the Waboomberg Formation has been demonstrated to possibly represent an Os sub – environment based on marine Malvinokaffric Realm Fauna (Theron, 1972, Cooper, 1986) and combined with the Tra-Tra – Boplaas System could represent a large T.S.T with a R.S.T extending somewhere within the overlying arenaceous Wupperthal Formation. This parasequence would represent a large transgression within the Bokkeveld Group comparable with the Rietvlei, Gydo transition and needs to be addressed, if proven correct, it would support hypothesised sea level curves and T-R cycles predicted by Cooper (1986).

5.7.3 Towards a Devonian Sea Level Curve for South Africa

The presented sea level curve is representative for Ceres Subgroup sedimentation only (Fig. 5.21). The transition between the Rietvlei and Gydo Formations is entirely hypothetical and is based on the assumption that the Rietvlei Formation represents sedimentation within a mixed fluvial – beach environment (strandplain) after Theron and Basson (1989). This sea level curve closely mimics that produced by Cooper (1986). The presented sea level curve and lithofacies data suggests that Gydo – Gamka System sedimentation was deposited during a period of eustatic sea level high with little to no discernable fluctuation. The Voorstehoek – Hex River System represents sedimentation during a sea level high (although not as great as the Gydo – Gamka System) with a great degree of sea level fluctuation. Three sea – level rises and falls appear to have been present at time of Voorstehoek – Hex River sedimentation. The Tra-Tra – Boplaas System records a large transgression and sea level rise hypothesised to have continued into Waboomberg, or even Wupperthal Formation time sedimentation.

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Chapter 6 Conclusions

From fieldwork intensive research on the sedimentology, stratigraphy and palaeontology of the Ceres Subgroup, a total of 12 lithofacies and 8 lithofacies associations have been recognised and described within the study area. Based on comparison with modern and ancient marginal – marine and shallow – marine depositional environments, the lithofacies associations recognised in this project are representative of sedimentation within 8 distinct sub – environments. These sub – environments, in turn, may be grouped into 2 larger environments of deposition, namely a wave and storm dominated shoreline and a wave and tidal influenced estuary. The lateral continuity and almost tabular nature of Lithofacies Associations and their inferred sub palaeo – environments across the entire outcrop area suggest sedimentation was within a broad and shallow eperic seaway.

From wave ripple and coarse grained ripple crest trends, parting lineation trends and lee slope dip directions of asymmetric current ripples and cross beds two potential palaeo – shorelines for the Ceres Subgroup have been identified during Gydo – Gamka and Voorstehoek – Hex River System sedimentation, these are a NE – SW and a NW – SE trending shoreline.

Majority palaeo – shoreline constraints from wave ripple and coarse grained ripple crest data suggest a NE – SW orientation for the Gydo – Gamka and Voorstehoek –

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Hex River Systems. In this scenario, lee slope dip directions of asymmetric current ripples and cross beds suggest distribution of sediment and progradation of systems in response to strong longshore drift currents trending sub – parallel to the shoreline at time of sedimentation in a southwesterly to south-southwesterly direction. This shoreline trend is supported by geochemical work and inferences of provenance areas by Fourie (2010) who demonstrate a Mesoproterozoic aged source, most likely being from the Namaqua - Natal Belt based on flow vectors published by Theron (1972). If the shoreline orientation presented within this project is correct, it is suggested that sediment may have been derived from the Namaqua - Natal Belt, or an equivalent in South America. This hypothesis is based upon data from fence diagrams by Theron (1972) which show a continuity of all lithological units northwestwards to their outcrop limits with gradual thinning of argillaceous units, but no pinching out. This suggests that each lithological unit of the Bokkeveld Group onlapped onto the underlying Nardouw Subgroup. If this was the true basin margin of the Bokkeveld Group, a solitary thick arenaceous unit is assumed to have been present. Thicknesses of arenaceous formations recorded in this project and by Theron (1972) also show a NW thickening trend, supporting a NW – SW shoreline trend and N to NW sediment source.

