African Basins, Volume 3 (Sedimentary Basins of the World)

SEDIMENTARY BASINS OF THE W O R L D An Introduction to the Series Etymology reveals much about the essence of a word. S...

Author: R.C. Selley

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SEDIMENTARY BASINS OF THE W O R L D An Introduction to the Series

Etymology reveals much about the essence of a word. Science in German, Wissenschaft, is the art of observing whereas science in Chinese, koxue, is the study of classifying. Scientific observations, with the help of modem equipments, have made leaps and bounds in our century, but taxonomy seems irrelevant. Classification can be science. The rise of the natural sciences in Europe could be traced back to Carl Linnaeus in 1750 when he used criteria of mutual exclusiveness to establish the taxonomy of living organisms. Unfortunately, this prerequisite in dividing and subdividing is not always appreciated, and a common practice in geology has been to "classify" basins through reference to incidental attributes. So, we have coastal basins, back-arc basins, extensionally rifted basins, successor basins, deep-sea basins, flysch basins, etc. These qualifications describe the geography, tectonic setting, principal stresses, orogenic chronology, depositional environment or sedimentary association of a basin, but these all could be different aspects of one and the same. "A basin is a basin is a basin", paraphrasing Gertrude Stein. Even though no two basins are exactly alike, subsidence is the common denominator of all. A systematic classification of basins depends upon recognition of the mutually exclusive causes of subsidence. Forty years ago, when I first went to study in the United States, I was fascinated by the debate whether basin subsidence is isostatically induced by sedimentary load. Later, in the 1950's, I began to realize that depressions on Earth are underlain by thin crust and came to the conclusion that subsidence could be the surface manifestation of an endogenetic process of crustal thinning, in response to the Airy isostasy. Still later, in the early 1960's, after geophysical studies had revealed the heterogeneity of the Earth's mantle, subsidence could be related to mantle cooling and corresponding density change, in response to the Pratt isostasy. Meanwhile, we all concede that the weight of a sedimentary pile filling a subsiding depression will induce isostatic subsidence. The three different mechanisms of isostatic adjustment are, however, not mutually exclusive, and they may have operated concurrently. To classify basins on the basis of the various modes of isostasy can thus not be systematic. But, as indicated by gravity studies, not all sedimentary basins are isostatically adjusted, and not all subsidence is isostatic. For a first-order division we could thus recognize two classes of basins, those which have subsided isostatically and those which have not. Crustal thinning is a prerequisite to initiate Airy isostasy. Thinning can be induced by two different systems of principal stresses, with the least principal stress either horizontal or vertical. The former leads to the genesis of rifted basins under horizontal extension, and the latter results in pull-apart basins in transform or strike-slip fault zones. The orientation of principal stresses could thus be the criterion for second-order distinctions. Rift basins are commonly present in the continental interior. After the appearance of oceanic crust between separated continents, the rifts become narrow oceanic gulfs (like the Red Sea). Eventually the loci of subsidence are shifted to passive margins where coastal plains are underlain by thick basinal sediments. Rifted basins may also form on an active margin where a segment of continent is torn apart from the mainland to form islands arcs; those are back-arc basins. The position with respect to plate margins can thus be used as the criterion for third-order distinctions. Subsidence induced by horizontal compression, the third of the three possible configurations of principal stress, is not isostatic. Plate-tectonics theory relates the origin of trenches to the underthrusting of ocean lithosphere on an active plate margin and the origin of foreland basins to the underthrusting within the continental lithosphere. These two major basins of compressional origin are thus also distinguished by the third-order criterion concerning their position with respect to plate margins.

VI

SEDIMENTARY BASINS OF THE WORLD - - AN INTRODUCTION TO THE SERIES

When we first planned the series of the Sedimentary Basins of the World, we intended to adopt a genetic classification. Basins are subdivided into the three sets of criteria discussed above: 1. Isostastically adjusted basins 1.1. Extensional basins 1.1.1. Rift basins in the continental interior 1.1.2. Narrow oceanic gulf basins 1.1.3. Basins of deposition on passive margins 1.1.4. Rifted basins on active margins or back-arc basins 1.2. Transcurrent (transform or strike-slip) pull-apart basins 2. Isostastically not adjusted basins 2.1. Compressional basins 2.1.1. Foreland basins (in a continental interior) 2.1.2. Oceanic trenches (on a continental margin)

With this scheme in mind, the editors of Elsevier and I made up a list of the volumes for the projected series, and we started our search for volume editors. Our first priority, taking into account current demand, was to bring out a volume on China. We were concerned that the Chinese basins could not be fitted into our scheme, because we were told that they represent a special group of unclassifiable basins on "paraplatforms". However, as I became personally involved in researches on the geology of China, I came to the conclusion that the Chinese basins were not "unclassifiable". Rifted basins, back-arc basins, pull-apart basins, foreland basins, etc., exist in China, as they are present elsewhere in the world. The Chinese basins of different origins do share a common history in geologic evolution, and they are united by their geography. It would be illogical to discuss the various Chinese basins in separate volumes. Yet if we are to include all of them in one, we have to designate that volume by their unifying geography. When we started to work on our second volume on rifts, we were still trying to keep our genetic scheme, although we were resigned to make a single exception for the Chinese opus. We were thinking of the East African Rift Valleys, and the emphasis was on Africa. As chance would have it, I just happened to accept a consulting contract on Africa. After a year of working on the assignment, I realized that basins on that continent are as diverse in origin as those in China; there are foreland basins, pull-apart basins, as well as rift basins in Africa. About this time, our choice of the editor for a volume on rifted basins, Professor R.C. Selley, brought up another issue. He pointed out to me the impracticability of putting out volumes on the sole basis of their postulated origins. The purpose of the series is to provide information on the geology of sedimentary basins of the world in order to help a novice to start a project. Commonly, the one who seeks information knows the geographic extent of his interests, but not necessarily the genesis of his targets. Taking, for example, the case of a person who is to start an exploration venture in some region, how should he know if he is to study a monograph on pull-apart basins or one on foreland basins. He knows, of course, if the location is in China or in Africa, and could consult an opus on Sedimentary Basins of China, or that on Africa accordingly. The arguments by Professor Selley finally convinced me to change our scheme. The criterion of dividing the volumes will have to be geographical. In addition to those on China and Africa, volumes on sedimentary basins of Australia, South America, Central America, and the Soviet Union (Russian Platform and Siberia) are planned. Geographical groupings are satisfactory if the basins of various origins in a region share some common heritage, but to throw all heterogeneous entities into one big pot could be disconcerting. To produce, for example, a volume on the Sedimentary Basins of Europe to include all those in the Russian Platform, under the North Sea, and in the Prealps may make a good lexicon, but not an opus harmonized by a unifying theme. This consideration led us to the decision to place priority on certain natural boundaries, so that each volume would sustain a certain coherence in geology as well as in geography. The Cenozoic basins in the Tethyan orogenic belt, for example, are to be grouped under the title of Foreland Basins of the Alpine-Mediterranean Region. The pull-apart and back-arc basins on the shores of the Pacific will be included in two or more volumes of the

Sedimentary Basins of the Circum-Pacific Region. There are, of course, other sedimentary basins of the Earth, especially those of the Near East, North America and Antarctic, which should be included in the series if the coverage is to be complete. On the other hand, we shall also evaluate the demand of the profession for such volumes as the initial volumes of the series successively appear during the next decade.

SEDIMENTARY BASINS OF THE WORLD - - AN INTRODUCTION TO THE SERIES

VII

It is my hope that the series of volumes would not be compared to philately albums; there should be unity in style and in substance. Yet the accumulation of geological information has reached such immense proportions since Eduard Suess wrote his monumental work Das Antlitz der Erde that no single person could ever hope to master the geology of the world. The series of our volumes will, therefore, have to be collective efforts. Coordinations by volume editors are indispensable. My job, as the series editor, is to further enhance the unity and harmony of the whole. We have, however, to accept the fact that each article of a volume may "speak" a different dialect, and each volume of the series may "speak" a different language. Perfect consistency can only be achieved if a person has the time or the capacity to translate all those hundreds of articles in more than a dozen volumes into one universal script. This is not possible, and the practical alternative to the ideal is, therefore, to leave each author or group of co-authors a maximum freedom in their style of presentation and in their interpretations of geology. The articles are to be accepted as expressions of the present state of understanding of an area by leading geologists working in that area. They may or may not represent the understanding of the volume editor, or that of the series editor. Through my experiences in editing the first volume on Chinese basins, I appreciate the potential dangers of such freedom of expressions; a lack of precision in semantics could lead to grave misunderstandings. I felt impotent when I saw basic terms, like orogeny, platform, shelf sediments, intracratonic basins, etc., defined, in certain communities of our profession, on a basis distinctly different from that adopted by modem students of geology. The misconceptions in some instances are so deeply rooted, that nothing short of a rewrite could save the situation. Yet neither the volume nor the series editor could completely revise all the articles. To avoid complete chaos, I plan, therefore, to write a summary, at the end of each volume (or at least some), in my style, and to interpret the geology on the basis of my understanding. Such summaries may contain an overdose of personal opinions and may involve interpretative errors by a single geologist, but they should, at least, be consistent, and may eliminate misconceptions caused by the divergent meaning of the same words as they are used in various "dialects" and "languages". The preceding pages were written in January, 1988 for the first volume of the series on Chinese Basins. The reviewers of the Chinese volume gave me encouragement that I, as the series editor, should continue my role as an interpreter of different cultures. I have, therefore, taken the initiative to write the last chapter of the second volume m A Distant View of the South Pacific Geology. The volume edited by Peter Ballance and my summary are both written in English. There is little need for translation. Nevertheless, New Zealanders speak English with their local accent, and the same "words" are pronounced differently by one who speaks English with a strong Chinese and Swiss accent. It is not surprising that Peter, a dear old friend, found it difficult to "agree entirely" with me. On the other hand, he was tolerant enough not to protest too strongly, and I could have my Distant View for the reference of other distant readers. In reviewing the articles in the second volume of our series, I became more convinced than ever of the wisdom of Selley's advice that the basins should be grouped regionally. I have devoted most of my professional career with a process-oriented approach. When I edited a book on Mountain-Building Processes, the emphasis was on processes. Instead of regional syntheses, I adopted an analytical approach to look into the different processes involved in mountain-building, sedimentary, magmatic, deformational, and metamorphic. In editing this series on Sedimentary Basins, there is no better alternative than regional synthesis. Not only basins in China have diverse origins, those in the South Pacific are equally diverse. Yet they all seem to belong to the same set of diverse basins. If the geology of the South Pacific seems very different from that of China, the apparent distinction can be attributed to the fact that they have advanced to different stages of tectonic evolution. Four years have gone by since the publication of the second volume of the series of The Sedimentary Basins of the WorM. On the eve of the publication of the third volume, I am encouraged to find that the series will not be philately albums; they will not be collections of random observations. I saw the parallelism in the pattern of orogenic deformations in China and in South Pacific, and I could see the same pattern in Africa. There is a difference in the stage of the tectonic evolution. China, as a part of Eurasia has undergone a billion years of amalgamation. The South Pacific is still in an earlier phase of accretion. Africa has gone a long away since its separation of Pangaea, 7and it is being pushed toward Eurasia to its ultimate destiny of a place in a supercontinent. The figures and the colorations of the mosaic pieces are different, but they are all to be pieced together for a unifying theory of global tectonics. The next two volumes will be The

VIII

SEDIMENTARY BASINS OF THE WORLD m AN INTRODUCTION TO THE SERIES

Sedimentary Basins of the Former Soviet Union. I expect to find the same manifestations of the fundamental principles which give us The Face of the Earth. Golden, Colorado, April, 1997

KENNETH J. HSLr Series Editor Sedimentary Basins of the World

SEDIMENTARY BASINS OF AFRICA Introduction and Acknowledgements I speak of Africa and golden joys W. Shakespeare - - Henry IV, Part 2. V, iii, 100

I was very badly brought up as a boy, being allowed to read extensively from the currently politically incorrect works of Rudyard Kipling and Rider Haggard. Whatever else these authors may have done to me, they imbued in my boyish mind a deep fascination for Africa and all things African. One of several reasons that compelled me to become a geologist was the opportunity that it provided to travel the world in general, and Africa in particular. I first landed in Africa in 1963, have intermittently visited the continent many times, and lived there for several years. My particular interest has been in petroleum exploration, but I have also had assignments on gold and coal exploration. From the days of the Roman Empire onwards European explorers have constantly made new discoveries about Africa, as Pliny wrote in 78 AD, "ex Africa semper aliquid novi" (there is always something new out of Africa). Geologists, however, commonly remark on the uniformity of much of Africa's geology. Vast areas are occupied by fiat, though high, plains of Precambrian basement, composed of igneous and metamorphic rocks of immense and diverse ages. These basement rocks are locally overlain by a cover of shallow marine and continental sediments of Cambrian to Recent age. One of the points documented in this book is the regional uniformity of Lower Palaeozoic stratigraphy. This is best seen when tracing the outcrop from the Atlantic coast of Mauritania in the west, eastwards through Algeria and Libya (Chapter 1). Indeed this stratigraphy can be traced, with little variation, into Arabia. This similarity is not difficult to explain, because this terrain formed the southern shores of the Tethys, the ancient ocean that lay to the north of the old Gondwana continent. Geologists who have hammered the Lower Palaeozoic rocks of north Africa, however, will also feel at home in the Cape of Good Hope. The unconformity at Sea Point, described by Playfair and Hall in 1815, and again by Charles Darwin during the Beagle cruise of 1831-1836, invites direct comparison with the pan-Saharan Tassilian discordance whose regional significance was recognised by Kilian (1922). The overlying facies in the Table Mountain Group, with its braided alluvium, and Tiggillitiferous shallow marine sands capped by Silurian tillites, is directly comparable with the Lower Palaeozoic sediments of the Sahara (Tankard and Hobday, 1977). The sedimentary basins of Africa are largely of two types, sag basins, and failed rifts (Clifford 1986). Some of the sag basins, such as the Karoo, are clearly syn-depositional in origin, with clear evidence that subsidence was coeval with, and exerted a strong control on, sedimentation (Chapter 12). Many of the sag basins, however, notably those of north Africa, are clearly post-depositional in origin. The uniform stratigraphy of basins, such as those of Algeria, shows that their sediments were laid down on the gently sloping Saharan platform. Their present basinal shape developed during the late Carboniferous "Hercynian" tectonic phase. Similarly the Murzuk and Kufra basins of southern Libya, if they existed at all, were northerly plunging embayments, open to the Tethyan Ocean to the north, throughout the Palaeozoic, and much of the Mesozoic eras. They did not gain their present closed basinal shape until the middle of the Cretaceous Period (Chapter 2). The genesis of circular sag basins, such as those of Africa, have long attracted attention. Several modes of origin have been postulated. One of the most popular proposes that thermal doming over a mantle "hot spot" leads to the erosion of uplifted crustal rocks, followed by cooling and crustal collapse, initially into a rift, followed by gentle sag subsidence (Allen and Allen, 1990). More recently it has been suggested that crustal sags may result from "cold spots" due to mantle cooling, resulting in downwelling and a dignified gentle sagging of the crust (Hartley and Allen, 1994). This model implies that sag basins may lack a precursor rifting phase, and an early high heat

X

SEDIMENTARY BASINS OF AFRICA - - INTRODUCTION AND ACKNOWLEDGEMENTS

flux, important considerations when modelling basins for petroleum generation studies. Support for this mechanism is provided by the preservation of the uniform Lower Palaeozoic pan-Saharan stratigraphic sequence in a series of isolated basins separated by basement ridges (Chapter 1). Palaeocurrent data (presented in Chapter 2) clearly show that the basins post-date their sediment fill, and implies that the uniform northerly slope of the palaeo-Tethys was locally disrupted by crustal sag basins in the mid-Cretaceous. The Cretaceous Period was a very important time in the history of Africa. The break up of the Gondwana continent, and the concomitant opening up of the Atlantic Ocean, defined the present boundaries of the African continent. As the Cretaceous rifts extended across Gondwanaland some rifts opened to become the Atlantic Ocean, others failed, becoming infilled with thick sequences of sediments that often included organic-rich muds laid down in restricted marine or fresh-water environments. Failed rifts are characterised by high heat flow, due to crustal thinning. These failed rift basins are thus often important petroleum provinces. Epicratonic rifts, such as the Sirte embayment of Libya (Chapter 3), have largely carbonate reservoirs, while the intracratonic rifts,

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SEDIMENTARY BASINS OF AFRICA - - INTRODUCTION AND ACKNOWLEDGEMENTS

XI

such as those of the Sudan (Chapter 6) and east Africa (Chapter 9) are characterised by terrigenous sediments. Finally, the present Atlantic and Indian Oceanic coasts of Africa are defined by rifts that failed and extended by sea-floor spreading into oceans. The Atlantic coastal basins are significant petroleum provinces, because their sediments prograded out over organic rich muds that were deposited in the narrow anoxic Cretaceous seaway of the incipient Atlantic Ocean (Chapters 8 and 13). Where the failed rift system of the Benue trough cross-cuts the Atlantic coast, a vast influx of terrigenous detritus generated the major petroleum province of the Niger delta (Chapter 7). The analogous rift basins on the Indian Ocean coast seem to lack such rich source beds (Chapter 10). Thus the geological history of Africa, as revealed by its sedimentary basins, has gone from unity to diversity, a pleasing synthesis of the disparate views of Africa held by the modem and ancient world. Figure 1 shows the location and chapter numbers of the sedimentary basins of Africa that are described in this volume. Readers will note the uneven length and varied style of the contributions in the different chapters. There is good reason for this treatment. There are still many African sedimentary basins, such as the rifts of eastern Africa, that have barely been described in the literature to date. In such instances the value of this volume is that it publishes the first detailed accounts of their geology. Other basins, however, such as those of northwest Africa, have been described over many years in polyglot papers in many journals. It is perhaps useful, therefore, to produce concise review papers that give coherent accounts of these basins, and that direct the reader towards appropriate references for sources and further information. I thank the contributors to this volume, not only for their contributions, but either for their patience, while their co-contributors wrote their contributions, or thank the laggards, for finally contributing. It has been remarked that a million years means nothing to a geologist, and that this is why they should never be lent money. But this dictum also applies to some geological authors. Finally I must acknowledge the many people and organisations who have enabled me to study Africa and provided me with the geological perspective to assemble and contribute to this volume. These include Arabian Gulf Oil Company, B.H.P., Esso, Genmin, Island Oil Corporation, Mobil, Oasis Oil Company of Libya, Shell, and SOEKOR. I am also grateful to the many geologists who have either propelled me off into Africa on some geological venture or other, or who have travelled with me across the continent, sharing not only their geological knowledge, but even their last can of beer. I am in their debt. R.C. SELLEY (Editor)

REFERENCES

Allen, EA. and Allen, J.R., 1990. Basin Analysis Principles and Applications. Blackwell Scientific, Oxford, 451 pp. Clifford, A.C., 1986. African oil -- past, present and future. In: M.T. Halbouty (Editor), Future Petroleum Provinces of the World. Am. Assoc. Pet. Geol. Mem., 40: 339-372. Hartley, R.W. and Allen, P.A., 1994. Interior cratonic basins of Africa: relation to continental break-up and role of mantle convection. Basin Res., 6:95-113. Kilian, C., 1922. Aper~u grnrrale de la structure des Tassilis des Ajjers. C.R. Acad. Sci. Paris, 175: 825-827. Playfair, K.J. and Hall, B., 1815. Account of the structure of Table Mountain, and other parts of the peninsula of the Cape. Drawn up by Professor Playfair from observations made by Captain Basil Hall (then Lieutenant) R.N., ER.S. Edinb. (Read 31 May, 1813) Trans R. Soc. Edinburgh, 7: 269-278. Pliny (the Elder), 78. Natural History, V11, 17. Tankard, A.J. and Hobday, D.K., 1977. Tide-dominated back-barrier sedimentation, Early Ordovician Cape Basin, Cape Peninsula, South Africa. Sediment. Geol., 18: 135-159.

List of Contributors *

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M.A. ALA 8 Department of Geology, Royal School of Mines Imperial College of Science, Technology and Medicine London SW7 2BP, U.K. G.K. BRINK 13 Consultant P.O. Box 226 Franschoek 7690, South Africa D.S. BROAD 13

Exploration Department, SOEKOR P.O. Box 307 Parow 7500, South Africa A.D.M. CHRISTIE 12 Geological Survey of South Africa P.O. Box X112 Pretoria 0001, South Africa D.I. COLEI2 Geological Survey of South Africa P.O. Box X112 Pretoria 0001, South Africa

K.J. HSU Institute for Resource and Environmental Geosciences Colorado School of Mines Green Center 1500 Illinois Street Golden, CO 80401, USA J.J. MAIER~3 Exploration Department, SOEKOR P.O. Box 307 Parow 7500, South Africa E.I. MBEDE l~ Department of Geology University of Dar-es-Salaam P.O. Box 35052 Dar-es-Salaam, Tanzania I.K. McMILLAN 13 De Beers Marine (Pty) Ltd P.O. Box 87 Foreshore Cape Town 8000, South Africa

A. DUALEH l~ Department of Geology Mogadishu, Somalia

R. McG. MILLER II National Petroleum Corporation of Namibia Private Bag 13196 Windhoek, Namibia

A.S. EL HAWAT4 Department of Earth Sciences Garyounis University P.O. Box 543 Benghazi, Libya

R.T.J. MOODY 5 Department of Geology Kingston University Penrhyn Road Kingston-upon-Thames, Surrey KT1 2EE, U.K.

L.E. FROSTICK 9 Research Institute for Environmental Science and Management University of Hull Hull HU6 7RX, U.K.

C.S. NWAJIDE 7 Shell Petroleum Development Company of Nigeria Ltd. XGSW/3 Warri, Nigeria

M.R. JOHNSON 12 Geological Survey of South Africa P.O. Box X112 Pretoria 0001, South Africa

S.W. PETTERS 7 Department of Geology University of Calabar Calabar, Nigeria

* Superior ciphers refer to the chapter number.

XIV T.J.A. REIJERS 7 Shell Petroleum Development Company of Nigeria Ltd. XGSW/3 Warri, Nigeria D.L. ROBERTS 12 Geological Survey of South Africa EO. Box X112 Pretoria 0001, South Africa R.B. SALAMA 6 CSIRO, Division of Water Resources Private Bag EO. Wembley, WA6014, Australia R.C. SELLEY 1'2'3'8 Department of Geology, Royal School of Mines Imperial College of Science, Technology and Medicine London, U.K.

LIST OF CONTRIBUTORS C.J. VAN VUUREN 12 Geological Survey of South Africa EO. Box X112 Pretoria 0001, South Africa J.N.J. VISSER12 Geological Survey of South Africa EO. Box X112 Pretoria 0001, South Africa H. de V. WICKENS 12 Geological Survey of South Africa EO. Box X112 Pretoria 0001, South Africa

Chapter 1

The Sedimentary Basins of Northwest Africa" Stratigraphy and Sedimentation

R.C. SELLEY

preserved within a series of gentle sag basins, the similarity of stratigraphy implies deposition on a uniform shelf which underwent subsequent warping, thus leaving the strata preserved within the basins, and eroded from the intervening arches (Fig. 1). The sedimentary wedge of the Sahara Platform is separated from the third zone, the Atlas Fold Belt, by a major tectonic break, termed the Sahara Flexure. In the Atlas Mountains the sedimentary section attains its maximum thickness, with rocks of all geological periods represented. It is believed, however, that the Atlas Fold Belt accreted onto the African Shield during late lateral crustal movement, as the

INTRODUCTION

Northwest Africa consists essentially of three main structural units. From south to north these are: the Precambrian cratons of the central Sahara, the Sahara Platform and the Atlas Fold Belt. The Precambrian basement of the central Sahara dips gently northwards towards the present day Mediterranean beneath a cover of Phanerozoic sediments. These sediments were deposited on the southern shores of the ancient Ocean of Tethys. The strata of the Sahara platform possess a remarkable uniformity of stratigraphy and facies. Though these strata now lie

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are dealt with in chapters l and 2, respectively. The Sirte basin is described in Chapter 3.