NW – SE trending shorelines would suggest that sediment was distributed normal to the shoreline and not redistributed by longshore drift currents. Progradation would still be in a southwesterly to south-southwesterly direction. Sediment in this scenario would be derived from a NE source area. This shoreline trend and progradation sense scenario is also supported by Fourie (2010) who identifies a major Namaqua – Natal Belt provenance area in the Clanwilliam Sub – basin at time of Bokkeveld Group sedimentation. Alternatively both NE – SW and NW – SE trending shorelines are permissible and the Clanwilliam Sub – basin was a shallow embayment as depicted by Theron (1972).

Shoreline indicators form the Tra-Tra – Boplaas System indicate a possible E –W orientation of the shoreline at time of sedimentation and is still supported by detrital zircon data by Fourie (2010) who present the notion of a Namaqua – Natal Belt source for sediment within the Clanwilliam Basin during Bokkeveld Group sedimentation.

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A total of five parasequences are present within the Ceres Subgroup. These are bound by M.R.S. sequence stratigraphic surfaces and have been associated with 3rd order T-R cycles. Parasequences 1.1 – 1.4 represent T-R cycles within the Gydo – Gamka and Voorstehoek – Hex River Systems and are associated with the 1st 2nd order flooding event into the Cape Basin marking the start of Bokkeveld Group sedimentation. The Gydo – Gamka and Voorstehoek – Hex River Systems are regressive and conformable with each other with the Gamka-Voorstehoek transition representing a T.S.T. The contact between the Hex River and Tra-Tra Formations is paraconformable and appears to represent a S.R. – U associated with a large and 2nd 2nd order flooding event within the Cape Basin. The Tra-Tra – Boplaas System is a part of parasequence 2.1 and is overall transgressive representing a T.S.T. which is hypothesised to continue into the Waboomberg Formation and terminate at an M.R.S. somewhere in the Wupperthal Formation. The sequence stratigraphic framework for the Ceres Subgroup presented in this project shows support for the hypothesised eustatic trends for the South African Devonian sea-level curve by Cooper (1986) which show a large global eustatic rise during the late Emsian (Gydo – Gamka System sedimentation, possibly starting within the Rietvlei Formation) and Eifelian (start Tra-Tra – Boplaas System sedimentation).

Palaeontological finds are most numerous in L.A. A and hence suggest a major offshore lifestyle for Malvinokaffric Realm Biota within the Clanwilliam Sub – basin. Fossils found within this sub – environment are almost always incomplete, but with little to no damage suggesting a calm environment with some reworking of shelly material by weak currents. Lithofacies 1 has the highest diversity of fossil fauna and is only present in the Gydo Formation and is thought to represent a distal Os sub - environment. Lithofacies 2 is found in both the Gydo and Voorstehoek Formations, with fossils again being common in the Gydo Formation but rare in the Voorstehoek and is thought to represent a proximal Os sub - environment. It is suggested that fossils are rare in L.F. A of the Voorstehoek Formation because this is an extreme proximal equivalent of the Os sub – environment.

Fossils in LFA B and C tend to be damaged and incompletely preserved are restricted to Australocoelia cf. tourteloti and Australocoelia sp. indet. and crinoid

135 ossicles. Damage to shells is interpreted to be in response to turbulent hydrological conditions in response to storm and wave activity in a demonstrable OT – dLSF and pLSF sub – environments. It is thought that fossils may have been entrained either from the immediate OT – dLSF and pLSF environment, or were entrained from the Os sub – environment and reworked within wave and storm flows and deposited in OT – dLSF and pLSF sub – environments.

Plant fossils found in L.A. A, B and C are interpreted to be allochthonous and transported to the respective Os, OT – dLSF and pLSF sub – environments during offshore directed storm and wave return flow events.

In washover flat sub – environments (L.A. F), brackish water conditions appear to have been prevalent based on the presence of marine Australospirifer sp. indet. fossils (often within communities) and Skolithos sp. and other surface trails in addition to terrestrially derived plant fossils and plant rootlet ichnofossils. Wasover fans are suspected to have been deposited in response to washover events from strong waves breaking through the estuary barrier into the lagoon. Plant rootlet ichnofossils as well as the highly bioturbated nature of L.A. F suggest that these washover fan events were periodic and conditions were calm enough to support marsh development as well as bioturbation.

A biostratigraphy for the Bokkeveld Group does not appear possible as the majority Malvinokaffric Realm Biota are strongly facies controlled and exist en masse within an Os sub – environment.

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References

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