African Basins. Sedimentary Basins of the World, 3 edited by R.C. Selley (Series Editor: K.J. Hsti), pp. 3-16. 9 1997 Elsevier Science B.V., Amsterdam. All rights reserved.

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Tethyan Ocean closed into the Mediterranean sea (Clifford, 1986, Klitgord et al., 1988). Because the Atlas Mountains are newcomers, and thus not really part of mainland Africa, they are not described in the present volume. For over a century geologists have remarked on the remarkable uniformity of geology in Arabia and North Africa. As Doughty wrote, following his epic field work in Arabia: "The geology of the land of the Arabs is truly of the Arabian simplicity, a stack of plutonic rock, where upon lie sandstones, and upon the sandstones limestones. There are besides great landbreadths of lavas and spent volcanoes ... " (Doughty, 1888) This sentence expounds a fundamental truth about the geology of Arabia and North Africa, lands that together once formed the southern shores of the great Tethyan Ocean. A geotraverse southwestwards from the Gulf of Arabia, or southwards from the Mediterranean, reveals a broadly similar stratigraphy (Fig. 2). This commences with Pleistocene beach rock. Moving inland one traverses a flat sarir or dune field surface, every now and then descending a limestone escarpment. Intermittently one encounters Tertiary lava flows and volcanoes. Finally one descends the last limestone scarp, normally a Cretaceous chalk, at the base of which are some ubiquitous Cretaceous Greensands. The traverse continues across a dissected terrain of sandstone scarps, jebels and wadis, before finally reaching Precambrian metamorphic basement. Detailed examination of the stratigraphy reveals great similarities, particularly in the Palaeozoic strata. Because of the uniformity of stratigraphy it is sensible to deal with this aspect in a single chapter to encompass all the basins of north west Africa. The next chapter then describes the structural evolution of the main sedimentary basins, and notes local variations in stratigraphy. This organisation avoids the unnecessary repetition of stratigraphic detail. The only exception to this arrangement is the Sirte basin of Libya, whose stratigraphy and structural history is

markedly different from the rest. It is thus dealt with in a separate chapter.

PRECAMBRIAN BASEMENT AND INFRA-CAMBRIAN SEDIMENTS

It is beyond the scope of this volume to discuss the Precambrian basement on which the Phanerozoic sedimentary basins lie. Detailed accounts will be found elsewhere (Fullagar, 1980; Ghuma and Rogers, 1980; Bowen and Jux, 1987). Briefly the Precambrian basement is composed of a complex series of igneous and metamorphic rocks that reflect the gradual cratonisation of continental crust, culminating in a widespread tectono-thermal event dated at about 580 my BP (Fig. 3). This basement includes several sedimentary sequences that have undergone regional metamorphism in general proportion to their age. The identification of the Precambrian: Cambrian boundary thus presents something of a problem. Across much of the Sahara the Precambrian basement is unconformably overlain by a thick broadly conformable sequence of sediments whose upper part contains Lower Palaeozoic fossils. The unconformity at the base of this sequence is termed the Tassilian discordance (Killian, 1922). Where the base of the Palaeozoic sequence unconformably overlies igneous and metamorphic basement the boundary between the Precambrian and the Palaeozoic rocks is easily delineated. Locally, however, barren sedimentary rocks lie beneath the major sub-Palaeozoic unconformity and above igneous and/or metamorphic basement. Such sedimentary formations are given local lithostratigraphic names and have been referred to as "Infra-Cambrian". Subsequently, however, some of the "Infra-Cambrian" sedimentary formations have been found to be of Cambrian age, on the basis of radiometric dates from associated igneous rocks. Thus, though the Tassilian discordance extends across much of the Sahara, it ranges widely in age, and can no longer be assumed to mark the Proterozoic-Palaeozoic boundary (Legrand, 1985).

THE SEDIMENTARY BASINS OF NORTHWEST AFRICA: STRATIGRAPHY AND SEDIMENTATION

5

Fig. 3. Outcrop of Precambrian granite basement, Tibesti, southern Libya. Note fractures and characteristic pan-African elephant buttock weathering

Fig. 4. Generalised Palaeozoic stratigraphic column for the Sahara platform (from Selley, 1996). This sequence, with only minor local variations, can be found preserved in all the sedimentary basins of northwest Africa from the Tindouf basin in the west, to the Kufra basin in the east. It can also be recognised with little change, apart from different formation names, on the Arabian platform.

CAMBRO-ORDOVICIAN

As noted earlier, geologists have long remarked on the similarity of Lower Palaeozoic stratigraphy in Arabia and the Sahara. There is a remarkably uniform sequence of facies that may be traced from basin to basin, from the Atlantic Ocean to the Arabian Gulf (Fig. 4). The Tassilian discordance is normally overlain by coarse pebbly channelled sands. These pass up into better sorted finer sands,

commonly bioturbated. This facies passes up in to graptolitic shales, which in turn pass up, often via turbidite sands in to prograding deltaic, and finally fluvial sands (Bennacef et al., 1971; Clark-Lowes and Ward, 1991). This is a gross generalisation, but a very useful one to memorise. There are two complications. First, though this sequence is broadly correct, there are gradations between the sand facies. Furthermore the various facies are occasionally interbedded, notably

6

R.C. SELLEY

the sandstones, but sometimes there are tongues of shale within sandstone sequences. The second problem is that of age. The sandstones contain an interesting assemblage of trace fossils, but are largely devoid of body fossils that can be used biostratigraphically, The shales, however, contain graptolites and microfossils, such as acritarchs, that can be used to establish a biostratigraphy. Because the barren sandstones are commonest at the base of the sequence, and the shales become more abundant upwards, so the stratigraphy gradually becomes better refined up the sequence. As mentioned earlier, the base of the Cambrian can only be roughly related to the Tassilian discordance. Sediments can only be attributed to a Cambrian age by inference from radiometric dates from the Precambrian basement beneath the Tassilian discordance, and from the occurrence of Ordovician fossils in overlying shales. When traced from basin to basin the rock units have been given local formation names. When actually seen in the field, it at once becomes obvious that these are of the same facies, though their age and lateral relationships may be unclear. The three facies will now be described, and their environments deduced. The lowest facies is composed of coarse pebbly channel sands (Fig. 5). Formations of this facies have essentially sheet geometries with little regional thickness variation. The base of the sequence is commonly marked by the Tassilian discordance. The unconformity is a mature pediment surface with occasional residual inselbergs. The overlying sediments range from pebbles to silt, but are largely of

gra

silt

/

\ \

f f f

Fig. 5. Representative sedimentological log of Cambro-Ordovician cross-bedded pebbly channel sand facies. Interpreted as braided alluvial outwash, composed almost exclusively of active braided channel sands, with rare abandoned channel shales.

medium to very coarse sand grade. Intraformational conglomerates of reworked siltstone occur throughout the sequence. Petrographically these sands include arkoses, especially at the base of the sequence. But they often pass up in to more mature quartz arenites. They are normally red coloured, though many formations are bleached during subsequent meteoric flushing, but the early red colour is retained by impermeable claystones and siltstones. This facies is composed of a series of superimposed channel complexes. Three main types may be recognized. Two types are broad and shallow, either with heterogeneously infilled sandstones, or shale infilled. The third type of channel has steep sided walls, and contains a sequence that fines up from a basal intraformational conglomerate, via sand to shale. Channels of the first type comprise about 90% of the facies. They are typically some 300 m wide and 5 m deep and infilled by sandstones with no apparent regular vertical arrangement of grainsize and structure. The channels are floored by a thin conglomeration of pebbles, and are infilled by various types of cross-bedding. Cross-bedding dips are unimodal at any one locality. When plotted regionally they indicate palaeocurrents flowing northwards off the Sahara shield. Flat bedding is also present, sometimes with scattered pebbles. Quicksand deformation structures are common, including both recumbent foresets as well as convolute bedding. Channels of the second type are rare. They are similar in scale and profile to the sand-infilled channels, internally they are quite distinct. The channel floors are marked by an extraformational pebble lag, abruptly overlain by siltstone, whose laminae drape over the pebbles. The fill is almost entirely composed of laminated micaceous siltstone. These shale-infilled channels are overlain by the more common sand-infilled channels. Body fossils are normally absent from this facies. Obscure unidentifiable trails and tracks occur. Bilobate trails attributable to Cruziana are sometimes found in the shale-filled channels (Seilacher, 1991). From the preceding account it is clear that this facies was deposited in channels. The coarse texture and cross-bedding indicates sedimentation from the bed load of extremely powerful currents. The small variability of cross-bed orientations and channel trends shows that the currents were unidirectional and regionally persistent. The siltstones clearly originated from the infilling of abandoned channels. The predominant red colour suggests early diagenesis above the water table. There are no unequivocal marine fossils. This facies would thus appear have been deposited from nonmarine channel-confined unidirectional traction currents. These conditions are to be found in braided alluvial channels such as occur

THE SEDIMENTARY BASINS OF NORTHWEST AFRICA: STRATIGRAPHY AND SEDIMENTATION

7

Fig. 6. Large scale (>2 m) tabular planar cross-bedding in Tigillitiferous Ordovician shallow marine shoal sand. Jebel Eghei, southern Libya.

on alluvial fans, or on extensive braid plains formed from coalesced alluvial fan systems. Palaeomagnetic global reconstructions show that the western Sahara was close to the South Pole in the Early Palaeozoic, and there is abundant evidence of glaciation (to be presented later) in overlying late OrdovicianSilurian sediments. It can be argued, therefore, that these braided alluvial sands and gravels were deposited as glacial outwash on the flanks of polar ice caps The braided alluvial pebbly sand facies is overlain by, and sometimes interbedded with, the non-pebbly sheet sand facies. The non-pebbly sheet sand facies is overlain by, and sometimes interbedded with graptolitic shales. The non-pebbly sheet sand facies is composed largely of well-sorted medium and fine grained proto-quartzites. This facies is generally cross-bedded, with tabular planar cosets up to 3 m high with sets 5-15 cm high. Sometimes, however, individual sets up 2 m height occur (Fig. 6). Troughs are generally rare. Foresets are homogeneous and accretionary, reflecting the good sorting of this facies. Heterogeneous and avalanche foresets, such as occur in the poorly sorted pebbly sands beneath, are rare (Fig. 7). Cross-bed dip directions in the sandstones are generally unimodal, and indicate deposition from northerly currents flowing down the depositional slope from the Sahara shield towards the Tethyan Ocean. There are very few body fossils in these sandstones. There are, however, occasional trace fossils. The most characteristic type is the vertical burrow known throughout the Sahara as Tigillites, and elsewhere in the world as Sabellarifex, Scolithos, or

Fig. 7. Representative sedimentological log of Cambro-Ordovician Tigillitiferous nonpebbly sheet sand facies. Interpreted as shallow marine shelf environment, composed almost exclusively cross-bedded shoal sands, with shale sheets deposited on tidal flats.

Monocraterion. These burrows are sometimes so abundant that they destroy any original sedimentary structures that may have once existed. The sandstones are interbedded with rare grey argillaceous micaceous siltstones with thin very fine sandstone layers. These units are each between 1-3 m thick and have sheet geometries, in contrast to the abandoned channel silts of the facies beneath. The

8 siltstone sheets are generally laminated throughout with occasional thin beds of very fine sand and isolated sand ripples. These sandstones are rippled throughout. Micro-cross-laminated cosets are generally absent; the sands being composed of congeries of tippled lenses separated by argillaceous laminae and clay drapes. The bases of the siltstone sheets are generally transitional, their tops are abrupt, rarely erosional. These shale units contain a diverse trace fossil assemblage that includes Cruziana, Tigillites and Harlania. The abundance and orientation of cross-bedding in this facies points to deposition from unidirectional lower flow regime traction currents. The fine grainsize, however, shows these currents had significantly lower velocities than those which deposited the braided alluvial sands beneath. The predominance of tabular planar cross-beds, and the absence of channelling indicates that these were open-flow currents unconfined by channel banks. The vertical burrows characterize the Scolithos ichnofacies which is diagnostic of shallow marine conditions. The laminated shale units indicate sporadic lower energy conditions when suspended sediment settled out. The associated rippled very fine sands show that gentle traction currents sometimes occurred; while the absence of cross-laminated cosets and the presence of clay drapes on ripple crests suggests that these currents pulsated gently. Such conditions are more likely to be found in tidal realms rather than in the more regular flows of river channels. A shallow tidal environment is also suggested by the suite of trace fossils. Tigillites, Cruziana and Harlania are all characteristic of shallow marine deposits. The weight of the evidence suggests therefore that the Cambro Ordovician non-pebbly sheet sand facies of the Sahara and Arabia originated in a marine shelf environment. The cross-bedded sands were probably deposited from migrating megaripples and bars similar to those described from modern shelf sand waves. The thin shales and very fine sands suggest intermittent regressive phases when shallower tidal fiat deposits were laid down. Across Arabia and the Sahara the non-pebbly sheet sand facies normally overlies the pebbly channel sand facies. Locally, however, interbedded formations of the two facies types occur, and in some instances the two facies are interbedded. Sometimes, for example, braided channels are found capped by Tigillitiferous horizons. The later sequences indicate the deposits of the actual coastline, where braided channels reached the sea and were subjected to marine influences. Study of the boundaries between sequences of the 2 facies reveals the nature of transgressions and regressions acros the Arabian and Saharan shields. The sequence boundaries where the fluvial forma-

R.C. SELLEY tions overly the shallow marine sands are usually planar erosional surfaces, marked by a thin basal conglomerate, that are intermittantly dissected earlier by steep-sided channels. They are thus analogous to the sequence boundaries described earlier from within the pebbly sand facies. They thus imply a similar origin, namely a pediment retreating across a lithified substrate, succeeded by braided alluvial outwash deposits. Where sequences of pebbly channel sands are overlain by the non-pebbly sheet sand facies, by contrast, the sequence boundary though also planar, is unchannelled. This implies that the sea transgressed across a wave cut bench that was then overlain by the non-pebbly sheet sand facies. Figure 8 displays the relationship between sequence boundaries and fluctuating sea level envisaged for interbedded formations of the Cambro-Ordovician pebbly channel sand and non-pebbly sheet sand facies. Note that these are superimposed over a gradual marine transgression across the Saharan and Arabian shields.

SILURIAN The non-pebbly shallow marine sheet sands are overlain by and pass down northwards into black graptolitic shales. Over most of southern Algeria and Libya there is a major discordance between the shallow marine sands and the overlying black shales. This transgresses older formations to directly overly the Precambrian basement in the centre of the Sahara. The main tongue of this shale is termed

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FACIES Non-pebbly sheet sand Pebbly channel sand facies PreCambrian basement

ENVIRONMENT Shallow marine shelf Braided alluvial outwash

SEA LEVEL FALLING RISING

(a)

1. Falling 2. Slowly rising 3. Rising fast 4. Slowly rising 5. Slowly falling 6. Falling fast 7. Slowly falling

Fig. 8. Geophantasmograms to illustrate the relationship between Lower Palaeozoic facies, sequence boundaries and inferred sea level change in north Africa and Arabia. In general terms braided alluvial sands are overlain by and pass Tethys-ward into Tigillitiferous marine shoal sands, with occasional interfingering. Fluvial sands typically overly channelized (Type 1) sequence boundaries, Tigillitiferous shoal sands overly planar sequence boundaries.

THE SEDIMENTARY BASINS OF NORTHWEST AFRICA: STRATIGRAPHY AND SEDIMENTATION ment shield

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Fig. 9. The Silurian Tannezuft Shale-Acacus Sandstone transitional sequence, Jebel Eghei, southern Libya. Though diachronous when traced off the shield towards the palaeo-Tethys, this upward-coarsening progradationai sequence can be traced along the strike from Mauritania, across north Africa and into Arabia. the Tannezuft Formation, from Wadi Tannezuft in Algeria. It can be traced from basin to basin across much of the Sahara, being variously referred to as the Tannezuft or Gotlandian (French for Silurian) shale. It is, as this name suggests, largely of Silurian age. In many parts of the Sahara geomorphic features indicative of glacial erosion occur cut into exhumed surfaces of divers pre-Silurian rocks. Glacial features such as striated pavements and roche moutonn~es have been recorded from the Atlantic coast to the southeastern part of the Kufra basin (Beuf et al., 1969). In some localities these surfaces can be traced beneath the base of the Tannezuft shale. Furthermore, thin pebbly mudstones and debrites, attributable to glacial moraines, occur between the graptolitic shales and the underlying discordant the base of the Tannezuft Formation. Locally diamictites infill steep sided valleys cut into the sub-Tannezuft discordance. Elsewhere diamictites exhibit the sinuous form of exhumed eskers. Because the base of the Tannezuft shale contains Lower Llandoverian graptolites, these erosional and depositional features demonstrate a pre-Early Llandoverian glaciation for much of North Africa. Analogous broadly coeval glacial features have been extensively documented in Arabia (e.g., Abed et al., 1993). The Tannezuft shale is often rich in organic matter. It is an important petroleum source rock for petroleum in the Hassi Messaoud and other fields of Algeria (Tissot et al., 1975). The black shales pass up into a sequence of upward-coarsening shale to sand increments in which the sand percentage grad-

ually increases upward (Fig. 9). The upper sand is termed the Acacus Sandstone throughout the Sahara. The type section being Wadi Acacus on the western flank of the Murzuk basin. Regional biostratigraphic studies by Klitzsch (1965) and Bellini and Massa (1980) using graptolites show that the shale-sand transition is strongly diachronous down the depositional slope tYom the Sahara shield towards the Tethys (Fig. 10). The transition from the Tannezuft shales into the Acacus sands is often marked by interbedded sequences of shales and thin sands with erosional bases, graded bedding, and fragmentary Bouma sequences. These are the characteristic features of turbidites. The Tannezuft Shale grades up transitionally into the Acacus Sandstone Formation (Fig. 11). This is a medium to very fine grained clean well sorted sandstone. As sand content increases up the section, trace fossils become common. They include Cruziana, Harlania and sparse Tigillites. These sands are overlain by cross-bedded channel sands with shale pellet basal conglomerates. The Tannezuft-Acacus boundary is often composed of several upward-coarsening genetic increments with a general pattern strongly suggestive of deltaic sedimentation, viz.: Unit

Facies

Environment

I II

Cross-bedded channel sands Interlaminated,load-casted, bioturbated cross-laminated sands and shales Graptoliticlaminated black shales with occasional turbidites

Distributarychannels Mouth bar and delta slope

III

Pro-delta offshore mud zone

THE SEDIMENTARY BASINS OF NORTHWEST AFRICA: STRATIGRAPHY AND SEDIMENTATION "', Tunisia North

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Fig. 10. Cross-section to show the diachronous nature of the Tannezuft-Acacus boundary across the Saharan platform. This reflects the gradual progradation of the Acacus delta system over the deeper water muds of the Tannezuft Formation. (Based on Bellini and Massa, 1980.)

DEVONIAN

Acacus Formation

cross-bedded multi-storey channel sands

assorted trails & burrows

upward coarsening increment >lOOm

graded sands with fragmentary Bouma sequences

Tannezuft Formation

graptol ites

Fig. 11. Sedimentological log through the Tannezuft-Acacus boundary showing the transitional nature. Interpreted as offshore muds overlain by a prograding delta slope and alluvial flood plain.

The repetition of these increments may be attributed to global or tectonic causes, or to the autocyclic progradation and abandonment of successive delta lobes as the shoreline gradually regressed across the Sahara Platform towards the Tethyan Ocean.

The remarkable similarity of Saharan Lower Palaeozoic facies and stratigraphy begins to break down in the Upper Palaeozoic. In the north and west the Silurian Acacus delta is conformably overlain by Devonian sands, shales and limestones with a diverse marine fauna. This suggests a gradual deepening of the sea. To the south and east, however, the base of the Devonian is strongly discordant. Early Devonian sediments are absent, and Emsian and younger sediments unconformably overstep older formations down to the Cambro-Ordovician fluvial sands. In these areas the unconformity is overlain by the sandstones that are attributed to the Tadrart Formation. The Tadrart Formation is a coarse to medium grained well sorted sand with a sparse kaolin matrix. It contains rare fine sand and thin shale horizons. Cross-bedding occurs in a variety of forms which are often seen to infill channels. Much less frequently finer, flatbedded sands with Tigillites occur. Plant fragments (Lepidodendron sp.) and spores have also been found. These have been used to place the Tadrart Formation in the Siegenien and Emsian stages. The Tadrart Sandstone is often broadly comparable in facies to the Cambro-Ordovician rocks. A similar environment is proposed, of braided alluvial sedimentation with occasional marine incursions that reworked the sands on beaches, bars and shoals (Turner, 1980). The quartzose composition, finer grain size, lack of pebbles and good sorting of the sands of the Tadrart Formation indicate a polycyclic derivation from Lower Palaeozoic strata cropping out on the shield margin, rather than a direct derivation from the basement hinterland. The Tadrart Sandstone is overlain conformably by the Ouan Casa Formation. Klitzsch (1969) has published the type section from the Wadi Ouan Casa in southwest Libya. The Ouan Casa Formation consists essentially of laminated siltstone with, occasionally, thin laterally

12

R.C. SELLEY g,'aveJ

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Libya in the east, a distance of some 3,000 km. Two types of ironstone can be recognised according to their distribution. One type is thin, and laterally extensive, cropping out for distances of 200 to 300 km. The second type is stratigraphically restricted to the Lower and Upper Devonian. It occurs in localised lenses some 20-30 m thick. It is this second type that is economically important (Guerrak, 1987).

.

.

CARBONIFEROUS-PERMIAN

.

-....__

1-2 m

Fig. 12. Sedimentological log of typical Devonian upward-coarsening genetic increment. Interpreted as upward-shoaling offshore mud to coastal sand bank environment, capped by a significant bioturbated pause horizon (maximum flooding surface).

extensive laminae of very fine sand. It contains horizons of Tigillites and rare brachiopods. The latter are taken by Klitzsch (1969), Collomb (1962) and Burollet and Manderscheid (1967) to indicate an age from Emsian to Eifelian. The Uan Casa Formation is disconformably succeeded by the Aouinet Ouenine Formation over much of the Sahara. There is no marked angular contact, and no overstepping relationship. The type locality is at the well of Aouinet Ouenine on the western flank of the Gargaf arch. Collomb (1962) describes five laterally extensive upward-coarsening shale to sand cycles each about 30 metres thick. When studied in detail these are often composed of smaller upward-coarsening genetic increments (Fig. 12). Comparable sedimentology and faunas imply that the Aouinet Ouenine and Uan Casa formations share a similar depositional environment. The palaeontology shows that this was shallow marine to intertidal. The sedimentology points to deposition in a series of tidal flats and sand banks that prograded cyclically down slope over deeper water offshore muds. There is, however, one other distinctive facies that reaches its apogee in the Devonian. Ironstones occur intermittently throughout much of the northwest Sahara in rocks of Ordovician to Lower Carboniferous age. This facies is particularly well-developed in the Devonian. The ironstones are generally oolitic, and occur interbedded with the shallow marine sands and shales previously described. They crop out intermittently from Zemmour in the west to the Fezzan of

The type sections for Saharan Carboniferous rocks are within the Illizi-Gahadames basin where they attain a thickness of over 500 m. The contact with the underlying Devonian is disconformable. The Carboniferous sediments are broadly comparable to those of the Devonian in facies and distribution. Marine limestones and shales in the northwestern Sahara and Atlas, pass south and east into progressively more sandy and continental deposits. The limestones include oolitic, bioclastic and algal Collenia varieties. These are interbedded with, and pass southwards and eastwards into, upward-shoaling shale: sand sequences analogous to those of the Devonian. These in turn pass south and eastwards into coarse, cross-bedded and channelled sandstones, with Lepidodendron and other unidentifiable plant fragments. These are comparable in facies to the Tadrart and Lower Palaeozoic sands which were earlier identified as of braided alluvial outwash origin. Late Carboniferous evaporites occur in the Reggane and Bechar basins of Algeria, heralding the onset of the Triassic phase of aridity. Preserved Saharan Carboniferous rocks generally range in age from Tournasian-Westphalian, though the evaporites in the Bechar basin are dated as late as Stephanian. The generally barren nature of the continental sands and evaporites means that they may extend in age into the Permian. By the end of this phase of deposition, however, the sea retreated northwards, never ever to return to the Saharan Platform with the depth and extent that prevailed through much of the Palaeozoic Era. This retreat of the sea was followed by a major warping of the Palaeozoic strata, sometimes referred to as the Hercynian Orogeny. Where preserved, the Mesozoic sediments overlie Palaeozoic strata with a strong regional unconformity.

MESOZOIC

Two main facies were deposited during much of the Mesozoic Era. A thick section of marine Mesozoic sediments crops out in the complex folds of the Atlas Mountains. These can be traced southwards in

THE SEDIMENTARY BASINS OF NORTHWEST AFRICA: STRATIGRAPHY AND SEDIMENTATION the subsurface in the Algerian sedimentary basins. Over much of the Sahara, however, the Palaeozoic strata are unconformably overlain by nonmarine, largely, unfossiliferous, sandstones, variously termed "Continental Post-Tassilian" and "Continental Intercalaire" in Algeria, "Continental Mesozoic" in Libya, and "Nubian", in Sudan, Egypt and Arabia. The dateable marine Mesozoic of the northwest will first be described, followed by an attempt to produce a coherent account of the nonmarine Mesozoic sediments. In the Atlas Mountains, and in the subsurface of the Algerian basins, the Mesozoic sequence begins with a series of evaporites. The Stephanian evaporites of the Bechar basin have been already noted. Evaporites are notoriously difficult to date palaeontologically, so it is quite possible that evaporite deposition commenced in the Atlas area in the Permian Period or earlier. The evaporites include both anhydrite and halite, and are interbedded with dolomites, red shales and occasional sandstones. The evaporites provide a regional barrier to petroleum migration across the Algerian basins, acting as the seal to the Hassi Messaoud oil field. Basal sands (dated as Triassic) between the Hercynian unconformity and the evaporites serve as petroleum reservoirs in the Hassi er R'Mel and other Algerian fields. In the Atlas Mountains the Triassic sediments are overlain by Lower Jurassic dolomites and limestones, heralding a return to the open marine conditions that continued throughout much of the Mesozoic. The marine transgression advanced diachronously across the northwestern part of the Sahara shield in Algeria, Tunisia and northern Libya.

13

Arid conditions continued intermittently, with Early Jurassic and Senonian evaporites occurring in the Polignac and Timimoun basins of Algeria. For the most part, however, shallow marine limestones, shales and sands were deposited throughout the Jurassic in all these areas. Over much of the Sahara the dateable (to some extent) Palaeozoic strata are unconformably overlain by continental deposits that are largely barren of fossils. These are in turn locally overlain by fossiliferous Cretaceous limestones, commonly of Cenomanian age. As mentioned earlier, these barren sandstones were given a range of local names. They continue to be a stratigrapher's nightmare (e.g., Banerjee, 1980, and Klitzsch and Squyres, 1990). The sandstones are largely unfossiliferous, apart from wood fragments attributable to the genus Dadoxylon (Fig. 13), and leaf imprints attributable to the genus Cladophlebis. These plant remains are taken to indicate an Early Cretaceous (Neocomian) age. Sandstones of this type were first described from the Sudan by Russeger (1837), who gave them the appropriate name of the "Nubian Sandstone", and attributed them to an Early Cretaceous age, because they were overlain by fossiliferous, and hence dateable, Cretaceous limestones. Thereafter, for a century or more, the term "Nubian" was given to barren sandstones beneath dateable Cretaceous limestones and above basement, from Libya in the west, to Arabia in the east. Gradually, as research progressed, and fossils were discovered, it became apparent that the "Nubian", though largely continental in origin, and commonly a very distinctive

Fig. 13. Fossilized wood (Dadoxylonsp.) in the Continental Mesozoic sandstones of the Messak scarp, northern Murzuk basin.

1

4

R

.

C

.

SELLEY

Fig. 14. Coset of tabular planar cross-bedding in braided alluvial Continental Mesozoic sandstones of the Messak scarp, northern Murzuk basin.

facies of braided alluvial origin, actually ranged in age from Early Cretaceous back to Cambrian or earlier (Van Houten, 1980; Klitzsch and Squyres, 1990). The term "Nubian" became abandoned, and indeed, in the late nineteen-sixties more papers published about Nubian semantics, than about the rocks themselves. Nowadays if researchers are in doubt about the stratigraphical affinity of a barren Saharan sandstone they .provide it with a new formation name. Sedimentologically the "Continental Mesozoic" (for want of a better term) consists of a range of facies indicating deposition in a corresponding range of different depositional environments. The facies include conglomerates that formed as fans around the basin margins. The fanglomerates are overlain by, and pass basinwards into, thick sequences of cross-bedded pebbly channelled sands (Fig. 14), comparable to the braided alluvial deposits described in earlier formations (Fig. 15A). There is also a facies of upward-fining increments of conglomerate-sandstone-red shale (Fig. 15B). Superficially this suggests deposition in a meandering alluvial environment. With the sands being deposited as channel point bars, and the shales as flood plain

deposits. Sometimes, however, shale laminae on the channel floors exhibit desiccation cracks, indicating that discharge was ephemeral (Fig. 16). The shales are commonly massive, with a conchoidal fracture, and they contain rare scattered grains of well-rounded quartz and feldspar. This distinctive texture is characteristic of playa lakes, across which isolated sand grains blow, as in parts of the present day Sahara. Thus it is more probable that these upward-fining increments are more probably indicative of alternations of braid plain and playa lake, than of meandering alluvial flood plain conditions. There are also occasional thick units of red and variegated mudstones, with rare scattered well-rounded sand grains that are diagnostic of a playa lake environment. In the Kufra basin typical continental Mesozoic sandstones formations overly and underly a curious silicified limestone, termed the Chieun Formation. This locally contains abundant Hydrobia, indicating a lacustrine environment, and a post-Triassic age. A description of the dateable marine limestones will be given in the chapter on the Sirte basin, since this is where they are best developed. It should be noted, however, that in some parts of the Sahara

T H E S E D I M E N T A R Y BASINS OF N O R T H W E S T AFRICA: S T R A T I G R A P H Y A N D S E D I M E N T A T I O N

15

gravel llJlJllsandi siltl clay

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.

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Fig. 15. Representative sedimentological logs of two of the most common facies of the Continental Mesozoic. (A) Cross-bedded pebbly channel sand facies. Interpreted as braided alluvial outwash, composed almost exclusively of active braided channel sands, with rare abandoned channel shales. (B) Upward-fining genetic increments of coarse cross-bedded channel sands overlain by shales9 Superficially these deposits are comparable to the point bars and flood plains of humid meandering fluvial channel systems. Note, however, the desiccation cracks on channel floors, indicative of ephemeral flow, and the scattered rounded (wind blown?) sand grains in the massive siltstones. Interpretable as ephemeral braided channels passing out into playa lakes.

there are isolated jebels of chalk, with a thin basal greensand, unconformably overlying a broad range of stratigraphic units, right down to Precambrian basement. These outliers are crucial markers of the break up of Gondwanaland, serving to delineate the extent of narrow Cretaceous seaways that extended right across the Sahara shield and into what is now N i g e r i a ( K o g b e , 1980).

Fig. 16. Desiccation cracks in abandoned braided channel shale within the Continental Mesozoic sandstones of the Messak scarp, northern Murzuk basin9

16

SELECTED BIBLIOGRAPHY Naturally an area as large as north-west Africa, with its great mineral wealth, has attracted a large literature. Important source books include: Holland, C.H., 1985. Lower Palaeozoic of North-Western and West-Central Africa. Wiley, Chichester, 512 pp. Salem, M.J. and Busrewil, M.T. (Editors), 1980. Geology of Libya, Vols. I-III. Academic Press, London, pp. 1-1712. Salem, M.J. and Belaid, M.N. (Editors), 1991. Geology of Libya, Vols. IV-V. Elsevier, Amsterdam, pp. 1713-2095. Salem, M.J., Sbeta, A.M. and Bakbak, M.R. (Editors), 1991. Geology of Libya, Vol. VI. Elsevier, Amsterdam, pp. 20992491. SONATRACH, 1979. Geology of Algeria. In: J.L. Chardac (Editor), Well Evaluation Conference 1979. Schlumberger, Paris, pp. 1-26.

REFERENCES Abed, A.M., Makhlouf, I.M., Amireh, B.S. and Khalil, B., 1993. Upper Ordovician deposits in southern Jordan. Episodes, 16: 316-328. Banerjee, S., 1980. Stratigraphic Lexicon of Libya. Socialist People's Libyan Arab Jamahiriyah Industrial Research Centre, Tripoli, 300 pp. Bellini, E. and Massa, D., 1980. A Stratigraphic contribution to the Palaeozoic of the Southern Basins of Libya. In: M.J. Salem and M.T. Busrewil (Editors), Geology of Libya, Vol. I. Academic Press, London, pp. 3-56. Bennacef, A., Beuf, S., Biju-Duval, B., De Charpal, O., Gariel, O. and P. Rognon., 1971. Example of cratonic sedimentation: Lower Palaeozoic of Algerian Sahara. Bull. Amer. Assoc. Petrol. Geol., 55: 2225-2245. Beuf, S., Biju-Duval, B., Stevaux, J. and Kulbicki, G., 1969. Extent of Silurian Glaciation in the Sahara: its influences and consequences on sedimentation. In: W.H. Kanes (Editor), Geology, Archaeology and Prehistory of Southwestern Fezzan, Libya. Petrol. Explor. Soc. Libya, Tripoli, pp. 103-116. Bowen, R. and Jux, U., 1987. Afro-Arabian Geology. Chapman and Hall, London, 295 pp. Burollet, P.F. and Manderscheid, G., 1967. Le Devonian en Libye et en Tunisie. Int. Symp. Devonian System, Calgary, Vol. 1, pp. 205-213. Clark-Lowes, D.D. and Ward, J., 1991. Palaeoenvironmental evidence from the Palaeozoic "Nubian Sandstones" of the Sahara. In: M.J. Salem, A.M. Sbeta and M.R. Bakbak (Editors), Geology of Libya, Vol. VI. Elsevier, Amsterdam, pp. 2099-2154. Clifford, A.C., 1986. African oil m past, present, and future. In: M.T. Halbouty (Editor), Future Petroleum Provinces of the World. Mem. Amer. Assoc. Petrol. Geol., 40: 339-372. Collomb, G.R., 1962. Etude g6ologique du Jebel Fezzanet de sa bordure Palaeozoique. Notes et Mem. Comp. Fr. Petrol, No. 1, 35 pp. Fullagar, P.D., 1980. Pan-African age granites of northeastern Africa: New or reworked sialic material? In: M.J. Salem and M.T. Busrewil (Editors), Geology of Libya, Vol. III. Academic Press, London, pp. 1051-1058. Ghuma, M.A. and Rogers, J.J.W., 1980. Pan-African evolution

R.C. S E L L E Y in Jamahiriya and north Africa. In: M.J. Salem and M.T. Busrewil (Editors), Geology of Libya, Vol. III. Academic Press, London, pp. 1059-1064. Guerrak, S., 1987. Palaeozoic oolitic ironstones of the Algerian Sahara: a review. J. African Earth Sci., 6: 1-8. Kilian, C., 1922. Aper~u g6nerale de la structure des Tassilis des Ajjers. C.R. Acad. Sci. Paris, 175: 825-827. Klitgord, K.D., Hutchinson, D.R. and Schoton, H., 1988. U.S. Atlantic continental margin; structural and tectonic framework. In: R.E. Sheridan and J.H. Grow (Editors), The Geology of North America, the Atlantic Continental Margin, U.S. Decade of North American Geology Series. Geological Society of America, Washington, DC, Vol. 1-2, pp. 19-55. Klitzsch, E., 1965. Die Gotlandien-Transgression in der Zentrul Sahara. Z. Deusch. Geol. Ges. Hannover, 117: 492-501. Klitzsch, E., 1969. Stratigraphic section from the type areas of the Silurian and Devonian strata at western Murzuk basin, Libya. In: W.H. Kanes (Editor), Geology, Archaeology, and Prehistory of Southwestern Fezzan. Petrol Explor. Soc. Libya., pp. 83-90. Klitzsch, E. and Semtner, E., 1993. Silurian palaeogeography of NE Africa and Arabia m an updated interpretation. In: U. Thornweihe and H. Schandelmeier (Editors), Geoscientific Research in Northeast Africa. Springer-Verlag, Berlin, pp. 341-344. Klitzsch, E. and Squyres, C.H., 1990. Paleozoic and Mesozoic geological history of northeastern Africa based upon new interpretation of Nubian Strata. Bull. Amer. Assoc. Petrol. Geol., 74:1203-1211. Kogbe, C.A., 1980. The Trans-Saharan Seaway during the Cretaceous. In: M.J. Salem and M.T. Busrewil (Editors), The Geology of Libya, Vol. I. Academic Press, London, pp. 91-96. Legrand, P.H., 1985. Lower Palaeozoic rocks of Algeria. In: C.H. Holland (Editor), Lower Palaeozoic of North-Western and West Central Africa. J. Wiley, Chichester, pp. 5-90. Russeger, J. 1837. Kreide und Sandstein: Einfluss von Grait auf letzteren. N. Jb. Min, pp. 665-669. Seilacher, A., 1991. An updated Cruziana stratigraphy of Gondwanan Palaeozoic sandstones. In: M.J. Salem, M.T. Busrewil and A.M. Ben Ashour (Editors), The Geology of Libya, Vol. IV. Elsevier, Amsterdam, pp. 1565-1582. Seilacher, A., 1993. Problems of correlation in the Nubian Sandstone Facies. In: U. Thornweihe and H. Schandelmeier (Editors), Geoscientific Research in Northeast Africa. SpringerVerlag, Berlin, pp. 329-340. Selley, R.C., 1996. Ancient Sedimentary Environments and Their Subsurface Diagnosis. Chapman and Hall, London, 4th ed., 300 pp. Tissot, B., Deroo, G. and Espitali6, J., 1975. Etude compar6e de l'6poque de formation et d'expulsion du petrole dans diverses provinces g6ologique. Proc. 9th World Petrol. Cong. Tokyo, Applied Sci. Pubs, London, 2: 159-169. Turner, B.R., 1980. Palaeozoic sedimentology in the Southeastern part of the AI Kufrah Basin, Libya: A model for oil exploration. In: M.J. Salem and M.T. Busrewil (Editors), Geology of Libya, Vol. I. Elsevier, Amsterdam, pp. 351-374. Van Houten, F.B., 1980. Latest Jurassic-Early Cretaceous regressive facies, Northeast Africa Craton. Bull. Am. Assoc. Petrol. Geol., 64: 857-867. Whiteman, A.J., 1971. Cambro-Ordovician rocks of A1-Jazair (Algeria). A Review. Bull. Am. Assoc. Petrol. Geol., 55: 1295-1335.

Chapter 2

The Basins of Northwest Africa" Structural Evolution

R.C. SELLEY

INTRODUCTION

As described in the previous chapter, the sedimentary basins of Northwest Africa share a remarkably uniform Early Palaeozoic stratigraphy, that gradually became more varied regionally until the great Late Carboniferous marine regression. It was because of this uniformity that their stratigraphy was described, in general terms, in a single tidy chapter, to avoid unnecessary repetition. In this chapter the structure of the various basins will be described one by one, in a general west to east direction. Only the Sirte basin will be omitted from this chapter and dealt with in one of its own. This is because it is younger than the other basins, formed in a different way, and has a different stratigraphy and structural history. Considered at is simplest north west Africa consists of the Sahara shield. This is a vast area of Precambrian continental crust that has undergone little structural deformation since the end of the Proterozoic Era. When traced northwards towards the present day Mediterranean the Sahara shield dips gently beneath the cover of Phanerozoic sediments that was described in the previous chapter. These sediments were deposited on the southern shores of Tethys. In the Atlas Mountains the sedimentary section attains its maximum thickness, with rocks of all geological periods represented. It is believed, however, that the Atlas fold belt accreted onto the African Shield during late lateral crustal movement as the Tethyan Ocean closed in to the Mediterranean sea (Clifford, 1986). Because the Atlas Mountains are newcomers, and thus not really part of mainland Africa, they are not described in the present volume. It is not easy to clearly define the sedimentary basins of the northwest Sahara. Some are easily recognised as quadripetally closed basins, but many are northerly plunging embayments that originally opened out into the great Tethyan ocean, to become subsequently closed. Furthermore some

basins change name where they cross an international boundary, or according to different authors. Bearing these qualifications in mind the main northwestern Saharan sedimentary basins from west to east are the Tindouf, Reggane, Ahnet, Mouydir and Illizi/Ghadames, Murzuk and Kufra (Fig. 1). The structure and tectonic evolution of these basins will now be defined.

TINDOUF BASIN

The Tindouf basin is the most westerly of the Saharan basins. It is elongated east to west. Its western edge is truncated by the Cretaceous Atlantic coastal sag basin in Western Sahara. It is closed off some 700 km to the east in western Algeria. The basin is asymmetric in cross-section, with a steep northerly limb, and a gentle southerly limb. It is infilled by some 8km of sediments of Cambrian to Carboniferous (Namurian) age, that broadly conform to the pan-Saharan stratigraphy described in the previous chapter. Unusually for much of the Sahara, the presence of archaeocyathid limestones at the base of the section demonstrates a (Lower) Cambrian age. There are several horizons of oolitic glauconite and phosphate in the Lower Ordovician section. A major glaciated erosion surface of pre-Ashgillian age transgresses across earlier sediments on to Precambrian basement (Deynoux et al., 1985). There are some Late Devonian dolerite and andesite lavas. The Tindouf basin has a history comparable to that of many of the Saharan basins. It was a northerly plunging embayment into which the Tethyan seas advanced intermittently from the Ordovician until the Carboniferous. The Tindouf embayment became a closed basin in the Late Carboniferous. This was a widespread phase of tectonic movement that occurred across much of the Sahara. This episode is commonly, but rather unfortunately, referred to

African Basins. Sedimentary Basins of the World, 3 edited by R.C. Selley (Series Editor: K.J. Hsti), pp. 17-26. 9 1997 Elsevier Science B.V., Amsterdam. All rights reserved.

18

R.C. S E L L E Y

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F i g . 1. Structural map to shown the distribution of the various sedimentary basins of Algeria. Simplified from S O N A T R A C H ( 1 9 7 9 ) . Note that there is a considerable variation of terminology for the different basins and subbasins. In particular the Illizi basin metamorphoses into the Ghadames basin when traced eastwards into Libya.

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Fig. 2. Pre-Mesozoic subcrop map of Algeria, simplified from Legrand (1985). This demonstrates that many basins had formed during the Hercynian tectonic event at the end of the Carboniferous Period, and before the deposition of Mesozoic strata. Lines A - B , C - D and E - F refer to cross-sections illustrated in Fig. 3.

THE BASINS OF NORTHWEST AFRICA: STRUCTURAL EVOLUTION

19

Fig. 3. Cross-sections to illustrate the architecture of Algerian sedimentary basins. Locations shown in Fig. 2. A-B, and C-D based on SONATRACH (1979), E - F based on Chiarelli (1978).

as the Hercynian Orogeny, though it involved little more than the uplift and northerly closure of many of the Saharan basins. There are no post-Carboniferous sediments in the Tindouf basin, apart for some 100 m of limestones, sandstones and shales of Miocene to Pliocene age.

REGGANE BASIN

The Reggane basin lies between the Tindouf and Ahnet basins. It is elongated northwest to southeast, with a length of some 350 km and a width of some 150 km. Like the Tindouf basin, it is also asymmetric, with a gentle southwestern limb that dips gently off the Reguibat massif, and a steeper northeastern limb that is bounded by the Ougarta range. The Reggane basin is infilled by some 5 km of sediments of Cambrian to Carboniferous (Namurian) age, that

broadly conform to the pan-Saharan stratigraphy described in the previous chapter. Dolerite and andesite lavas near the base of the section give radiometric dates that indicate a Cambrian age. There are also Late Devonian and Late Carboniferous (Namurian) lavas intercalated with the sedimentary section. Thin evaporites occur interbedded with dolomites and limestones in the Upper Visean section. The Reggane basin has a history comparable to that of the Tindouf and many of the Saharan basins. It was a northeasterly plunging embayment into which the Tethyan seas advanced intermittently from the Ordovician until the Carboniferous. The Reggane embayment became a closed basin by the uplift of the Ougarta range in the Late Carboniferous during the Hercynian orogenic event. Unlike the Tindouf basin, however, the Reggane basin contains over 200 m of "Continental Intercalaire", the local name for the largely barren

20

R.C. SELLEY

continental Mesozoic clastics. There are also intermittent outcrops of Plio-Pleistocene alluvial and eolian sands, and lacustrine marls and limestones that locally attain thicknesses of 100 m or so.

AHNET, MOUYDIR AND ILLIZI/GHADAMES BASINS

The Tindouf and Reggane basins are clearly defined centripetally dipping features. Further east, however, the structure becomes more complex. As noted earlier, during the Palaeozoic Era the Sahara platform consisted of northerly plunging embayments that opened out into the great Tethyan ocean. These embayments become closed basins during the Hercynian tectonic phase. East from the Reggane basin the southern ends of the Ahnet, Mouydir and Illizi/Ghadames basins are clearly defined by their unconformable contact with Precambrian basement. When traced northwards into the subsurface their relationships with one another, and adjacent basins and subbasins become unclear. The Illizi basin of Algeria metamorphoses into the Ghadames basin where it extends northeastwards into Libya. These basins contain the classic pan-Saharan Palaeozoic stratigraphy described in the previous chapter. Indeed, many of the type sections occur where the formations crop out around the flank of the Hoggar Massif. Similarly, it was in this region that the evidence for the late Ordovician glaciation was first recognised. When traced northwards the Palaeozoic sediments of the Ahnet, Mouydir and Illizi basins are overlain by Mesozoic sediments. A sub-Mesozoic map is a useful way of demonstrating how the Hercynian tectonic event defined the architecture of the various basins and subbasins (Fig. 2). Beginning with Triassic evaporites, and following with shallow marine clastics and carbonates, the Mesozoic sediments form a cover that thickens northwards towards the Atlas Mountains. The Algerian basins host major oil and gas fields, such as Hassi Messaoud (Balducci and Pommier, 1970; Bachellen and Peterson, 1992) and Hassi er R'Mel (Magliore, 1970) respectively. These fields occur on the north-south aligned ridges that separate the various basins and subbasins. The negligible variation in Palaeozoic stratigraphy shows that these ridges did not exist at that time. As shown by the pre-Mesozoic subcrop map and cross-sections (Figs. 2 and 3) they developed during the Hercynian phase. This event defined the present boundaries of the various basins and subbasins. There is also evidence of rejuvenation of these positive features during the Cretaceous Period (Sonatrach, 1979). This last phase of movement was crucial, because it predated the maturation of petroleum in the Silurian Tannezuft shale and its migration into adjacent

Fig. 4. Cross-sections to illustrate the evolution of the Triassic salt basin of Algeria (after SONATRACH. 1979). Note that the Cambro-Ordovician reservoir sands of the Hassi Messaoud oil field are capped by Triassic evaporites. Thus the Tannezuft Shale (Silurian) source rock could not have generated petroleum until post-Triassic time.

petroleum reservoirs that range in age from CambroOrdovician to Triassic (Fig. 4). Thus the structures were formed in time to trap the migrated petroleum (Tissot et al., 1975, 1984; McGregor, 1996) When the Ahnet, Mouydir and Illizi/Ghadames basins are traced northwards it becomes hard to resolve basin architecture beneath the ever-thickening sequence of monoclinally dipping Mesozoic strata. This region is best described under the broad regional term of "the Triassic Salt basin". This includes within its bounds the Bechar basin, the Oued Mya basin, the Western Great Erg basin, the Mac Mahon basin, and the Polignac basin. The northern limit of the Triassic Salt basin is defined by a major tectonic feature at the foothills of the Atlas Mountains. This is termed the Atlas flexure, or more dramatically and appropriately, by French geologists "Le Accident sud-Atlasian". This line marks an abrupt increase in sediment facies and thickness from the Sahara platform in the south to the Atlas trough in the north. Palaeogeographic maps commonly show that this line coincides with major

THE BASINS OF NORTHWEST AFRICA: STRUCTURAL EVOLUTION

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changes in sedimentary facies or a lateral offset. The Atlas flexure is believed to be a major complex of wrench faults by which the sediments of the Atlas trough were laterally accreted on to the old African continent (Clifford, 1986; Klitgord et al., 1988). It is for this reason that an account of the Atlas trough has not been included in this volume.

MURZUK BASIN

Crossing from Algeria to Libya, the change from the Illizi to the Ghadames basin has been already noted. Southwestern Libya is occupied by a classic sag basin, whose sub circular shape is clearly visible on satellite photographs. This is termed the Murzuk basin, or occasionally the Djado basin. It covers some 40,000 square km, extending southwards into Niger. The Murzuk basin is separated from the Illizi basin to the west by a north-south ridge, the Tihemboka mole, to the north by the Gargaf arch, to the east by the Tibesti-Sirte arch, and to the south by the Precambrian basement of the Sahara (Fig. 5). The Murzuk basin contains some 5,000 m of sedimentary fill (Thomas, 1995). Along the southern margin of the basin the preCambrian basement is unconformably overlain by continental red beds termed the Mourizidie Formation (Jaque, 1963). This is attributed to the Infra-Cambrian, and correlated with the Purpre d'Ahnet of Algeria. The Mourizidie

Formation is followed by the typical pan-Saharan Palaeozoic sequence (Bellini and Massa, 1978). This crops out in a series of intermittent subconcentric escarpments around the basin margins. All the main formations are present. On the north flank of the basin, at Wadi es Shatti on the southem limb of the Gargaf arch, there are sedimentary iron ore deposits. These are magnetite oolite beds within the Upper Devonian Aouinet Ouenine Formation (Turk et al., 1980). The most dramatic escarpment of the Murzuk basin rim is provided, however, not by Palaeozoic formations, but by the Continental Mesozoic sandstones, known locally as the Messak Sandstone. This takes its name from the Messak escarpment on the northern flank of the basin (Klitzsch and Baird, 1969). This formation may attain a maximum thickness of over 1,500 m in the basin centre (Fig. 6). The age of these barren continental sediments is typically hard to establish. Spores of Triassic-Jurassic (sic) age have been recovered from the E1-NC58 well on the southwestern flank of the basin (Pierobon, 1991). Wood and leaf fragments recovered from the Messak scarp indicate an Early Cretaceous (Wealden) age (Klitzsch and Baird, 1969). These palaeontological data imply a prolonged, if intermittent, history of continental deposition throughout the Mesozoic Era. The basin centre contains an indeterminate fill of Holocene eolian dunes and Pleistocene alluvium.

22

R.C. SELLEY D

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1200 km T i b e s t i - S i r t e arch ( T e r t i a r y ba.salts & C r e t . T e r t i a r y c a r b o n a t e s locally o v e r l v basement)

r

East Kufra basin sea level

4 000 m Continental

I +•++

Mesozoic

palaeozoic

Precambrian

basement

Fig. 6. Cross-sections to illustrate the architecture of the Murzuk and Kufra basins. Based on Goudarzi (1980). For locations see Fig. 5.

Viewed from space, or from a geological map, the Murzuk basin looks like a typical intracratonic sag basin. Palaeocurrent studies show, however, that, like the Algerian basins to the west, the Murzuk basin was once a northerly plunging embayment that opened to the Tethyan ocean (Figs. 7-9). On the Gargaf arch, the structure that defines the present northern limit of the basin, the Melez Chogranne Shale Formation (Ordovician) is locally cut out by the Memouniat Sandstone Formation (Ordovician). Similarly the Tannezuft Shale Formation (Silurian) cross-cuts both of these to unconformably overly the Hassaouna Formation (Cambro-Ordovician) (Collomb, 1962). These unconformities show that the Gargaf arch was an intermittently positive feature in the Early Palaeozoic. Nonetheless palaeocurrents interpreted from cross-bedding in the fluvial Hassaouna (Cambro-Ordovician) sandstones show a uniform northerly flow direction (Burollet and Byramjee, 1969). Similar northerly palaeocurrents have been recorded from the continental Mesozoic sandstones of the Messak scarp by many geologists from McKee (1963) onwards. But Lorenz (1980), though recording northerly palaeocurrents, noted facies changes that suggested that the Gargaf arch was sufficiently uplifting to pond continental Mesozoic sediments within the Murzuk basin, and to only permit limited sediment transport northwards across the arch.

These data show that, like the Algerian basins described previously, the Murzuk basin was a northerly plunging embayment during episodes of Palaeozoic deposition. Intermittent uplift of the Gargaf arch is proven though, by the unconformities over its crest. This alternation of quiescence, when sediment was transported northwards across it, and uplift, when sediment was trapped by it, seems to have continued throughout the Mesozoic Era. Whereas the Algerian basins were closed off from the Tethys by the Late Carboniferous Hercynian tectonic event, the final closure of the Murzuk basin postdates the (Neocomian?) deposition of the continental Mesozoic sediments. Little is published about the internal subsurface structure of the Murzuk basin (Fig. 10). Klitzsch has published papers illustrating a complex system of Palaeozoic troughs and arches (e.g., Klitzsch, 1970). These features appear to have been recognised, not from geophysics or well control across the basin, but from thickness variations of formations measured from outcrops around its margin. It can be argued that these "arches" and "troughs" are more apparent than real. The thickness variations being due to the dissected nature of the basin margin scarps. Thus sections measured up the wadis towards the basin centre reveal greater thicknesses than are recorded from those measured on outlying jebels, where the section is naturally thinner.

23

THE BASINS OF N O R T H W E S T AFRICA: S T R U C T U R A L E V O L U T I O N

J

TheMediterranean ( ~" (

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Fig. 7. Map to show palaeocurrent directions for parts of northwest Africa. These include over 3000 readings taken from slope-controlled fluvial deposits of the Hassaouna Formation (Cambro-Ordovician), the Tadrart Formation (Devonian), diverse Carboniferous fluvial sandstones, and the Continental Mesozoic sandstones. Compiled from Clarke-Lowes and Ward (1991), Burollet and Byramjee (1969), Collomb (1962), Lorenz (1980), McKee (1963), Turner (1980) and other sources.

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.

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.

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Fig. 8. Outcrop map of the Palaeozoic sediments of southern Libya and environs showing regional palaeostrike lines, and depositional slope (black arrows), based on the data in Fig. 7. Note the northerly palaeoslope over the Gargaf arch, indicating that it was not a positive feature during the deposition of Palaeozoic sediments. Note the northeasterly dip on the southwestern flank of the Kufra basin, indicating the existence of the Tibesti-Sirte arch. Palaeocurrent data show that neither the Murzuk, nor the Kufra basins were structurally closed during the Palaeozoic.

0

50(~m

I

1

Fig. 9. Outcrop map of Continental Mesozoic sediments of southern Libya and environs showing regional palaeostrike lines, and depositional slope (black arrows), based on the data in Fig. 7. Note the northerly palaeoslope on the north flank of the Murzuk basin, indicating that it was not a closed basin during the deposition of these sediments. Note the northeasterly dip on the southwestern flank of the Kufra basin, indicating the existence of the Tibesti-Sirte arch. Palaeocurrent data show that neither the Murzuk, nor the Kufra basins were structurally closed during the deposition of the Continenal Mesozoic sediments.

24

R.C. SELLEY

Fig. 10. Structure contour map of Libya to show the architecture of the Murzuk (lower left), Kufra (lower right) and Sirte (upper centre) basins. From Goudarzi and Smith (1978), see also Goudarzi (1980). D-E and G-F are the locations of the cross-sections shown in Fig. 6.

Drilling for oil has gone on in the Murzuk basin for many years, and indeed it was the site of Libya's first oil discovery. In 1957 oil shows were discovered in Devonian and Carboniferous sands in the Atshan wells on the edge of the Gargaf arch (Colley, 1963). Recently oil has been discovered in the Memouniat (Ordovician) sandstone on the northwestern part of the basin (Meister et al., 1991; Thomas, 1995). Sadly there is negligible published information on the subsurface architecture of the Murzuk basin.

THE KUFRA BASIN

The Kufra basin is the last and most easterly of the Saharan basins to be considered in this chapter. The Kufra basin occupies a large part of south east Libya, though it extends northeastwards into Egypt, and southeast into Sudan and southwest into Chad. The Kufra basin is separated from the Murzuk basin to the west by the Tibesti-Sirte arch, and from the Sirte basin to the north by the Calanscio

THE BASINS OF NORTHWEST AFRICA: STRUCTURAL EVOLUTION arch. Its eastern limb forms the westem edge of the Arabo-Nubian shield to the east. The Precambrian Ennedi craton defines its southern rim (Fig. 5). The Kufra basin is one of the least accessible and least known of all the Saharan basins. Nonetheless there are published accounts of its Palaeozoic sediments, notably by Turner (1980, 1991), and of its continental Mesozoic sediments, notably by Van Houten (1980) and Klitzsch and Squyres (1990). The Palaeozoic sediments are broadly comparable in facies, and in their gross stratigraphic sequence to those further west. Precambrian basement is unconformably overlain by braided alluvial cross-bedded sands comparable to those of the Hassaouna Formation. They are commonly interbedded with Tigillitiferous sands of typical Haouaz type. It is tempting to attempt to recognise the Memouniat and Melez Chogranne formations of western Libya in the Kufra basin, but, though the facies are comparable, lithostratigraphic continuity may not exist. The Tannezuft shale is present, however, and clearly recognisable and datable due to the presence of graptolites. It is, however, notably silty, and lacks the oil source rock characteristics that it shows in Algeria and western Libya. The Tannezuft Shale Formation passes up into the Acacus Sandstone Formation with the gradational sequence of upward-coarsening cycles that are seen to the west. The Tadrart Sandstone Formation (Devonian) is also recognisable, with both shallow marine and fluvial deposits being present. Later Devonian sediments are locally termed the B inem Formation on the western side of the basin, and the Blita Formation on the eastern side. These cyclically arranged sands and shales compare in facies with the shallow marine sediments of the Aouinet Ouenine and Ouan Casa formations of the Murzuk and Hammada basins to the west and northwest. The Carboniferous rocks of the Kufra basin are, however, markedly different from those of western Libya. Carbonate facies, such as occur in the Assedjefer and Dembaba formations of the Murzuk basin, are absent. They are replaced by coarse cross-bedded and channelled fluvial sands with plant remains such as Lepidodendron. The centre of the Kufra basin is infilled with continental Mesozoic clastics (Van Houten, 1980, and Klitzsch and Squyres, 1990). These are broadly comparable in facies to those of the Murzuk basin, as described in detail in the previous chapter. Bellini et al. (1991) recognised three units of continental Mesozoic sediments. The lowest is composed of typical fluvial sands and shales. They apply the Algerian term "Continental Post-Tassilian" to these sediments, and suggest a "Permian to Jurassic" age, citing the existence of a rich Permian pollen assemblage in borehole samples. This is overlain by a curious silicified limestone, that they term the

25

Chieun Formation. This locally contains abundant

Hydrobia, indicating a lacustrine environment, and a post-Triassic age. The topmost continental Mesozoic sediments they describe as "Nubian sandstone", and suggest an Early Cretaceous age. Similarly K1itzsch and Squyres (1990) in their synthesis of the "Nubian" sandstones of northeastern Africa offer no age for the continental Mesozoic sediments of the Kufra basin beyond "undifferentiated Jurassic to Cretaceous". Thicknesses of the various formations of the Kufra basin have been published from surface sections measures around the basin margin. With negligible oil exploration in the basin little is known about the thicknesses of the formations in the basin centre, and still less is publicly available (Fig. 10). Using a combination of surface and borehole data Bellini et al. (1991) suggest that there is a total thickness of some 3,500 m of sediment in the Kufra basin. The structure and tectonic history of the Kufra basin is also little known. The Kufra basin shares many features with the basins to the west. These include a similarity of facies and stratigraphy, together with predominantly northerly palaeocurrents in fluvial sediments. These features all suggest that the Kufra basin formed part of the undifferentiated Sahara platform, from the Cambrian until at least the Cretaceous Period. Thus the present synclinal shape of the Kufra basin most probably postdates the deposition of the sediments with which it is filled, and was coeval with the break up of Africa and the collapse of the Tibesti-Sirte arch in the mid-Cretaceous (Fig. 6). Trans-Saharan facies belts tend to show southwest-northeast trends. The more continental aspect of the facies in the Kufra basin compared with those of the western basins suggest that it lay further from the shores of the Tethyan Ocean.

SELECTED BIBLIOGRAPHY

Naturally an area as large as northwest Africa, with its great mineral wealth, has attracted a large literature. Important source books include: Holland, C.H., 1985. Lower Palaeozoic of North-Western and West-Central Africa. Wiley,Chichester, 512 pp. Salem, M.J. and Busrewil, M.T. (Editors), 1980. Geology of Libya, Vols. I-III. Academic Press, London, pp. l- 1712. Salem, M.J. and Belaid, M.N. (Editors), 1991. Geology of Libya, Vols. IV-V. Elsevier, Amsterdam, pp. 1713-2095. Salem, M.J., Sbeta, A.M. and Bakbak, M.R. (Editors), 1991. Geology of Libya, Vol. VI. Elsevier, Amsterdam, pp. 20992491. SONATRACH, 1979. Geology of Algeria. In: J.L. Chardac (Editor), Well Evaluation Conference 1979. Schlumberger, Paris, pp. 1-26.

26 REFERENCES Bachellen, W.D. and Peterson, R.M., 1992. Hassi Messaoud Field. In: N.H. Foster and E.A. Beaumont (Editors), Structural Traps V. American Society of Petroleum Geologists, Tulsa, OK, pp. 211-226. Balducci, A. and Pommier, G., 1970. Cambrian Oil field of Hassi Messaoud. In: M.T. Halbouty (Editor), Geology of Giant Petroleum Fields. Am. Assoc. Petrol. Geol. Mem., 14: 477488. Bellini, E. and Massa, D., 1980. A stratigraphic contribution to the Palaeozoic of the southern basins of Libya. In: M.J. Salem and M.T. Busrewil (Editors), Geology of Libya, Vol. I. Academic Press, London, pp. 3-56. Bellini, E., Giori, I., Ashuri, O., and E Benelli. Geology of AI Kufrah Basin, Libya. In: M.J. Salem, A.M. Sbeta and M.R. Bakbak (Editors), Geology of Libya. Vol. VI. Academic Press, London, pp. 2155-2185. Burollet, P.E and Byramjee, R., 1969. Sedimentological remarks on the Lower Palaeozoic sandstones of southern Libya. In: W.H. Kanes (Editor), Geology, Archaeology and Prehistory of Southwestern Fezzan. Petrol. Explor. Soc. Libya, Tripoli, pp. 91-102. Chiarelli, A., 1978. Hydrodynamic framework of eastern Algerian Sahara - - influence on hydrocarbon occurrence. Bull. Am. Assoc. Petrol. Geol., 62: 667-686. Clarke-Lowes, D.D. and Ward, J., 1991. Palaeoen- vironmental evidence from the Palaeozoic "Nubian Sandstones" of the Sahara. In: Salem, M.J., A.M. Sbeta and M.R. Bakbak, The Geology of Libya. Vol. VI. (Editors), Elsevier, Amsterdam, pp. 2099-2154. Clifford, A.C., 1986. African oil-past, present and future. In: M.T. Halbouty (Editor), Future Petroleum Provinces of the World. Am. Assoc. Petrol. Geol. Mem., 40: 339-372. Coiley, B.B., 1963. Libya: petroleum geology and development. 6th World Petrol Cong. Frankfurt, Sect. 1, Paper 43, 10 pp. Collomb, G.R., 1962. Etude g6ologique du jebel Fezzan et de sa bordure Palaeozoique. Notes Mem. C.EP., 1, 36 pp. Deynoux, M., Sougy, J. and R. Trompette., 1985. Lower Palaeozoic rocks of West Africa and the western part of Central Africa. In: C. Holland (Editor), Lower Palaeozoic of NorthWestern and West-Central Africa. Wiley, Chichester, pp. 337496. Goudarzi, G.H., 1980. Structure m Libya. In: M.J. Salem, M.T. Busrewil and A.M. Ben Ashour (Editors), The Geology of Libya, Vol. III. Academic Press, London, pp. 879-892. Goudarzi, G. and Smith, J.P., 1978. Preliminary structure contour map of the Libyan Arab Republic and adjacebt areas. U.S. Geol. Surv. Misc. Geol. Invest., Map 1-350C. Jaque, M., 1963. Reconnaissance g6ologique du Fezzan orientale. Notes Mem. C.F.P., 5, 44 pp. Klitszch, E., 1970. Die Strukturgeschichte der Zentralsahara. Geol. Runsch., 59: 459-527. Klitzsch, E. and Baird, D.W., 1969. Stratigraphy and palaeohydrology of the Germa (Jarma) area, southwest Libya. In: W.H. Kanes (Editor), Geology, Archaeology and Prehistory of Southwestern Fezzan. Petrol. Explor. Soc. Libya, Tripoli, pp. 67-82.

R.C. S E L L E Y Klitzsch, E. and Squyres, C.H., 1990. Paleozoic and Mesozoic geological history of northeastern Africa based upon new interpretation of Nubian strata. Bull. Am. Assoc. Petrol. Geol., 74:1203-1211. Legrand, P.H., 1985. Lower Palaeozoic rocks of Algeria. In: C.H. Holland (Editor), Lower Palaeozoic of North-Western and West-Central Africa. Wiley, Chichester, pp. 5-90. Lorenz, J., 1980. Late Jurassic-Early Cretaceous sedimentation and tectonics of the Murzuk Basin. In: M.J. Salem, M.T. Busrewil and A.M. Ben Ashour (Editors), The Geology of Libya, Vol. II. Academic Press, London, pp. 383-418. Maggliore, P.R., 1970. Triassic gas field of Hassi er R'Mel. In: M.T. Halbouty (Editor), Geology of Giant Petroleum Fields. Am. Assoc. Petrol. Geol. Mem., 14: 489-501. McGregor, D.S., 1996. The hydrocarbon systems of North Africa. Mar. Pet. Geol., 13: 329-340. McKee, E.D., 1963. Origin of Nubian and similar sandstones. Geol. Runsch., 52:551-587. Meister, E.M., Ortiz, E.F., Pierobon, E.S.T., Arruda, A.A. and Oloviera, M.A.M., 1991. The origin and migration of fairways of petroleum in the Murzuq Basin, Lybia: An alternative exploration model. In: M.J. Salem, M.T. Busrewil and A.M. Ben Ashour (Editors), The Geology of Lybia, Vol. VII. Elsevier, Amsterdam, pp. 2725-2742. Pierobon, E.S.T., 1991. Contribution to the stratigraphy of the Murzuk basin, S.W. Libya. In: M.J. Salem, M.T. Busrewil and A.M. Ben Ashour (Editors), The Geology of Libya, Vol. V. Academic Press, London, pp. 1767-1783. SONATRACH, 1979. Geology of Algeria. In: J.L. Chardac (Editor), Well Evaluation Conference, 1979. Schlumberger, Paris, pp. 1-26. Thomas, D., 1995. Geology, Murzuk oil development could boost S.W. Libya prospects. Oil Gas J, March 6, pp. 41-46. Tissot, B., Deroo, G. and J. Espitalie., 1975. Etude compar6e de l'6poque de formation et d'expulsion du p6trole dans diverses provinces g6ologique. In: 9th World Petrol. Cong. Proc., Applied Sci. Publ., London, pp. 159-169. Tissot, B., Espitalie, J.Deroo, G., Tempere, C. and D. Jonathan., 1984. Origin and migration of hydrocarbons in the eastern Sahara (Algeria). In: G. Demaison and R.J. Murris (Editors), Petroleum Geochemistry and Basin Evaluation. Am. Assoc. Petrol. Geol. Mem., 35:315-334. Turner, B.R., 1980. Palaeozoic sedimentology in the southeastern part of the A! Kufrah Basin, Libya: A model for oil exploration. In: M.J. Salem and M.T. Busrewil (Editor), Geology of Libya, Vol. I. Academic Press, London, pp. 351-374. Turner, B.R., 1991. Palaeozoic deltaic sedimentation in the southeastern part of the AI Kufrah Basin, Libya. In: M.J. Salem and M.N. Belaid (Editor), Geology of Libya, Vol. III. Academic Press, London, pp. 1713-1726. Turk, T.M., Abdolrahman, K., Doughri, K. and S. Banerjee., 1980. A review of the recent investigations on the Wadi ash Shati iron ore deposits, northern Fazzan, Libya. In: M.J. Salem and M.T. Busrewil (Editor), Geology of Libya, Vol. III. Academic Press, London, pp. 1019-1043. Van Houten, EB., 1980. Latest Jurassic-Early Cretaceous regressive facies, northeast Africa Craton. Bull. Am. Assoc. Petrol. Geol., 64: 857-867.

Chapter 3

The Sirte Basin of Libya

R.C. SELLEY

positive feature throughout much of Palaeozoic and Mesozoic time, separating the Murzuk and Kufra embayments. With the break up of the African continent in the Cretaceous Period the northern part of the Tibesti-Sirte arch collapsed to form the Sirte basin. This is, more strictly, an embayment that opens out northwards into the Mediterranean basin. Thus the floor of the Sirte basin is a major unconformity, above which is a thick sequence of Late Cretaceous to Recent sediments. The following account deals first with the stratigraphy and sedimentology of the basin, and then with its structure and tectonic evolution.

INTRODUCTION The Sirte basin is a major sedimentary basin that extends southwards from the Gulf of Sirte in central Libya (Fig. 1). Unlike the basins of southern Libya and Algeria there is little surface expression of the Sirte basin. It lies beneath vast sarir gravel plains, with occasional sand seas and escarpments. The existence of the Sirte basin was unknown until gravity and magnetic surveys were carried out in the late nineteen-fifties as part of the quest for petroleum. Subsequent seismic surveys and drilling have proven up a major petroleum province, with estimated reserves in excess of 45,000 million barrels of oil and gas oil equivalent (Thomas, 1995). Despite over 35 years petroleum exploration, however, relatively little has been published about the Sirte basin. The Sirte basin occupies a collapsed north-south trending positive feature, the Tibesti-Sirte arch. As discussed in the previous chapter, this arch was a

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Mediterranean Sea

k N

STRATIGRAPHY AND FACIES

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,/ kEGYPT

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Fig. 1. Structure contour map on the mid-Cretaceous floor of the Sirte basin, showing the location of some of the major oil fields. Contour interval in kilometres (after Sanford, 1970). African Basins. Sedimentary Basins of the World, 3 edited by R.C. Selley (Series Editor: K.J. Hs~i),pp. 27-37. 9 1997 Elsevier Science B.V., Amsterdam. All rights reserved.

28

R.C. SELLEY Tunisia ~ l r . /"

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Fig. 2. Subcrop map of the floor of the Sirte basin, based on Hea (1971) and Clifford et al. (1980). This shows how the Sirte unconformity cross cuts the collapsed Tibesti-Sirte arch. Subcropping formations young to the west and to the east, away from PreCambrian basement on the crest of the former arch, now the axis of the Sirte basin. These subcropping formations and basement are productive petroleum reservoirs within the Sirte basin.

gressively younger formations moving westwards, and eastwards (Fig. 2). This reflects the truncation of the Tibesti-Sirte arch, prior to its collapse. The Precambrian basement includes a wide range of igneous and metamorphic rocks, that include granites and volcanics. When uplifted, weathered, and fractured these can serve as petroleum reservoirs, as in the Augila field (Williams, 1968, 1972). Another important rock type beneath the unconformity may be loosely termed quartzite. These are often quartzites in the sense that they are mineralogically very mature, and also in that they are extremely well-cemented. Primary intergranular porosity has been destroyed. But because of their brittle nature, secondary fracture porosity has often been induced. For this reason, like the basement igneous and metamorphic rocks, they too can serve as petroleum reservoirs, as in the Samah and Raguba oil fields (Brennan, 1992). The stratigraphic position of these "quartzites" is hard to resolve. It is possible that they are the epidiagenetic product of intense weathering operating across a wide range of stratigraphic formations during the prolonged uplift of the Tibesti-Sirte arch (Hea, 1971). A third common lithology beneath the floor of the Sirte basin are coarse-cross-bedded sandstones. These are largely unfossiliferous. As discussed at length in Chapter 1, the Sahara is covered by a discontinuous blanket of unfossiliferous braided alluvial sands. Dates obtained from fauna in rare intervening shales show that the barren fluvial sands range in age from Precambrian to Recent. This makes it very difficult, if not impossible, when trying to produce a subcrop map of the floor of the Sirte basin. There has been a tendency for oil companies to generate a new formation name every time they

encounter barren fluvial sands in a newly discovered oil fields. E.g. the Sarir Sandstone of the Sarir Field (Sanford, 1970), the Amal Formation of the Amal Field (Roberts, 1970), and the Hofra Formation of the Hofra Field (Barr and Weeger, 1972). These formations have been attributed to ages ranging from Cambro-Ordovician to Early Cretaceous. Palaeontological data sometimes support these ages. The fluvial sands make important petroleum reservoirs in the fields mentioned above, as well as Messla (Clifford et al., 1980), and many others. Primary intergranular porosity is often enhanced by the solution of feldspar and early carbonate cements. This has been attributed to leaching due to acid meteoric waters beneath the Sirte unconformity (Hea, 1971). On the flanks of the Sirte basin Devonian and Carboniferous sandstones and shales have been encountered. The oldest rocks encountered above the Sirte basin unconformity are of Cenomanian age. This is thus taken as the date for the collapse of the Tibesti-Sirte arch. It must be remembered, however, that these dates come from wells drilled on the crest of structural highs. It is probable that sedimentation began earlier in the intervening undrilled troughs. A Cenomanian age for the collapse of the Tibesti-Sirte arch would accord, however, with the regional picture of early Cretaceous trans-African rifting (Clifford, 1986). The floor of the Sirte basin is extensively faulted. These faults exerted a strong control on initial sedimentation, but this control gradually diminished through time, with the few exceptions noted later in this chapter. Immediately above the Sirte unconformity early Cretaceous sands thin out by nonde-

THE SIRTE BASIN OF LIBYA

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Fig. 3. Summary stratigraphic column for the Sirte basin. No attempt has been made to show formation thicknesses due to rapid lateral changes from trough to high, especially in the older part of the section. Definitions and details of the type sections of the various formations will be found in Barr and Weeger (1972).

position or truncation on to the crests of the fault blocks. Late Cretaceous and Palaeocene shales in the troughs pass up into reefal carbonates on the crests of the horsts (Fig. 3). The term Bahi Formation is given to the basal sands of the Sirte basin sequence (Barr and Weeger, 1972). Though barren, and comparable to the Continental Mesozoic sands of southern Libya, glauconite is present in their upper part, indicating the onset of the Sirte marine transgression. The early Cretaceous age is inferred from their position above the unconformity, and by the fact that they are immediately overlain by the Lidam Formation that contains Cenomanian fossils. The Lidam Formation was deposited by the advancing marine transgression as the Sirte trough subsided. It consists of thinly interbedded sands, shales, limestones, dolomites with nodular anhydrite arranged in typical sabkha cycles. The Lidam Formation contains glauconite, phosphate and a diverse and locally abundant assemblage of marine fossils. The Lidam Formation is abruptly overlain by the Sirte Shale Formation. Though less than 200 m thick in the type section, the Sirte Shale thickens

rapidly off the horsts on which it is drilled, into the intervening troughs. On the crests of the horsts it locally oversteps the Lidam and Bahi formations to directly overly the Sirte unconformity. The Sirte Shale is a dark grey to dark brown organic rich shale. It is the main source rock for the petroleum within the Sirte basin, with TOCs averaging 2% (Parsons et al., 1979). Kerogen is of Type I and Type II, (oil and oil and gas prone) varieties (Hamyouni, 1984). It contains abundant pelagic foraminifer that indicate open marine conditions, and a Maastrichtian to Campanian age (Barr and Weeger, 1972). The abrupt change from the sabkha environment of the Lidam Formation, the fine grain size, open marine fauna, and transgressive base of the Sirte Shale Formation, all point to a rapid rise in sea level, and rapid subsidence of the basin floor. Pelagic mud deposition continued in the troughs throughout the Palaeocene, giving rise to the overlying Hagfa Shale Formation. This also contains organic matter, but it is not as rich as the underlying Sirte Shale. The Palaeocene Hagfa shales pass up dip into carbonates on the crests of the horsts. These carbonates consist of a diversity of facies. They include, coralgal, bryozoan, foram, and other bioclasts in a range of textures from grainstone to wackestone. Ooids are very rare, and oolitic limestone formations unknown. This implies that there were negligible tidal currents within the Sirte embayment. There is no marked development of bioherms beyond a few isolated patch reefs. The grain types and textures of these sediments suggest that they were deposited on broad shallow sand and mud banks. The carbonates have generally undergone extensive diagenesis that has resulted in porosities of up to 40%. This was caused by meteoric flushing of the limestones during phases of emergence, leading to the development of secondary solution porosity and dolomitization (Bebout and Pendexter, 1975). The carbonates on the crests of the horsts are major petroleum reservoirs. As with the basal clastics, there was a tendency for each oil company to generate a new formation name for every new field reservoir. Thus in the Nasser (formerly Zelten) field it was termed the Ruaga Limestone, in the Gialo and Sahabi fields, the Sabil Formation, and in the Bahi and Dahra "B" fields, the Satal Formation. This pattern of shale and marl deposition in the troughs, and shallow water carbonate deposition on the horsts, continued intermittently from the Maastrichtian until the Late Palaeocene (Garea and Braithwaite, 1996). Occasional sea level rises allowed deeper water marls, such as the Khalifa Formation, to extend over the horst crests, while drops in sea level lead to emergence, leaching and porosity enhancement of the carbonates. In the Late Palaeocene a complex of pinnacle reefs developed

30

R.C. SELLEY

DEPOSITIONAL

DETRITAL LIMESTONE FACIES cross-bedded shell s e n d s deposited by s h o r e w a .rd migrating megaripples

ENVIRONMENTS

L A M I N A T E D SHALE FACIES lamination, oyster beds, open closed lagoon lagoon

INTERLAMINATED SHALE & SAND FACIES rippled , b u r r o w e d & w i t h mudfiiled channels. intePtidal flat & c r e e k

GEOMORPHOLOGY

o f f s h o r e b a r s 8, barrier beaches

~

Ls & creeks Fluviatile coastal plain 9

CROSS BEDDED SAND & SHALE FACIES l i g n i t e s , r o o t l e t beds & palaeosols. fluviatile

".' 9~:~.

CALCAREOUS SANDSTONE CHANNEL FACIES r a d i a t i n g s e a w a r d t r e n d i n g s h o e s t r i n g complexes, b i p o l a r cross bedding, m i x e d continental 8, m a r i n e fossilS, e s t u a r i n e channels

Fig. 4. Geophantasmogram to illustrate the facies and the sedimentary environments of the Miocene shoreline of the Jebel Zelten area (from Selley, 1969). This model can be applied to most of the post-Eocene sediments of the Sirte basin.

on the eastern flank of the Marsa Brega trough. Some of these host the giant Intisar (formerly Idris) oil fields (Terry and Williams, 1969; Brady et al., 1980). The complex lateral facies variations seen in the Palaeocene sediments, are replaced by a greater lateral continuity of stratigraphy in Eocene sediments. This indicates that fault movement in the Sirte basin began to be less dramatic. The oldest Eocene rocks comprise the Gir Formation. This consists of limestones, dolomites and evaporites, often arranged in a regular cyclic motif. The evaporites include gypsum (near the surface), anhydrite, and halite in the south western part of the basin. The Gir Formation passes up into the Gialo Limestone Formation. This is a Nummilitic limestone, that is also laterally continuous across much of the Sirte basin. This formation is the main reservoir of the Gialo field. There is then a major disconformity between the Gialo and the overlying Augila Formation. This too consists of Nummilitic limestones, but these are interbedded with thin shales and glauconitic sands. A further depositional break separates the Eocene sediments from overlying Oligocene sediments. These commence with several upward-coarsening cycles of shale and glauconitic sands laid down in coastal barrier conditions. These, the Arida Formation, serve as a petroleum reservoir in the Gialo field. The Arida Formation passes conformably up into the glauconitic sands and shales of the Dibba Formation. There is a major depositional break between the Oligocene and the Miocene sediments of the Sirte basin. Though there is no angular discordance

between the two, there is locally evidence of an extensive palaeosol, penetrated by tree roots, beneath the overlying Miocene Marada Formation. Whereas all the previously described formations of the Sirte basin are only known in the subsurface, the Marada Formation crops out extensively in jebels and scarps around Marada Oasis and the Jebel Zelten. Detailed sedimentologic analysis of the Marada Formation shows that it was deposited in a suite of environments that paralleled the shoreline of the Sirte basin (Fig. 4). Basinal marls pass south into skeletal shoal sands. These separated the open sea from a lagoon, that passed landwards, via tidal flats, into an alluvial coastal plain which supported a rich fauna indicative of savannah conditions (Selley, 1969; EI-Hawat, 1980). The Sirte basin is extensively dissected by a complex of Late Miocene deep steep-sided sandfilled channels that drained northwards into the Gulf of Sirte. This feature, termed the Sahabi Channel, provided early impressive evidence to support the argument for the Messinian salinity crisis of the Mediterranean (Barr and Walker, 1972). Coastal Pliocene lagoonal sediments in the Sahabi area are noted for their vertebrate fauna. Elsewhere continental Pliocene sediments can only be differentiated from Quaternary ones by palaeontology. Both consist of fluvial sands and lacustrine shales, marls and limestones. These beds crop out intermittently across the Sirte basin, whose surface, as mentioned earlier, consists largely of the flat gravel plain that is termed Sarir.

THE SIRTE BASIN OF LIBYA

31

STRUCTURE AND TECTONIC EVOLUTION

The progressive opening up of the Atlantic ocean was associated with the development of an extensive rift systems across much of Africa (Fairhead and Binks, 1991). The African rifts developed in response to radial lithospheric divergence from the village of Abong M'Bang in central Cameroon (Fig. 5) (Pavoni, 1993). The Sirte basin may be regarded as one of the Cretaceous collapse rifts that failed to open (Van Houten, 1983). As demonstrated in the previous chapter, however, the Sirte basin actually occupies the axis of the Tibesti-Sirte arch, a positive feature that separated the Kurzuk and Kufra embayments from Cambrian times up until the mid-Cretaceous collapse. The Sirte embayment can therefore be regarded as a major inversion structure (Fig. 6). Detailed mapping of the Sirte basin reveals that it is composed of a series of gently curved horsts and graben (Fig. 7). The faults that define these features are characterized by abrupt dog-legs, or local sudden curves. Many of them are scissor faults, whose direction of down-throw changes along the length of the fault. The most active time of fault movement was in the mid-Cretaceous, gradually dying out towards the end of the Palaeocene. Neogene and

AbongM'Bang .~r.. / South America

0

3 0 0 0 km

maximum e x t e n t of marine Cretaceous

~ "~

direction of crustal divergence

Cretaceous rifts

Fig. 5. Map of Africa to show the tectonic setting of the Sirte basin during the mid-Cretaceous phase of Pan-African rifting (according to Pavoni, 1993), together with the Cretaceous seaways at the time of maximum transgression (according to Clifford, 1986).

N

4-

I Palaeozoic

-I-,.. -~- +

J

Ti best i- S i r t e arch

+

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Murzuk basin Sirte embayment .

.

.

.

.

.

.

Gargaf arch IT[ m i d - C r e t a c e o u s -Tertiary

.

.

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.

.

.

Kufra basin

~, collapsed arch

Fig. 6. Geophantasmograms to show the structural evolution of Libya, dominated by the mid-Cretaceous collapse of the Tibesti-Sirte arch.

younger faulting is largely restricted to the western part of the basin, apart from the Miocene inversion phase outlined later. The faults define a series of troughs, or subbasins. The main troughs are, from west to east, the Hon, Zella (Tagrifet), Hagfa (Marada) and Agedabia (Marsa Brega) rifts (Fig. 8). As noted earlier these are the "kitchens" in which the Upper Cretaceous and Palaeocene petroleum source rocks are thickest, richest, and, because of crustal thinning, hottest. Burial curves produced by Gumati and Kanes (1985) and Gumati and Nairn (1991) show that subsidence began in the Cenomanian, accelerated throughout the Palaeocene, and then slowed down again (Fig. 9). At their present burial depth the Palaeocene source rocks are immature, and the Late Cretaceous source rocks are just entering the oil window. This seems superficially bizarre, since the source rocks have demonstrably generated large quantities of petroleum within the Sirte basin. Two possible explanations may be advanced. Either the source rocks were matured by a hot flush, due to a spell of heightened heat flow, sometime between the Palaeocene and the present day, or the Sirte basin has undergone inversion (Gumati and Schamel, 1988; Gumati et al., 1995). Both of these are possible. A pulse of high heat flow has been recognised in sedimentary basins elsewhere, such as the North Sea. Tectonic inversion in the S irte basin

I"J

THE MEDITERRANEAN SEA M

E

D

I

7"

E'

R

R

A

GULF OF

N

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S

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l

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i

o

,

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.-

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.

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~

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i

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I

I

I

Fig. 7. Left: Structure contour map of the Sirte basin, drawn on the Sirte unconformity (approximate base mid-Cretaceous), from Goudarzi and Smith (1978), see also Goudarzi (1980). Right: Map of the Sirte basin to show the major tectonic features, and the locations of the cross-sections in Figs. 8 and 10. Abbreviations: Ho Gr: Hon graben, Ze/Ta Tr: Zella/Tagrifet trough, Da-Ho PI: Dahra-Hofra platform, Ra-Sa PI" Raguba-Samah platform, Ze-De PI" Zelten-Defa platform, Ja PI: Jahama platform, Ag/Br Tr: Agedabia/Brega trough, Am/Gi H: Amal/Gialo High, Ma/Ha Tr: Marada/Hagfa trough.

~0

t" t"' ,.<

THE SIRTE BASIN OF LIBYA

A South West

33

B North West

575 km

-

Zella trough

Dahfra Hofra platform

Hagfa Jamama trough platform

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Cyrenaica platform

sea

sea

level

Iev e l

depth

(km)

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~

shale

~

carbonates

//

assorted super& sub-Sirte unconformity sandstones largely continental age Cambro-Ordovician to mid Cretaceous Precambrian basement

evapo ri t es Fig. 8. Cross-section across the Sirte basin along the line A-B in Fig. 7.

Cretaceous

~cTene T

sea level

100

1000

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2000

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90 ~

70

time of ~source rock ~ f~176

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50

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10

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sea level

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2000

3000

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4000

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Fig. 9. Burial curve for the Sirte basin basal unconformity. Based on Gumati and Nairn (1991). Note how maximum subsidence in the Palaeocene correlates with maximum faulting and lateral sedimentary facies and thickness changes. is well known, as will be detailed shortly. Local geothermal anomalies are reported over some of the Sirte basin fields at the present time (Ibrahim, 1995). This implies that fluids are still actively migrating within the basin. The sedimentary sequence, outlined in the earlier section of the chapter, showed how rapid lateral facies changes in the Late Cretaceous and Palaeocene demonstrate active faulting co-eval with rapid basin subsidence. From the Eocene onwards, however, the Sirte basin is characterised be a uniform stratigraphy, implying little local movement on the faults. Where depositional breaks do occur, as at the end of the Eocene, and the end of the Oligocene, they are widespread and only disconformable. This indicates a uniformly horizontal basin floor, with negligible topographic expression over the basement faults. The

Tertiary formations, especially the younger ones, crop out in a series of southerly facing escarpments concentric with the present day coast of the Gulf of Sirte. Allowing for subsequent erosion, this implies that successive marine transgressions became less and less extensive. This can be interpreted as either a progressive uplift of the Sirte basin, or a gradual global drop in sea level. The faults were not completely quiescent, however, and there is evidence for an episode of inversion during the Miocene (Selley, 1971). Figure 1 shows how oil fields lie in trends along the crests of basement fault blocks. One major trend is the Zelten-Waha ridge, a second, less obvious one, is the Bazuzi high on which lie the Reguba-Samah trend of fields. The Miocene shoreline mentioned earlier was locally interrupted by two rivers that scoured es-

34

R.C. SELLEY

Fig. 10. Series of cross-sections along the line C - D in Fig. 7, to show the structural evolution of the Sirte basin. Note the evidence for Miocene inversion (developed from Sclley, 1971).

tuarine channels and infilled them with some 200 m of sand. These rivers flowed south to north from the Sahara platform towards the Mediterranean. Their locations coincide with the axes of the two basement structures (Fig. 10). This clearly demonstrates a phase of Miocene structural inversion within the Sirte basin (Selley, 1971). It is not clear, however, just how widespread this inversion actually was. The Miocene estuarine channels are overlain by horizontal younger strata, indicating a return to quiescence, that continued with little interruption until the present day. The only dramatic exception to this statement is to be found on the western side of the Sirte basin, where the Hon Graben contains a thick fill of Pliocene to Recent sediment.

IGNEOUS ACTIVITY

Rifts, and failed rift basins, are normally accompanied by volcanic activity, and in this the Sirte basin was no exception. Basalts are encountered in the subsurface, and in at least one instance overly commercial accumulations of petroleum. The surface manifestation of igneous activity is much more dramatic, however, ranging from isolated volcanic craters to vast plains of basalt lava (Fig. 11 ). The distribution of the volcanic centres correlates closely with regional tectonic structures of North Africa, tending to occur where the Sirte rift faults intersect older, Palaeozoic and even Precambrian lines of crustal weakness (Vail, 1971). The main volcanic centres lie along the western flank of the Sirte basin. From northwest to southeast these are the Gharian, Jebel es Soda and Haruj al Aswad complexes. They have been described by

THE SIRTE BASIN OF LIBYA

35

Tripoli

\

Gulf of Sirte

Hon graber

/'~

\

\

,

it

Jeb~

Sirte \ basin \

es soda '1

-

/)

Harouj al Aswar

c

\)

(

/ (Namus /~eanw~

Algeria ' ////,

/

~

t

500 km

I

I

~-

\ ]

Jebel, 1 Eghei\ "I

Jebel Tibesti

0

basin, just beyond the limits of petroleum prospectivity. The Haruj al Aswad, however, covers large areas of the Zella (Tagrifet) trough. There are two other areas of volcanic activity that deserve mention, though they lie far to the south beyond the boundary of the Sirte basin. These are at Jebel Eghei, on the west flank of the Kufra basin, and in the Tibesti Mountains. Though beyond the limit of the Sirte basin they are clearly genetically related to its formation, bearing testimony to the collapse of the Tibesti-Sirte arch that gave it birth.

/

Tunisia .f ~

I l--

I I r__...J I I

Tertiary_Quaternary I volcanics .-~

Fig. 11. Map to show the extent of Neogene-Recent volcanic activity associated with the Sirte basin and its environs.

Busrewil and Wadsworth (1980a, b) and Woller and Fediuk (1980) respectively. Surface volcanic activity is largely Neogene to Recent in age. On the flanks of the Haruj al Aswad complex plateau basalts can be seen to have flowed down wadis cut in to Palaeocene limestones (Fig. 12) show them to be Pliocene in age. The fresh state of volcanic cones in both the Gharian dome and of the isolated volcanic vent of Waw en Namus (Fig. 13), suggest a very young, post-Pleistocene age (Pesce, 1966). The Gharian dome and the Jebel es Soda lie on the western edge of the Sirte

SELECTED BIBLIOGRAPHY Despite its vast petroleum reserves relatively little has been published on the Sirte basin. Important source books include: Gray, C. 1971. Symposium on the Geology of Libya. University of Libya, Tripoli, 520 pp. Salem, M.J. and Busrewil, M.T. (Editors), 1980. Geology of Libya, Vols. I-III. Academic Press, London, pp. 1-1712. Salem, M.J. and Belaid, M.N. (Editors), 1991. Geology of Libya, Vols. IV-V. Elsevier, Amsterdam, pp. 1713-2095. Salem, M.J., Sbeta, A.M. and Bakbak, M.R. (Editors), 1991. Geology of Libya, Vol. VI. Elsevier, Amsterdam, pp. 20992491. Also the field trip guide books of the Petroleum Exploration Society of Libya.

Note added in proof But note the publication of a 3-volume work on the Sirtc Basin: Salem, M.J., A.J. Mouzoughi and O.S. Hammuda (Editors), 1996. The Geology of the Sirte Basin, Volume I. Elsevier, Amsterdam, 564 pp.

Fig. 12. Weathered post-Palaeocene basalt lava flow on the eastern edge of the Harouj al Aswad.

36

R.C. S E L L E Y

Fig. 13. The caldera and inner cone of the Waw an Namus Quaternary volcanic crater. For location see Fig. 11. Salem, M.J., A.S. EI-Hawat, and A.M. Sbeta (Editors), 1996. The Geology of the Sirte Basin, Volume II. Elsevier, Amsterdam, 578 pp. Salem, M.J. Busrewil, M.T., A.A. Misallati and M. Sola (Editors), 1996. The Geology of the Sirte Basin, Volume III. Elsevier, Amsterdam, 380 pp.

REFERENCES Barr, E.T. and Walker, B.R., 1972. Late Tertiary channel system in northern Libya and its implication on Mediterranean Sea level changes. In: W.B.F. Ryan, K.J. Hsti et al. (Editors), Initial Reports of the Deep Sea Drilling Project, Washington, DC, XIII, pp. 1244-1251. Barr, F.T. and Weeger, A.A., 1972. Stratigraphic nomenclature of the Sirte basin, Libya. Petrol. Explor. Soc. Libya, Tripoli, 179 PP. Bebout, D.G. and Pendexter, C., 1975. Secondary carbonate porosity as related to early Tertiary depositional facies, Zelten Field, Libya. Bull. Am. Assoc. Petrol. Geol., 59: 665-693. Brady, T.J., Campbell, N.D.H. and Maher, C.E., 1980. Intisar "D" oil field, Libya. In: M.T. Halbouty (Editor), Oil and Gas Fields of the Decade 1968-78. Ame Assoc. Petrol. Geol. Mem., 30: 543-564. Brennan, P., 1992. Raguba Field. In: N.H. Foster and E.A. Beaumont (Editors), Structural Traps VII. American Association of Petroleum Geologists, Tulsa, OK, pp. 267-90. Busrewil, M.T. and Wadsworth, M.J., 1980a. Preliminary chemical data on the volcanic rocks of A1 Haruj area, Central Libya. In: M.J. Salem and M.T. Busrewil (Editors), Geology of Libya, Vol. III. Academic Press, London, pp. 1077-1080. Busrewil, M.T. and Wadsworth, M.J., 1980b. The basanitic volcanoes of the Gharyan area, NW Libya. In: M.J. Salem and M.T. Busrewil (Editors), Geology of Libya, Vol. III, Academic Press, London, pp. 1095-1106. Clifford, A.C., 1986. African oil - - past, present, and future. In: M.T. Halbouty (Editor), Future Petroleum Provinces of the World. Mem. Am. Assoc. Petrol. Geol., 40: 339-372. Clifford, H.J., Grund, R. and Musrati, H., 1980. Geology of a

stratigraphic giant: Messla oil field, Libya. In: M.T. Halbouty (Editor), Oil and Gas Fields of the Decade 1968-78. Am. Assoc. Petrol. Geol. Mem., 30: 507-524. Colley, B.B., 1963. Libya: petroleum geology and development. 6th World Petrol Cong., Frankfurt, Sect.l, Pap. 43, 10 pp. EI-Hawat, A.S., 1980. Carbonate-terrigenous cyclic sedimentation and palaeogeography of the Marada Formation (middle Miocene) Sirt basin. In: M.J. Salem and M.T. Busrewil (Editors), Geology of Libya, Vol. II. Academic Press, London, pp. 427-448. Fairhead, J.D. and Binks, R.M., 1991. Differential opening of the Central and South Atlantic Oceans and the opening of the West African rift system. Tectonophysics, 187: 191-203. Garea, B.B. and Braithwaite, C.J.R., 1996. Geochemistry, isotopic composition and origin of the Beda dolomite, Block NC74F, SW Sirte Basin, Libya. J. Pet. Geol., 19: 289-304. Goudarzi, G.H., 1980. Structure m Libya. In: M.J. Salem, M.T. Busrewil and A.M. Ben Ashour (Editors), The Geology of Libya, Vol. III. Academic Press, London, pp. 879-892. Goudarzi, G. and Smith, J.P., 1978. Preliminary structure contour map of the Libyan Arab Republic and adjacebt areas. U.S. Geol. Surv. Misc. Geol. Invest., Map 1-350C. Gumati, Y.D. and Kanes, W.H., 1985. Early Tertiary subsidence and sedimentary facies, northern Sirte Basin, Libya. Bull. Am. Assoc. Petrol. Geol., 69: 39-52. Gumati, Y.D. and Nairn, A.E.M., 1991. Tectonic subsidence of the Sirte basin, Libya. J. Pet. Geol., 14: 93-102. Gumati, Y.D. and Schamel, S., 1988. Thermal maturation history of the Sirte basin, Libya. J. Pet. Geol., 11: 205-217. Gumati, Y.D., Kanes, W.H. and Schamel, S., 1996. An evolution of the hydrocarbon potential of the sedimentary basins of Libya. J. Pet. Geol., 19:95-112. Hamyouni, E., 1984. Source and habitat of oil in Libyan Sirte basins. In: Habitat of Petroleum in the Arab Countries. OPEC, Kuwait, pp. 125-180. Hea, J.P., 1971. Petrography of Paleozoic-Mesozoic sandstones of the southern Sirte basin, Libya. In: C. Gray (Editor), Symposium on the Geology of Libya. University of Libya, Tripoli, pp. 107-126. Ibrahim, M.W.I., 1995. Geothermal gradient anomalies of hydrocarbon entrapment in the Middle East and North Africa.

T H E SIRTE BASIN OF LIBYA Megerisi, M.F. and Mamgain, V.D., 1980. The Upper-Cretaceous-Tertiary formations of northern Libya: a synthesis. Dept. Geological Researches and Mining, Bull. 12, 85 pp. Parsons, M.G., Zagaar, A.M. and Curry, J.J., 1979. Hydrocarbon occurrences in the Sirte basin, Libya. In: A. Miall (Editor), Facts and Principles of World Petroleum Occurences. Can. Soc. Petrol. Geol. Mem., 6: 723-732. Pavoni, N., 1993. Pattern of mantle convection and Pangea break-up as revealed by the evolution of the African plate. J. Geol. Soc. Lond., 150: 953-964. Pesce, A., 1966. Uaw en Namus. In: J.J. Williams (Editor), South-Central Libya and Northern Chad. Petroleum Exploration Society of Libya, Tripoli, pp. 47-52. Roberts, J.M., 1970. Amal Field, Libya. In: M.T. Halbouty (Editor), Geology of Giant Petroleum Fields. Am. Assoc. Petrol. Geol. Mem., 14: 438-448. Sanford, R.M., 1970. Sarir oil field - - desert surprise. In: M.T.Halbouty (Editor), Geology of Giant Petroleum Fields. Am. Assoc. Petrol. Geol. Mem., 14: 449-476. Selley, R.C., 1969. Nearshore marine and continental sediments of the Sirte basin, Libya. Quart. J. Geol. Soc. Lond., 124: 419-460. Selley, R.C., 1971. Structural control on Miocene Sedimentation of the Sirte basin, Libya. In: C. Gray (Editor), Symp. on the Geology of Libya. University of Libya, Tripoli, pp. 99-106.

37 Terry, C.E. and Williams, J.J.1969. The Idris "A" bioherm and oil field, Sirte Basin, Libya. In: E Hepple (Editor), The Exploration for Petroleum in Europe and North Africa. Inst. Pet. London, pp. 31-48. Van Houten, F.B., 1983. Sirte Basin, north-central Libya: Cretaceous rifting above a fixed mantle hot spot? Geology, 11: 115-118. Vail, J.R., 1971. Dike swarms and volcanic activity in northeastern Africa. In: C. Gray (Editor), Symp. on the Geology of Libya. University of Libya, Tripoli, pp. 341-347. Van Houten, F.B., 1983. Sirte Basin, north-central Libya: Cretaceous rifting above afixed mantle hot spot? Geology, 11: 115-118. Williams, J.J., 1968. The stratigraphy and igneous reservoirs of the Augila field, Libya. In: T.E Barr (Editor), Geology and Archaeology of Northern Cyrenaica, Libya. Petroleum Exploration Society of Libya, pp. 197-206. Williams, J.J., 1972. Augila Field, Libya, depositional environment and diagenesis of sedimentary reservoir and description of igneous reservoir. In: R.E. King (Editor), Stratigraphic Oil and Gas Fields. Am. Assoc. Petrol. Geol. Mem., 16: 623-632. Woller, E and Fediuk, F., 1980. Volcanic rocks of Jebel as Sawda'. In: M.J. Salem and M.T. Busrewil (Editors). Geology of Libya, Vol. III. Academic Press, London, pp. 1081-1094.

Chapter

4

Sedimentary Basins of Egypt" An Overview of Dynamic Stratigraphy

AHMED S. EL HAWAT

Initially, the questions may be elementary as 'What?' and 'Where?' but soon the 'How?' and 'When?' of process, history and geometry take center stage . . . . one of the most rewording questions that can be asked is 'So w h a t ? ' . . , or 'Let us see what it means'

R.N. Ginsburg (1982)- Seeking Answers

INTRODUCTION Information and literature on the geology of Egypt has been gathered and published by Said (1962, 1971, 1990), summarized by E1-Shazly (1977) and an updated annotated geological bibliography was published by E1-Baz (1984). The latest volume on the Geology of Egypt edited by Said (1990) is the most comprehensive reference on the subject to date. During the last few decades, exciting new research has been done by geoscientists from Egyptian and international academic institutions, oil companies, and government departments, such as the Geological Survey of Egypt. These publications have contributed a great deal towards our current understanding of the sedimentary basins of Egypt. In my view, some of the most significant contributions to the geology of Egypt were based on the application of the concepts of plate tectonics, global sea level changes, stratigraphy and sedimentology. The solution of the Nubian Sandstone problem in North Africa and the Middle East (Klitzsch et al., 1979; Klitzsch, 1990a; Klitzsch and Squyres, 1990; Van Houten, 1980; Van Houten et al., 1984, to mention just a few) is a case of a point. Because of the volume and the detailed nature of published material on the Egyptian geology, the present work attempts to retrace basinal development in Egypt, and concisely present it in the light of dynamic regional and global events through time. Since Libya and Tunisia were influenced by the same events as Egypt a conscious attempt is made

in this paper to extend synthesis across into these areas when possible. Stratigraphic subdivisions and nomenclature are often confusing and names are therefore not the objective in this paper. They are used here as a reference to other workers in Egypt. Stratigraphic sequences are treated instead in terms of time, space, events, cycles and facies. The first part of this paper deals with an overview of the Proterozoic evolution of the Afro-Nubian craton and the development of Phanerozoic sedimentary basins. The following sections trace basinal development through the Palaeozoic, Mesozoic and Tertiary times. Each of these sections is concluded with a subsection highlighting the relationship between sedimentation and significant geologic events which have taken place during that particular time. The last section is a summary and synthesis of common themes of the geologic history of the depositional basins of Egypt.

SEDIMENTARY BASINS T h e A f r o - N u b i a n craton

The basement complex of Egypt is exposed north and west of the Red Sea and the Gulf of Suez, in the areas of Sinai and the Eastern Desert. It extends south of latitude 24 ~ north of Jabal Uweinat and Aswan. Basement exposures represent about 10% of the total area of Egypt, the rest are covered by Phanerozoic sediments which increase progressively in thickness northwards following the regional dip

edited by R.C. Selley (Series Editor: K.J. Hsti), pp. 39-85. 9 1997 Elsevier Science B.V., Amsterdam. All rights reserved.

African Basins. Sedimentary Basins of the World, 3

40 of African plate towards the Mediterranean Sea. Irregular thickness distribution of the Phanerozoic sedimentary cover and basin development were influenced by basement structures and tectonic history since the Precambrian. The basement complex of Egypt was recently reviewed by E1 Gaby et al. (1990), Richter and Schandelmeier (1990) and Hassan and Hashad (1990). In their study of the structural development of the basement complex of the northeast African plate, Schandelmeier et al. (1987) have subdivided the basement into two distinctive parts, the pre-PanAfrican eastern Saharan craton and the Pan-African Nubian shield (Fig. 1). The eastern Saharan craton is located west of the river Nile. The oldest of these pre-Pan-African rocks are exposed in Jabal Uwainat which is located southwest of Egypt and southeastern Libya. These are dominated by granulite metamorphic rocks dated as Late Archean (2673 Ma) in age (Klerx, 1980). These rocks may represent a protocrust which was developed as a result of compressive tectonics during continent to continent collision (Morgan, 1990). Outwards from the Archean nucleus the craton consists of younger rocks that range from Early to Middle Proterozoic in age (Richter and Schandelmeier, 1990). These polymetamorphic and granitoid rocks are 2300 to 1800m years old and are arranged in a regional NW-SW, N-S, and E - W trending metamorphic belt extending throughout northeast Africa (Schandelmeier et al., 1987). These authors recognized three stages of deformation that include: (1) initial folding (2100 Ma), (2) development of transcontinental shear zones and sigmoidal bending of metamorphic foliation (2000 Ma), and (3) the development of Jebel Uweinat-Jebel Kamil, 55 ~ dextral wrench faults, and 150~ trending structural lineaments (1800 Ma). During Late Proterozoic, a second deformational event known as the Pan-African thermo-tectonic event took place. It resulted in thermal expansion and uplift of the basement, and in accretion of the oceanic crust assemblage of the Nubian shield to the eastern margin of the Saharan craton. Extensive studies were published on the Nubian shield exposure in Sinai and the Eastern Desert east of the Nile (El Gaby et al., 1990; Morgan 1990). Five main evolutionary development stages representing a complex history of crustal subduction of arc-trench system leading to accretion onto the east African continental nucleus are recognized (El Shazly, 1977). These stages are: (1) Eugeosynclinal flysch sedimentation associated with island arc volcanism and crustal subduction (1195-856 Ma). These sediments and associated ophiolites were regionally metamorphosed and folded by later orogenic events. (2) Orogenic and syn-orogenic intrusion of granite

A.S. EL HAWAT and granodiorite plutonic bodies. These were later partly metamorphosed into gneisses (890-876 Ma). (3) Post-orogenic geosynclinal volcanic phase (665654 Ma). (4) Development of post orogenic foreland basins and molass sedimentation. (5) Late orogenic and post orogenic plutonic intrusions of granites and granodiorites (656-480 Ma). It was noted that in the high mountain areas, the Nubian craton are formed of orogenic granites and develop a 30 to 40 km thick crust. Whereas the crust at the margin of the Red Sea trough is 20 km thick and consists of alkaline metasomatic granites belonging to the latest plutonic intrusions (EI-Shazly, 1977). The crustal thickness of the east Saharan craton also, exhibits similar order of magnitude from Jabal Uweinat area to the Mediterranean coast. The Proterozoic Pan-African crustal accretion and subsequent continent to continent collision was associated with major development of transcontinental shear zones caused by strike-slip motion parallel to the west African continental nucleus (Morgan, 1990, fig. 7.4). Reactivation of these shear zones and lineaments had a significant influence on the development and evolution of sedimentary basins during the Phanerozoic. In contrast to the spectacular Proterozoic events, tectonics and sedimentation during the Phanerozoic were relatively placid. During the Phanerozoic, events were consistent with the cratonal evolutionary process. Events were often initiated by uplifting of cratonic areas due to the thermal expansion resulting from the development of hot spot anomalies beneath the craton. Sedimentary basins were developed following crustal attenuation and fracturing of uplifted areas and subsequent faulting and subsidence. Basinal filling often produce symmetrical or asymmetrical depositional cycles which usually starts with basal clastics followed by shallow marine clastics and carbonate and, in some cases, evaporites. These depositional cycles reflect stages of basinal subsidence and subsequent marine transgression. This is a common theme of tectonics and sedimentation in the Egyptian basins. Based on their individual tectonic and depositional history a variation on the theme may occur. Sedimentary basins of Egypt are either intracratonic basins, pericratonic or rifts basins. Intracratonic basins

The Dakhla and Upper Nile basins are broad intracratonic depocentres which were developed in southern and central Egypt during the Palaeozoic and Mesozoic (Fig. 2). They have evolved as a result of structural differentiation and subsidence of the rigid cratonic plate. Morgan (1990) suggested that subsidence in Dakhla and the adjacent Kufra basin in Libya was initiated in response to cooling and

SEDIMENTARY BASINS OF EGYPT: AN OVERVIEW OF DYNAMIC STRATIGRAPHY 26 ~

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thermal relaxation of the crust at the closing stages of the Pan-African tectono-thermal event. Others (Neev 1975, 1977; Keeley 1989)however, attribute the origin of these basins to the major trans-African shear along the "Pelusium Line".

The southern margin of these basins is marked by Precambrian crystalline basement high forming a line north of the Egyptian-Sudanese border (Fig. 1). Southwest of Dakhla basin, near the corner of the Egyptian-Sudanese-Libyan border point, the base-

42

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(-iooo) ISOPACH CONTOURS Fig. 2. Tectonic map showing intracratonic basins of Egypt and Northern Sudan (After Klitzsch, 1984).

ment around Jabal Uweinat mountains consists of Precambrian metamorphic and igneous rocks and younger volcanics. They are partly covered by up to 600m of Palaeozoic clastics that extend north and northeast to the Gilf Kebir area in Wadi Abel Malik and Abu Ras Plateau (Klitzsch, 1986). To the southeast of the upper Nile basin, and south of Aswan, the Precambrian outcrops exhibit WSW-ENE alignment. Towards the east, it joins the NW-SE trending Rayan swell that constitutes the Eastern Desert Precambrian basement, and separate the Upper Nile basin from the Red Sea graben. The Dakhla and Upper Nile basins are separated from each other by the Kharga and Dakhla basement uplifts which are aligned in N-S and WNW-ESE directions respectively. During Late Cretaceous, the Dakhla and Upper Nile basins were joined as a result of relief inversion and subsidence of the Kharga high. To the north, these intracratonic basins are separated from the northern pericratonic marginal basin belt by the occurrence of the extension of Kattania and Farafra high. These structures are recognized in the subsurface from exploratory wells and geophysical surveys. During the Palaeozoic, the Dakhla

basin was separated from its western counterpart, the Kufra basin of Libya, by NNW-SSE trending Uweinate-Gardeba-Sirt high. However, during the Mesozoic this high was cut by the N W - S W trending fault system that joined the two basins along a common trough (Fig. 2). Klitzsch (1986, 1990a) and Klitzsch and Squyres (1990) recognized three major sedimentary cycles filling the intracratonic basin of Egypt and northern Sudan (Fig. 2). These cycles correspond to recognizable tectonic events in the NE African craton. The basal cycle was developed during Early Palaeozoic and was terminated at the end of the Visean. This cycle is dominated by clastic fluvial sediments, alternating with marine and marginal marine clastics associated with transgressive events. These transgressions took place during the Early Cambrian, Early and/or Middle Ordovician, Early Silurian and Early Carboniferous. Uplifting of central and southern Egypt during Late Carboniferous to Early Jurassic led to erosion of Early Palaeozoic deposits. In the Upper Nile basin, Early Palaeozoic deposits were found only at the centre of the basin in the area between Wadi Qena and Wadi

SEDIMENTARY BASINS OF EGYPT: AN OVERVIEW OF DYNAMIC STRATIGRAPHY Dakhla (Klitzsch, 1986, fig. 4). Because of the structural set up, fluvial transport was directed mainly southwards towards the Sudan. Together with the Late Carboniferous glacial facies association found in southwest Egypt and northwest Sudan, they are assigned to the second depositional cycle referred to as the Karoo cycle (Klitzsch, 1986). This cycle was followed by a regressive phase interrupted by a single transgressive event during the Aptian. In turn, this sequence is overlain by transgressive phase which was interrupted by regressive events during Turonian and Santonian times. These phases belong to the Early and Late Cretaceous respectively, which constituted the Nubian depositional cycle (Klitzsch, 1986; Klitzsch and Squyres, 1990). Pericratonic basins The pericratonic basins were developed on the northern continental margin of the Afro-Nubian craton following the opening of the Tethys ocean during the Middle Jurassic time. They occur on an attenuated continental crust consisting of structural system which run parallel to the northeast African coast of Egypt and northeast Libya (El Hawat and Shelmani, 1993). As a part of this system, these basins are bound to the north by a subsurface structural high trending parallel to the coastline in the Mediterranean offshore (Fig. 1). They are also, separated from the intracratonic basins to the south by the subsurface Rayan-Nafsha and Farafra highs. The pericratonic sedimentary basins were generally affected by the same prevailing regional structural and tectonic conditions as the intracratonic basins. However, being located on the margin of the craton, they were more susceptible to tectonic influences induced by collision with the Eurasian plate than their southern counterparts on the stable craton. The subsurface basement highs and basins of the pericratonic area were developed as a result of folding and associated faulting in two major directions. The older E S E - W N W trending Palaeozoic structures are intersected by younger Mesozoic ENEWSW trending set. These basins and highs exhibit considerable changes in shape and size through time. It should be noted that because of the changes of centres of sedimentation some authors used different names for the same basin at specific times (e.g. Keeley, 1989). These basins include Siwa, Matruh, Abu Gharadig, Faiyum, lower Nile or Gindy, and the Nile Delta (Fig. 1). Siwa basin is one of the main Palaeozoic depocentres in the Western Desert of Egypt. The basin is bound to the east by Mamura-Farafra ridge, and extends westwards into Gaghboub basin across the border into Libya, where the basin attains greater depth (El Gazeery et al., 1975). The Palaeozoic sec-

43

tion encountered in Faghur #1 Well, in the vicinity of Siwa oasis, is about 1700 m thick, but reaches a thickness of 2283 m in Gaghboub's Cori well G1-83. A hiatus at the top of the Palaeozoic marks the Hercynian tectonic movement. The Mesozoic section, on the other hand, is not significant in thickness in comparison to that of the Palaeozoic and Cenozoic. The Matruh basin is located northwest of the Western Desert and northeast of Siwa basin, it is bound to the south by the Qattarah ridge. The basin axis slopes to the northwest in the direction of the E - W trending Mediterranean offshore ridge (Fig. 1). Like A1 Jabal al Akhdar trough in Cyrenaica, subsidence in Matruh basin seems to have started during the Jurassic. During the Cretaceous, however, the basin was extended as far south as Qattarah depression and Jabal Agila as it was connected to Imbarka subbasin where 1830 m of sediments were accumulated. To the north subsidence along the basin's hinge belt accounts for the sudden increase of sediment thickness which reaches 3445 m (El Gazeery et al., 1975). One of the major sedimentary basins in Egypt that exhibits a great hydrocarbon potential is Abu Gharadig basin (Figs. 1 and 3). Currently, it produces 40,000 barrels of oil and 0.4 bcf of gas per day from Cretaceous reservoirs. The basin is located in the central part of the Western Desert. The basin was opened during the Early Cretaceous as a result of right-lateral movement of E-W to ENE trending normal faults, leading to the development of pull-apart grabens (Abdel Aal and Moustafa, 1988). This basin is also considered to be an extension of the E-W trending arm of the Sirt basin rift system (Guiraud and Maurin, 1992). It has evolved as a result of the development of wrench faulting and structural inversion related to the Late Cretaceous Syrian Arc fold system. The basin attains asymmetrical north-south cross-section with increased depth to the north. It is separated from the Matruh basin to the north by the Qattarah ridge, and from the Faiyum basin to the east and south by the Kattaniya high. The latter merges southwards into the Baharya basement platform. Abu Gharadig basin exhibits a general northeasterly trend with easterly tendency towards the centre (El Gazeery and Taha, 1971). The centre of the basin, whose depth is estimated to be 12,200m (Awad, 1984), is filled by more than 1300 m of Palaeozoic sediments above the basement complex. The basement is structurally higher at the southern border of the basin as it rises to a depth of about 1265 m (Awad, 1984). Generally, the Palaeozoic section decreases in thickness and pinches out south and east of the basin towards the Faiyum and Nile Delta basins due to the occurrence of a Palaeo-

44

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zoic basement high (Fig. 5), and to truncation by the Hercynian unconformity. The sequence is followed by the Jurassic which develop a single asymmetrical depositional cycle consisting of a basal sandstone facies grading upwards into carbonates, reflecting progressive marine influence. Laterally however, the western and southwestern portions of the basin are

dominated by continental clastic facies that changes into a thick sequence of marine carbonates to the northeast (Bayoumi et al., 1997). Sedimentation in Abu Gharadig basin during the Early Cretaceous was dominated by fluviatile and fluviomarine clastics. During the Late Cretaceous transgression, however, the basin was enlarged and

SEDIMENTARY BASINS OF EGYPT: AN OVERVIEW OF DYNAMIC STRATIGRAPHY extended over an extensive area at the centre of the Western Desert. Also, because of active basin subsidence history, it has received the thickest Late Cretaceous sequence throughout Egypt. The thick organic rich Senonian-Turonian carbonates and shales were considered to be the main hydrocarbon source rocks in the basin. On the other hand, the Cenomanian fluviatile and fluviomarine sands form the principle reservoir (Awad, 1984; E1 Gazeery and Taha, 1971). Bound by Abu Gharadig basin to the west and the Nile Delta basin to the east, E1 Faiyum basin was developed during the Jurassic as a northerly sloping basement. Like most pericratonic basins it was developed during the opening of the Tethys. Jurassic and Cretaceous depocentres were located between the Jabal Rissu-Katratania and Mubaraka area at the centre of the basin. However, this area was uplifted during the Late Cretaceous Syrian Arc tectonic activity. Consequently, the basin was effectively subdivided into two major northern and southern Palaeocene-Eocene depocentres, which correspond to Tiba and Gindy (Lower Nile) basins respectively (El Zarka, 1983). During the Palaeozoic, the N N W SSE trending axis of the Lower Nile basin was located between the Gulf of Suez and the Nile River due to the occurrence of a major high to the west (Fig. 6) that extended from Charge to Nafsha. The Palaeozoic is represented by 1520 m in Abu-Hammad well east of Cairo. Subsidence of the basin continued during the Jurassic, as a 2350 m thick sequence was deposited. During the Cretaceous the basinal axis was shifted in a northwest-southeast direction and the Lower Nile basin was connected to Faiyum basin to the north. However, the deepest area of sedimentation was established further south between Asyut and Minia. Following the Late Cretaceous Syrian Arc tectonics, Faiyum basin was divided and the newly developed Lower Nile (Gindy) basin and was separated from its northern counterpart Tiba basin. As the Lower Nile basin was established as the major Palaeogene depocentre in Egypt, the Nile Delta is recognized as the main Neogene depocentre in the Pericratonic area of Egypt. Due to its intermediate location, the Nile Delta area constituted a part of the Lower Nile basin during Palaeozoic, part of the Faiyum basin during Jurassic and Cretaceous, and was part of the Tiba basin during Palaeogene time. During the Neogene, the Nile delta basin was firmly established as a result of the development of two hinge faults arranged in a "V" shape with its head north of Cairo and its base opened towards the Mediterranean (Fig. 1). Subsidence of the delta east-west trending hinge belt started during the Middle Miocene due to sediment loading following the Palaeonile shift from

45

the west to its present day position (Salem, 1976). At mid-delta, the Late Cretaceous-Tertiary strata exhibit steep dips north of the hingeline. Younger Neogene marine and deep water deposits, on the other hand, are cut by a series of normal rotational fault system, taking a trend parallel to the Mediterranean continental margin. Two major coarsening-up depositional cycles belonging to the Miocene and Plio-Pleistocene followed by the Holocene constitute the Nile Delta sequence. Three major gas fields were discovered in the basin, including E1 Wastani, Abu Madi and Abu Qir gas fields. In these fields the organic rich open marine and basinal shales are recognized as the main hydrocarbon source rocks, whereas the top part of the early cycle, and the basal transgressive sands of the second cycle are recognized as forming the main reservoir rocks (Marzouk, 1974). The Gulf of Suez rift basin

Recent hydrocarbon exploration activities and research have led to extensive publications on the Gulf of Suez and the Red Sea rifts. These included comprehensive reviews, synthesis and original data on the dynamic tectonic evolution and sedimentation in the rift system (Phillobos and Purser, 1993; Bosworth, 1994; Patton et al., 1994; Moustafa and Khalil, 1995). The Gulf of Suez basin is a tensional tectonic rift that forms the northern extension of the Red Sea graben. The basin is 60 to 80 km wide and consists of two major tilted blocks found on each side of the rift (Thiebaud and Robson, 1979). The Gulf of Suez basin is separated from the Nile basin to the west by Um E1 Tenassib-Elba ridge of the Nubian plate and is bounded to the east by the Sinai basin of the Arabian plate. Deposition was maintained in the basin following the NNW-SSE trend of the basement structural lineaments since Palaeozoic (Said, 1962; Schandelmeier et al., 1987). The rift, however, is considered to be a postEocene feature that became a fully developed basin during the Miocene as it received more than 3700 m of sediments (Scott and Govean, 1985; Sellwood and Netherwood, 1984; Schandelmeier et al., 1987). The Gulf of Suez basin contains several sedimentary units that are regarded as good source rock for oil. These include Kareem and Rudeis shales of the Gharandal Group, the Belayim shales of Ras Malaab evaporites, and Esna Shale of the Palaeocene-Eocene sequence (Fig. 4). The Senonian and Eocene carbonates, however, are considered to be the main hydrocarbon source rocks in the Gulf of Suez basin. Oil was found in the Miocene, Cenomanian, Albian and Carboniferous sandstone reservoirs. Porous and fractured limestone and reefs of the Miocene, Eocene, and Late Cretaceous are

46

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also good reservoirs. The estimated oil reserves of the Gulf of Suez fields are amounted to 3800 million barrels.

THE PALAEOZOIC The Pan-African arc accretion and cratonization of the Nubian shield was followed by a period of extensive erosion and peneplanation resulting in

the development of a gently inclined land surface to the north and northwest. This surface formed a prominent unconformity over which the Palaeozoic clastic sequence was deposited. The unconformity surface exhibits strong erosional features, such as kaolinitization and soil development. The Early Palaeozoic sequence consists of an alternating siliciclastic-dominated succession of continental and marine stratigraphic units. The pattern of successive Palaeozoic marine incursions was

SEDIMENTARY BASINS OF EGYPT: AN OVERVIEW OF DYNAMIC STRATIGRAPHY 26 ~ 3

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controlled by the structural and palaeotopographic configuration of the craton which was inherited from the Pan-African event (Fig. 5). During the Palaeozoic times marine waters inundated Egypt in a southeasterly direction through the NNW-SSE trending Kufra-Darfur and Dakhla-Misaha troughs. These troughs were the main passageways through which the Early Palaeozoic seas extended as far south as northern Chad and the Sudan (Klitzsch, 1986; Schandelmeier et al., 1987). In the mean time, fluvial and fluvioglacial sediments were transported from the high areas from the south and east. The distribution and accumulation of continental and marine sediments were controlled by the position of the Pan-African highlands (Morgan, 1990). During the Palaeozoic transgression, the Egyptian craton was covered by a shallow to very shallow ma-

rine waters9 However, the Dakhla basin to the south and Siwa and Abu Gharadig basins to the north (Fig. 6) were the main foci of subsidence9 Palaeozoic sediment thickness in these depocentres reaches up to 3000 metres. Thinner sediment accumulations were found elsewhere in the cratonic and pericratonic areas in the Nile basin, the Delta, Gulf of Suez basin and Sinai. The age of the Palaeozoic sequence above the basement unconformity varies in different regions in Egypt, as elsewhere on the Nubian craton. This is attributed to marine transgression and continental sedimentation over structurally controlled and irregular palaeogeography, as much as the subsequent epeirogenic movements of the craton. Also, the paucity of time diagnostic fossils within the siliciclastic-dominated sequence led to difficulty in age determination, stratigraphic subdivision and cor-

48

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Fig. 6. Palaeozoic depocentres in Northern Egypt (after El Gezeery et al., 1972, 1975). relation. Abdallah et al. (1992) presented a review and discussion of nomenclature of the Palaeozoic stratigraphic units in the Eastern Desert of Egypt. By the uplift of central and southern Egypt at the end of the Visean during the Hercynian tectonics, the early Palaeozoic depositional regime was terminated. During this time a great part of the early Palaeozoic sequence was eroded, and with the destruction of the old structural configuration, fluvial drainage was reversed to the south (Klitzsch, 1983, 1986, 1990a, b). Cambrian

The Early Cambrian marine transgression in Egypt extended to latitude 27~ (Fig. 5). Marine Cambrian rocks found near Wadi Dakhla and Wadi Qena areas consist of fluvial and shallow marine arkosic sandstone and sandy shale (Klitzsch, 1986; fig. 4). At the northern edge of Dakhla Basin the Cambrian is 457 m thick marine sandstone and shale found in the Bahariya well. These beds contain brachiopods and trilobites remains of Early Cambrian age (Dakkak, 1988). Further north in Abu Gharadig basin, 443 metres of Cambrian sediments are recognized. In Siwa basin, the base of the Cambrian was reached in Fagur #1 Well. It consists of 851 metres thick sequence of quartz arenite and micaceous sandstone that grades upwards into shales. The top of the sequence is marked by a lithological break, where medium and fine sandstone with chert fragments

pass into micaceous shales. Fossils collected from cores of this well include linguloid brachiopods and trilobites of Late Cambrian age (Dakkak, 1988; Khalil et al., 1983). Elsewhere, the Cambrian (Araba Formation) was described from Araba-Druba area of southwestern Sinai (Said, 1971). It was later proven that Araba Formation and its lateral equivalents in the Eastern Desert to be Early Cambrian in age based on dating by trace fossils (Seilacher, 1990). The sequence is up to 120m thick and consists of basal lenticular conglomerates containing angular vein quartz and granitic pebbles, followed by a coarsening up cycle consisting of interbedded micaceous sandstone and variegated red and green sandy mudstone grading upwards into red-brown arkosic sandstone unit. These units are intensively bioturbated with abundant Skolithos burrows and trilobites traces of Cruziana (Said, 1971; Issawi and Jux, 1982; Abdallah et al., 1992). On the eastern side of the Gulf the Cambrian sequence is 100 metres thick. The lower part of the sequence consists of 1 to 5 m thick coarsening-up cycles of soft red and green mudstone, grading upwards into red-brown arkosic bioturbated sandstone. This unit exhibit desiccation cracks, wavy and lenticular bedding and low angle cross-lamination. Ichnofauna is dominated by Cruziana traces. The upper part of the sequence contains 10-15 m thick subunits of dark brown fissile, silty shale beds intercalated with coarse to very coarse often pebbly sandstone units exhibiting fining-up cycles with

SEDIMENTARY BASINS OF EGYPT: AN OVERVIEW OF DYNAMIC STRATIGRAPHY tabular, trough and herringbone cross-bedding showing extensive Skolithos and Bifungites burrows and some Cruziana traces. This facies extends into the Eastern Desert and to Wadi Qena area in the Nile basin (Klitzsch, 1986). The coarsening-up cyclicity of sand-shale units suggest deposition of sand bodies in near shore area; and the occurrence of interlaminated sands and shales indicate deposition in a shallow marine or protected lagoonal, tidal flat setting. The sequence may represent a prograding coastal plain sequence following the Early Cambrian an marine transgression (Bhattacharyya and Dunn, 1986; Abdallah et al., 1992). Ordovieian-Silurian Ordovician strata are relatively rare in Egypt. Ordovician outcrops, however, are found southwest of the Dakhla basin in Jabal Uweinat area unconformably underlain by crystalline basement. They consist of shallow marine sandstone beds containing abundant Cruziana rouaulti and Skolithos ichnofossils of Early Ordovician age (Klitzsch, 1986; Klitzsch and Lejal-Nicol, 1984). The Ordovician-Silurian deposits in the northern basins of the Western Desert were undifferentiated and unrecognized from the Early Palaeozoic sequence because of the paucity of index fossils. Therefore, some authors use the stratigraphic subdivision as those found in Ghadames and Murzuk basins of Libya because of lithological similarities (Khalil et al., 1983). Recently, north of Siwa Oasis in the Western Desert, the Silurian sequence was identified by means of palynomorphs in the subsurface of Basur #1 and Kohla #1 wells (Keeley, 1989). The Silurian forms a coarsening up regressive sequence bound by unconformities above and below. It consists of 626 metres of marine and marginal marine siltstone with minor amount of mudstone and sandstone grading upwards into 400 to 700 metres of braided stream and alluvial fan sandstone and conglomerate. In Foram #1 Well to the west and Sheiba #1 Well east of Qattarah Depression 200 to 300 metres of marine sandstone, siltstone and shale sequence was encountered. These rocks contain palynomorphs and ichnofossil Harlania dating the sequence as Silurian (Klitzsch, 1990b). In the Gilf Kebir area of southwestern Egypt, the Silurian forms the early 400 metres of the Palaeozoic sequence, where it rests unconformably over the basement. The sequence consists of thick bedded fluvio-deltaic and shallow marine sandstone interbedded with well bedded shale and siltstone. These strata exhibit trace fossils of Cruziana, Arthrophycus, and Skolithos burrows. This ichnofaunal assemblage is similar to that found in the Silurian sandstone of Libya. These deposits change

49

southwards into tillite consisting of white sandstone containing erratic quartz pebbles suggesting fluvioglacial origin (Beall and Squyres, 1980; Klitzsch, 1983). In the Gulf of Suez area the Ordovician-Silurian sequence, Naqus Formation (Said, 1971), is up to 462 metres thick. On the western side of the Gulf the sequence overlies the Cambro-Ordovician Araba Formation with an apparent gradational boundary. It is also, unconformably overlain by the Early Carboniferous Abu Thora Formation (Abdallah et al., 1992). The sequence is unfossiliferous and consists of fining upwards to noncyclic, often cross-bedded, well sorted coarse to medium grained, uniform feldspathic sandstone with few clay intervals. The sandstone contains occasional scattered vein-quartz pebbles and cobbles up to 10 cm in diameter. These pebbles in the sandstone exhibit a general northward decrease in size, as feldspar grains in the sandstone are kaolinitized giving the rock a speckled appearance. The basal unit of the sequence suggests emergence of the Cambrian marine sequence below. The top of the sequence, however, is marked by development of prominent mottled and kaolinitized soil horizon. The Ordovician-Silurian sandstone sequence consists of low-angle cross-bedded lenticular bodies which were deposited in northwesterly to north-northeasterly flowing braided streams on a prograding alluvial fan setting (Buhattacharyya and Dunn, 1986). Issawi and Jux (1982) interpreted this sequence as fluvio-glacial in origin. Devonian

In Siwa Basin of the Western Desert, the Devonian sequence was described from Zeitoun #1 Well. It forms a single, 288 metres thick, coarsening and shallowing up depositional cycle bounded by unconformities. The early part of the cycle contains interbeds of skeletal limestone grading up into pyritic claystone followed by sandstone and conglomerate at the top (McGarva, 1986). North of Siwa Basin, a more complex 166 metres thick sequence of the Late Devonian is reported. In this area the sequence consists of basal sandstone unit overlain by two marginal marine and deltaic coarsening-up cycles. Each of these cycle consists of argillite grading upwards into argillaceous sandstone and arenaceous sandstone. The sequence is terminated by fossiliferous shallow marine limestone. Late Devonian fossils found in the sequence include Theodessia aft. hungerfordi, Productella cf. hallina, Leptostraphia magnifica, Platyracheela cf. mesastrials and Fenestrillina aft. omaciata (Dakkak, 1988; fig. 7). The Devonian is found in Jabal Uweinat and Abu Ras Plateau area on the southwest margin of Dakhla

50

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basin (Klitzsch, 1983). It forms a 70 m thick succession overlaying the Ordovician-Silurian sequence. It consists of cross-bedded fluviatile sandstone which corresponds to the Early-Middle Devonian Tadrart sequence of Libya. Elsewhere, in Wadi Abdal Malik the sequence is 10-100 m thick shale and siltstone containing plant detritus. It is overlain by fine to medium-grained, bioturbated sandstone with Bifungites fezzanensis and Camerotoechia spp. trace fossils. It also contains brachiopods and pelecypods that indicate Devonian to Carboniferous age (Klitzsch, 1983).

Carboniferous The Carboniferous sequence in the subsurface of the Western Desert basins forms a symmetrical depositional cycle bounded by unconformities. It con-

sists of basal clastic unit followed by a unit of mixed carbonate and clastics, and is overlain by a clasticdominated unit at the top. These units are called Desouqy, Dhiffah and Sail formations respectively (Keeley, 1989). This cycle shows an increase in thickness from east to west as it develops 670 to 950 metres thick sequence towards Libya. The basal unit is 100300 metres thick. It consists of fining-up and deepening-up sequence of kaolinitic fluvial sandstone grading upwards into marginal marine, deltaic and prodeltaic siltstone and mudstone. The middle part of the Carboniferous cycle is 300 to 450 metres thick, and consists of oolitic and bioclastic limestone interbedded with mudstone and shale. As the carbonate facies decrease in thickness and abundance upwards, the sequence grades into a third unit at the upper part of the cycle. In Fagura #1 Well this unit forms a single coarsening-up argillite and arenaceous sandstone se-

SEDIMENTARY BASINS OF EGYPT: AN OVERVIEW OF DYNAMIC STRATIGRAPHY quence overlain by a carbonate unit. This sequence represents a prograding deltaic and near shore sequence followed by shallow marine, clastic free conditions suggesting marine transgression possibly related to subsidence of the delta. The occurrence of Polytaxis indicates Late Carboniferous age (Dakkak, 1988). The Carboniferous cycle may not be complete in some places due to erosion which have taken place during the Hercynian event. South of the Western Desert, in Wadi Abdal Malik, Klitzsch (1983, 1990b) reported the occurrence of Wadi Malik Formation. It is underlain by the Devonian sandstones of Tadrart Formation. Wadi Malik Formation develop 100 to 150 metres thick coarsening-up depositional cycle of alternating bioturbated silty shale, siltstone and fine grained sandstone grading upwards into cross-bedded coarse-grained sandstone and conglomerate. The basal units have yielded impressions of brachiopods, plant remains 2's

zb

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51

like Eremopteris whitei, and trace fossil such as Camerotoechia sp., Conostichus broadhedi, Asteriacite gugelhupf and Bifungites fezzanensis of Early Carboniferous age (Klitzsch, 1990b; Seilacher, 1983; 1990). The above sequence is disconformably overlain by 50 metres thick unit of Late Carboniferous ill-sorted tillite consisting of erratic blocks floating in mud matrix. The tillite grades to the south into sandstone containing erratic blocks and pebbles. Further south across the border in northern Sudan, the sandstone passes into several hundred metres thick sequence of varve deposits consisting of silty claystone and clayey siltstone with thin layers of fine grained sandstone containing trace fossils (Fig. 9). In Wadi Abdal Malik, however, the tillite sequence is overlain by sandstone and siltstone sequence rich in Carboniferous flora. The facies association described above represents a typical lateral change from glacial tillite into fluvioglacial and

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Fig. 8. Palaeography of Late Carboniferous to Early Jurassic and the structural impact of the Hercynian orogeny (compiled after Klitzsch and Wycisk, 1987; Schandelmeier et al., 1987).

52

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The Early Carboniferous section in the Gulf of Suez basin is 500 metres thick in the subsurface, and may reach a m a x i m u m of 80 metres in surface sections. It is assigned to U m m B o g m a Formation

SEDIMENTARY BASINS OF EGYPT: AN OVERVIEW OF DYNAMIC STRATIGRAPHY (Kostandi, 1959; Soliman and E1 Fetouh, 1970). It forms a transgressive sandstone sequence with basal conglomerate of well rounded vein-quartz pebbles and well sorted, well rounded sand grading upwards into ferruginous and magniferous claystone and shale. The top shale unit forms a regional marker horizon within the basin. It grades into dolomitic unit containing Early Carboniferous marine fauna. In turn, these carbonates grade laterally into multicoloured calcareous siltstone with poorly preserved brachiopods, and bryozoan shoreface sandstone indicating lagoonal bay or open marine conditions (Bhattacharyya et al., 1983). The Late Carboniferous Ataqa Formation in the Gulf area (Kostandy, 1959) consists of a single fining-up cycle of basal sand and shale facies. The lower sand facies consists of a series of fining upwards lenticular units with basal conglomerate and overlaying clays. These sands grade laterally northwards into thin-bedded, rippled and bioturbated sandstone that interfingers with fossiliferous marine shales. The sandstone, also, grades upwards into a coarsening-up sequence of interbedded sandstone and shale. These shales are either carbonaceous and contain plant remains such as Lepidodendron and Sigillaria (Weissbrod, 1969), or as in Umm Bogma area, it contains bituminous coal seams. Elsewhere, the late unit consists of green clays with well preserved brachiopods and bryozoa (Bhattacharyya et al., 1983). North of the Gulf of Suez area the cumulative thickness of the Early and Late Carboniferous sequence may reach up to 600 m of carbonates and marine shales. It is suggested that the Late Carboniferous sequence was deposited as a result of rapid progradation of fluviatile and fluviomarine facies, followed by gradual marine transgression leading to the deposition of shallow marine and intertidal sands and shales. The top of this sequence is marked by a regressive phase and the development of a soil horizon (Bhattacharyya et al., 1983). The Permian

Because of the effect of the Hercynian orogeny and the uplift of southern and central parts of Egypt (Klitzsch, 1983) there is no sedimentary record of the Permian in southern Egypt. However, north of the Western Desert in the Siwa Basin the Permian is represented in the subsurface by 70 m thick shallow water carbonates, sandstone and some shales. These sediments contain Early Permian fossils including Waagenocencha montepelierenses and Anisopyge cf. Prerassulata (Dakkak, 1988). In the Gulf of Suez basin the Permian and Triassic develop an undifferentiated fluvial sandstone sequence assigned to Quseib Formation (Abdallah et al., 1963). This section is discussed in the Triassic sequence.

53

Geologic events and sedimentation

Morgan (1990) noted that the Palaeozoic sequence in North Africa and Egypt is characterized by the relative paucity of carbonate sediments in comparison to siliciclastics when compared with other shelf areas of the world. He suggested that siliciclastic dominance over carbonate sedimentation in Egypt is attribute to sustained erosion of the PanAfrican mountains, epeirogenic movements of the craton and continental glaciation. Redistribution and reworking of these sediments, on the other hand, was attributed to the high eustatic sea level during the early Palaeozoic. Eustatic vs. tectonic control of sedimentation

Erosion, transportation, sedimentation and marine transgression during the Palaeozoic in Egypt was controlled by the development of wide grabens and smaller horst blocks following the NNW-SSE trend (Schandelmeier et al., 1987), and the occurrence of major Pan-African mountains to the south. Several depocentres were established due to postPan-African thermal relaxation and cooling and/or to shearing along major lineaments of the craton. The resulting troughs were submerged under marine conditions during the Ordovician, Silurian, Devonian and Early Carboniferous transgressions. Fluvial sediments were carried in the opposite direction by rivers after erosion from the Pan-African highlands. The resulting depositional cycle attain an asymmetrical motif consisting of basal fine-grained marine deposits grading upwards into coarse-grained progradational fluvial deposits. Otherwise, depositional cycles are symmetrical, as in the Carboniferous sequence in the pericratonic basins. In these depocentres the Carboniferous cycle begins and ends with coarse-grained continental clastics sandwiching a sequence of mixed siliciclastic-carbonate succession. These cycles are bounded by unconformities at the Ordovician-Silurian, Silurian-Devonian, DevonianCarboniferous and the end of the Permian. These erosional boundaries were caused by eustatic low stands associated with major tectonic movements of the craton. Bhattacharyya and Dunn (1986) demonstrated that marine transgression in the Gulf of Suez area during Cambrian and Early Carboniferous was interrupted by fluvial progradation during Ordovician-Silurian (Naqus Formation) and early Late Carboniferous (Ataqa Formation) times. These fluvial sandy progradation interrupted marine transgressions in different times and directions as a result of vertical block movement of the basement in the hinterland, along the old tectonic lineaments (Fig. 10). At the end of the Visean and as a result of the collision between Gondwana and the northern conti-

54

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The second most important reason for the dominance of siliciclastic sedimentation over carbonates in the Palaeozoic sequence of North Africa and Egypt is continental glaciation. Smith et al. (1981) suggested that during Late Ordovician-Early Silurian times the south pole was located in the central Saharan region. The Pan-African Hoggar mountains of Algeria and Tibesti and Jabal Uweinat mountains of Libya and Egypt were some of the main glacial centres throughout the Palaeozoic. Evidence of glaciation found in southern Algeria includes glacial moraines, fluvio-glacial and glacial-lacustrine deposits as well as glacial landforms exhibiting

glacially polished surfaces and striations. In the extreme southwest of Egypt, in Jabal Uweinat area, the occurrence of U-shaped valleys and rounded peaks in the crystalline basement are also, attributed to Early Palaeozoic glaciation sculpture (Bowen and Jux, 1987). The Ordovician-Silurian glaciation in southwestern Egypt was also, confirmed by the discovery of glacial tillite (Beall and Squyres, 1980). Similar glacial deposits and cold water fauna were also reported from Libya. Fluvio-glacial sedimentation in the Gulf of Suez area is indicated by the mineralogical and textural immaturity and the occurrence of erratic vein quartz pebbles and cobbles of Naqus Formation (Bowen and Jux, 1987). The tillite deposits found in Wadi Abdul Malik north of Jabal Uweinat in southwestern Egypt, on the other hand, were deposited during the Carboniferous glaciation (Figs. 8 and 9). Two hundred kilometres to the southeast across the border in the Sudan, these tillites change into several hundreds of metres of glacial-lacustrine varve sequence which are regarded by Klitzsch (1983) as equivalent to the Dwyka glaciation of southern and central Africa, where the pole was thought to be in the region of the Transvaal.

MESOZOIC The theme of tectonic and magmatic activities of the Hercynian orogeny continued unabated from

SEDIMENTARY BASINS OF EGYPT: AN OVERVIEW OF DYNAMIC STRATIGRAPHY the late Palaeozoic to early Mesozoic, leading to the breakup of the super continent of Pangea. During the Mesozoic several events of global magnitude have had a strong domino-effect tectonic influence on sedimentary basins of North Africa and Egypt. These included the opening of the mid-Atlantic ocean during the Early Jurassic, the opening of the Mediterranean Tethys during the Middle Jurassic; and the subsequent reversal of crustal movement leading to subduction of the Mediterranean oceanic crust due to collision between North Africa and Eurasia during the Late Cretaceous. These tectonics have led to movements of basement blocks in the cratonic and pericratonic areas of Egypt, shifts and changes of depocentres and inversion of subsiding basins into structural highs, or the reverse. The associated global eustatic events were also effective in influencing sedimentation in the North African and Egyptian basins. These include transgressions during the Middle Triassic, Middle and Late Jurassic, Aptian and Cenomanian. As we shall see later, recognition of these events by utilizing modern concepts of sedimentology and stratigraphy was used with great effect in solving endemic geological problems in North Africa and Egypt, such as the Nubian Sandstone problem.

Triassic The Late Carboniferous collision of Gondwana with the northern continents resulted in uplifting of central and southern Egypt and the development of an east-west trending basin to the north (Klitzsch, 1984; Klitzsch and Wycisk, 1987). The occurrence of Triassic deposits is restricted to northern Egypt, Sinai (Fig. 8), and northeast and northwest Libya (Shelmani et al. 1992, fig. 3). In these areas, deposition of the Triassic sequence was a result of gradual transgression, where depositional facies diachronously change both laterally northward and vertically across the Permo-Triassic time line, from continental sandstone and shales to marine limestones and marls. Most of the marine sediments were dated as Middle Triassic; the Permian if present, is mostly continental. South of Sinai and along the Gulf of Suez the Permo-Triassic sequence constitutes a fluvial succession. It consists of a fining-up cycles with basal lag of well-rounded vein-quartz pebble conglomerate, overlain by mid-channel trough cross-bedded, moderately well sorted sandstone and over-bank mudstone. Some cycles may exhibit well developed soil horizons at the top. Palaeocurrent measurements suggest NE and ENE flow direction (Darwish, 1992). Locally, greenish laminated kaolinite lenses up to 5 m thick occur within the sequence and are interpreted as lacustrine deposits (Bhattacharyya et

55

al., 1983). On the western side of the Gulf of Suez in E1-Galala E1-Bahariya Plateau, the Permo-Triassic is 80 m thick (Abdallah et al., 1963). It consists of ferruginous sandstone, siltstone and shale with minor gypsum and rock salt overlain by limestone and marl. The fluvial-dominated sediments of southern Sinai change northwards and upwards into a progressively marine siliciclastic and carbonate sequence of Middle Triassic age. In the subsurface (Halal #1 Well) of northeastern Sinai, the southern fluvialdominated facies change into marginal marine and deltaic facies complex interbedded with carbonates (Druckman, 1974). Southeast of this well in Arif E1 Naga anticline, 200 metres of Middle Triassic sequence are exposed. The sequence exhibits an upwards increase of marine influence, as the basal sandstone units grade upwards into carbonates and evaporites interbedded with algal stromatolites at the top. The lower units of the sequence consist of clean, multicoloured, coarse grained, trough cross-bedded sandstone containing vertebrate bone fragments, and variegated shale and siltstone with plant remains. Palaeocurrent analysis of the sandstone facies suggest a dominant north and northeast transport direction (Karcz and Zak, 1968). Up section the cross-bedded sandstone units are overlain by a succession of sandstone, shale, limestone and marl with Middle Triassic ammonite Ceralites (Awad, 1946). The carbonates consist of skeletal wackestone and grainstone grading upwards into ooskeletal packstone, mudstone and dolomites with algal stromatolites. These in turn grade into dolomitic shale and anhydrite (Jenkins, 1990). The sandstones and shale were interpreted as being deposited in a marginal marine estuarine-tidal fiat-beach complex. The limestone development in the sequence suggest increased carbonate productivity due to marine transgression and reduction of siliciclastic influx into the basin. The occurrence of evaporites algal stromatolites, and dolomitic shale at the top of the sequence, on the other hand, suggest terminal progradation of lagoonal and peritidal deposits over the high energy, marginal marine carbonate sand bodies.

Jurassic The east-west trending Hercynian uplift ceased to dominate central and southern Egypt in Early and Middle Jurassic time (Klitzsch, 1986; Schandelmeier et al., 1987). During Late Jurassic (145-132 Ma) reactivation of the Palaeozoic NNW-SSE trending basement configuration took place and dominated sedimentation throughout the Mesozoic. These were associated with continued extensional and strikeslip sinistral tectonics between Africa and Eurasia

56

A.S. E L HAWAT 26 ~

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SHORELINE

P A L E O C U R R E N T S OF EARLY CRETACEOUS R E G R E S S I V E DEPOSITS

A L L U V I A L DEPOSITS EARLY CRETACEOUS

Fig. 11. Palaeogeography of Late Jurassic-Early Cretaceous and Aptian transgression (compiled after Klitzsch and Wycisk, 1987; Van Houten et al., 1984).

(Smith, 1971), in conjunction with alkaline magmatism in Egypt (Meneisy, 1990). During the Mesozoic, the structural configuration and physiographic features of the basement of Egypt were not pronounced, allowing successive transgressions and regressions to take place over wide areas (Schandelmeier et al., 1987). These features extended westwards into Cyrenaica platform of northeastern Libya. North of this area, however, Cyrenaica trough was subsiding at a faster rate in response the extensional forces exerted during the opening of the Mediterranean Tethys. Sedimentation in the trough was dominated by bathyal debris flow deposits and turbidites through out the Jurassic and Early Cretaceous (Klitzsch, 1970; E1 Hawat and Shelmani, 1993). Further to the west, the initial

stages of breakdown and collapse of Sirt-Tibesti uplift was taking place and leading to the development of Sirt rift system during Late Jurassic. The marine Jurassic transgression in Egypt did not advance south further than latitude 29~ (Fig. 11). Jabal Maghara of northern Sinai offers the best Jurassic section in Egypt (A1 Far, 1966). The sequence has a maximum thickness of 1980 metres and consists of three major siliciclastic-carbonate cycles that extend from Early to Late Jurassic (Jenkins, 1990). The lower cycle is siliciclastic-dominated, the remaining cycles are arranged in an upwards increasing thickness and carbonate content. Each depositional cycle consists of basal sandstone, shale and occasional coal, overlain by limestone which marks the climax of marine transgressions during

SEDIMENTARY BASINS OF EGYPT: AN OVERVIEW OF DYNAMIC STRATIGRAPHY Early, Middle and Late Jurassic times respectively. These facies represent fluvio-deltaic, marsh and near shore clastic sedimentation, followed by open marine neritic depositional conditions associated with maximum transgression. The upper part of the Late Jurassic carbonate-dominated cycle (Masajid Formation) consists of coralline and stromatoporoid patch reefs associated with skeletal debris and oolitic shoal facies. Laterally, however, the Jurassic sequence forms a series of parallel facies belts arranged from south to north into siliciclastic fluvial and marginal marine facies, shallow marine shelf carbonates and shelf margin reefs and ooids shoal complex. These grade northwards in the Mediterranean offshore into deep marine carbonates (Jenkins, 1990, fig 19.7). Regardless of the difference in nomenclature used, the Jurassic section in the subsurface of the Western Desert maintains the same lateral northsouth facies pattern and vertical cyclicity as in Sinai. Hantar (1990), however, reported that the Early Jurassic cycle (Bahrein Fm.) in the Western Desert is dominated by continental clastics which passes laterally and upwards into marginal marine clastic (Khatatba Fm.) and followed by mixed clastic-carbonate facies (Wadi Athrun Fm.) of Middle Jurassic cycle. These change northwards and upwards into the Late Jurassic carbonate-dominated cycle (Masajid and Sidi Barrani Fms.). Elsewhere, in the Western Desert and west of the Nile Delta, the Faiuym Basin was the most prominent Jurassic basin on the unstable pericratonic area. In this basin Awad (1984) reported the occurrence of 2290 metres of Jurassic section in the subsurface. The Early Jurassic sequence consists of 360 m thick sandstone and shales of littoral affiliation followed by 1400 m thick shale sequence suggesting increased basinal subsidence during Middle Jurassic. The Late Jurassic, however, is characterized by a 500 metres thick shallow marine carbonate-dominated sequence. Further to the south on the stable craton, the Early and Middle Jurassic sections are missing due to the residual effect of the Hercynian structure and erosion. The Late Jurassic sequence, however, was found in a restricted area in the subsurface of Dakhla Basin (Wycisk, 1987). This sequence consists of 200 m thick fluviatile sandstone that passes laterally to the north into an alternation of sandstone and mudstone of shallow marine origin. The geographic distribution of sequence suggests deposition in a rapidly subsiding embayment with opening to the north along A1 Mesaha trough.

Cretaceous Early Cretaceous The Tethyan extensional tectonics and associated sinistral strike-slip movements which were initiated

57

during early Mesozoic continued to influence the North African continental margin during Early Cretaceous times. Sedimentary basins of the south were subjected to epeirogenic movements of fault blocks and maintained subsidence. These tectonic activities were responsible for sustaining regressive fluvial siliciclastic sediment influx in these basins. The balance of this regressive depositional style was briefly altered during a transgressive eustatic event during the Aptian. These tectonic activities were also responsible for the development of new depocentres and maintaining subsidence along the older ones on the unstable pericratonic margin area of northern Sinai, the Western Desert and northeast Libya. The new depocentres include Matruh and Abu Gharadig basins in Egypt (Fig. 12) and Sirt rift basin in Libya.

Neocomian-Barremian. South of the Western Desert, in the subsurface of the Dakhla basin, Wycisk (1987) recognized a 50-80 m section of fluvial deposits possibly of Neocomian age directly underlain by Late Jurassic sandstones, and is overlain by fluvial sandstones sequence of preAptian age. The Neocomian-Barremian sequence was designated as Six Hills Formation by Barthel and Bottcher (1978) for exposures on the southern margins of Dakhla and the Upper Nile basins. In southwestern Egypt the sequence consists of 100 to 500 m thick sandstones, unconformably underlain by peneplaned crystalline Precambrian basement or Palaeozoic rocks (Klitzsch et al., 1979). The Six Hills Formation forms a succession of fining-up cycles of ill sorted, medium to coarse grained sandstone, grading upwards into white, massive, mottled kaolinitic sandstone with well developed soil horizons showing root casts. The sandstone exhibits large scale tabular planar cross-bedding sets separated by conglomeratic layers of quartz pebbles. Dip direction of cross-bed foresets suggest northeast to northwest dispersal patterns. Fossils found in these rocks are mainly silicified wood fragments. These facies were deposited in fluvial-dominated, flood-plain environment in slightly sinuous channels (Klitzsch et al., 1979; Hendriks et al., 1984). Further to the south of Egypt, the base of the formation consists of conglomerates that, also, indicate deposition under terrestrial conditions. Sediments at the top portion of the sequence are generally different. They form a fining upwards sequence and exhibit a notable improvement of textural characteristics. These features, together with the occurrence of occasional burrows and Teredo borings in silicified wood, suggest marginal marine influence at the top of the sequence (Klitzsch et al., 1979; Hendriks et al., 1984). This formation is equivalent to the lower sandstone, fluvial-dominated, member of the

58

A.S. EL HAWAT

MEDITERRANEAN SEA

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tripartite subdivision of the Nubian (Sarir, Calanscio, Faragh) sequence of southeast Sirt basin of Libya (El Hawat, 1992; El Hawat et al., 1997). In the subsurface north of the Western Desert the Early Cretaceous sequence is referred to as Burg El Arab Formation. The lower part of formation consists of sandstone (Alam E1 Bueb) and shale (Matruh) members of Neocomian-Aptian age (Hantar, 1990). The sandstone forms a coarsening-up sequence as it grades up from shale units at the base. It is overlain by limestone units at the top. These marginal marine facies changes southwards into fluvial-dominated deposits of the Six Hills Formation and passes into carbonates to the north. The shale member is dark, calcareous, pyritic and contain lignitic layers suggesting deposition in marginal marine marsh conditions. These shale facies are limited in their distribution to Mersa Matruh area. Aptian. The Aptian was a time of transgression in Egypt (Fig. l 1) and Libya. The sequence of A1 Alamain Formation which was encountered in the subsurface of the northern Western Desert consists of two members. The lower clastic-dominated member develops a gradational relationship with the underlying Neocomian-Barremian clastics. It grades upwards into an upper, widespread carbonate-dominated member (Abdin and Deibis, 1972; E1-Zarka, 1983). Alamain Formation, however, does not extend further south than latitude 29~ which marks the limit of carbonate sedimentation during the Aptian transgression. This member forms

an important reservoir in A1 Alamain field and other oil fields in the Western Desert. It consists of lower limestone and upper dolomite units that are skeletal, oolitic and contain orbitulines, Bryozoa, echinoids and other fossils (Soliman and E1 Badry, 1970). Further south, these carbonates grade into the Aptian Abu Ballas Formation. This formation is clastic-dominated marine sequence which reaches a maximum thickness of 250 m at the centre of the Dakhla Basin. On the periphery of the basin, the Abu Ballas Formation is over 20 to 45 m thick and consists of five superimposed facies arranged in a single, coarsening-up, transgressive-regressive depositional cycle (Hendriks and Kallenback, 1986). The basal facies of this sequence is multicoloured, laminated, massive and bioturbated mudstone, rippled siltstone and cross-laminated fine grained sandstone. It contains brachiopod (Ligula sp.), pelecepods gastropods, and Rhizocorallium sp. burrows. Reworked mud conglomerate found in this facies suggests periods of high energy storm activities. The basal facies grades upwards into varicoloured siltstone and fine-grained sandstone exhibiting cross-lamination, ripple laminations and herring-bone cross-bedding in channels. Siltstone intercalations may contain plant fossils and exhibit desiccation cracks. The upper part of the sequence is coarsening-upwards sandy units alternating with tippled and bioturbated siltstone. These facies were deposited in a beachback shore environment which was possibly influenced by tidal and longshore currents (Hendriks and

SEDIMENTARY BASINS OF EGYPT: AN OVERVIEW OF DYNAMIC STRATIGRAPHY Kallenback, 1986). Abu Ballas Formation is disconformably overlain by fluviatile facies of the AlbianCenomanian, Sabaya Formation. Abu Ballas Formation is also exposed in the upper Nile basin to the east (Hendriks et al., 1984). The geographic distribution of the formation suggests, a shallow Aptian marine transgression covering most of Egypt, except for the basement high areas (Fig. 11). Abu Ballas facies also suggests that the southern edge of the Tethys reached a maximum water depth of about 20 metres (Bottcher, 1982). In Gilf Kebir area, southwest of Dakhla basin, the Gilf Kebir Formation, which rests unconformably on the crystalline basement (Fig. 13), consists of

f m c cor~g I I I I /

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59

70 m of cross-bedded sandstone deposits of fluvial, fluviomarine and coastal origin. These are laterally equivalent to Six Hills and Abu Ballas formations (Klitzsch and Wycisk, 1987). Abu Ballas Formation is equivalent to the middle sandstone member and its lateral variegates shale facies in the Nubian sequence of Sirt basin, Libya. The Aptian eustatic sea-level rise led to reworking and sedimentation of the middle sandstone member along E - W trending, fault controlled inlets which were cutting in N N W - S S E trending basement highs in the basin. The resulting lowering base level of equilibrium led to deposition of the variegated shales in the marginal troughs of the basin (El Hawat, 1992; E1 Hawat et al., 1997).

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Fig. 13. Lower-Upper Cretaceous sections in Gilf Kebir area SW Dakhla Basin (A) and the Upper Nile Basin south of Aswan (B) (compiled after Klitzsch and Wycisk, 1987; Van Houten et al., 1975).

60 Albian. Following the Aptian transgression, and before the onset of the Early Cenomanian transgression, continental sedimentation dominated the A1bian in Egypt. In the intracratonic basins of southern Egypt, however, continental sedimentation was extended to Early Cenomanian, as the Late Cretaceous transgression reached the Dakhla and upper Nile basins during the Late Cenomanian. Basaltic layers associated with these deposits give dates ranging from 84 to 100 Ma (El Shazly, 1977) and suggesting that sedimentation during the Albian was associated with intense tectonic activities and rejuvenation of fault movement. The Albian sequence encountered in the subsurface of the northern Western Desert develop a coarsening upwards cycle of sedimentation overlaying the Alamain carbonate unit. The cycle consists of basal greenish gray shale, siltstone and sandstone interbeds (Dahab mem.). These grade upwards into (Kharita mem.) fine to coarse grained sandstone, subordinate shale and carbonate interbeds that increase in thickness to the northwest (Hantar, 1990). To the south, on the eastern edge of the upper Nile basin, the Albian-Cenomanian sequence rests unconformably on deeply weathered Precambrian crystalline basement. It consists of 30-100 m thick, very coarse, pale yellowish to brownish gray, kaolinitic quartzose sandstone arranged in lenticular fining up units with basal lag of angular quartz pebbles (Ward and McDonald, 1979; Van Houten et al., 1984). The sandstone often grades upwards into paleosols of mottled kaolinitic sandstone showing vague roots and ferruginous nodules. Trough cross-bed sets indicate westward flow direction from the basement highland of the present day Red Sea (Fig. 11). Locally, however, cross-beds may indicate south to southwest flow directions due to the presence of local highs. Elsewhere, in the upper Nile and Dakhla basins palaeocurrents suggest northeast and northwest flow directions away from basement highs (Klitzsch et al., 1979). These sediments were deposited in low sinuosity braided stream systems associated with abundant development of paleosols (Ward and McDonald, 1979; Klitzsch et al., 1979; Van Houten et al., 1984). Similar facies belonging to the Sabaya Formation was described from other parts of the upper Nile (Hendriks et al., 1984) and the Dakhla basins, where it reaches a maximum thickness of 200 metres and rests disconformably on the Aptian, Abu Ballas Formation. All Authors emphasize the illsorted nature of the Albian-Early Cenomanian fluvial deposits and the occurrence of paleosols in southern Egypt. The diverse paleocurrents directions and the evidence of volcanic activities associated with this facies are indications of reactivation of basement highs surrounding the intracratonic basins of Dakhla

A.S. EL HAWAT and upper Nile (Van Houten et al., 1984). The upper sandstone member of the Nubian sequence of S irt basin is also regressive. They are associated with volcanic rocks and current transported pyroclastic deposits (El Hawat, 1992; E1 Hawat et al., 1997). In the Gulf of Suez basin, the Early Cretaceous Malha Formation which is exposed on the west side of the gulf is 50-100 metres thick and consists of continental to shallow marine sandstone and shale of Urgo-Aptian age (Abdallah et al., 1963; E1 Shazly, 1977). However, on the eastern side of the gulf this formation is thought to be of continental origin and of Albian age (Bhattacharyya and Dunn, 1986). In this area the Albian sequence consists of basal polymictic conglomerate containing angular to subrounded chert, jasper, volcanic rock fragments and vein quartz pebbles, associated with poorly sorted kaolinitic sandstone. It is thought that conglomerate clasts were derived from the nearby basement source as a result of uplift of the basement fault blocks (Van Houten et al., 1984). In the subsurface of central Sinai the upper portion of 520 metres thick sequence of sandstone interbedded with shale has yielded corals and ammonites of Albian age (Jenkins, 1990). This sequence passes further to the north of Sinai and the Mediterranean offshore into east-west trending shelf margin carbonates consisting of oolitic and bioclastic limestone facies grading into deep marine shale in the offshore area. Late Cretaceous The Late Cretaceous was a time of a major change in the depositional and tectonic history of the North African basins because of contemporaneous global tectonic and eustatic events. Basin development and sedimentation were influenced by three major factors, these are: (a) continued influence of the NNW-SSE basement structural trends; (b) the collision between the Africa and Eurasia; and (c) the global Late Cretaceous eustatic sea-level rise. The NNW-SSE basement structural elements continued to control basin development as during the Early Cretaceous. These structural elements were enhanced by increased differential movements and increased basinal subsidence during Late Cretaceous (Klitzsch, 1986). Meanwhile, the extensional tectonics and the associated sinistral strike-slip movement which was associated with opening of the Mediterranean Tethys during early Mesozoic was terminated. These tectonic movements were reversed and replaced during the Late Cretaceous-Palaeocene time by compressive and dextral shear movements contemporaneous with the opening of the north Atlantic (Smith, 1971). In Egypt, these tectonic reversals were associated with magmatism which were dated from the Albian to Campanian (100-80 Ma) and peaked during the Turonian (90 Ma) (Meneisy, 1990). These

SEDIMENTARY BASINS OF EGYPT: AN OVERVIEW OF DYNAMIC STRATIGRAPHY i

26 i

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i

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Fig. 14. Late Cretaceous depocentres in northern Egypt (after El Gezeery et al., 1972, 1975). tectonic changes have led to the subduction of the Tethys oceanic crust and the inversion of the subsiding pericratonic basins of Egypt into structural highs known as the Syrian Arc fold system. In Cyrenaica it was responsible for the inversion of the previously subsiding trough of A1Jabal al Akhdar into structural high (El Hawat and Shelmani, 1993). These new tectonic conditions have influenced basin orientation throughout northern Egypt during Late Cretaceous. The eastern basins in Sinai, the Gulf of Suez and the lower Nile (Gindy) basin were oriented in N N W - S S E direction following the trend of the rising basement high of the Red Sea. Basins of the Western Desert such as Abu Gharadig, Faiyum and other smaller basins were oriented in N E - S W direction (Fig. 14). Also, tectonic reactivation of faults, and the global marine transgressions were combined to produce a thick accumulation of Late Cretaceous sediments in most of these basins. The role of the Late Cretaceous eustatic sea-level rise was not only important in covering the Afro-Nubian craton by marine water. It has also altered the siliciclastic depositional system which was dominant since the Palaeozoic, into carbonate-dominated system. The latter system lasted until the Oligocene time when crustal tectonics reversed the depositional regime through out North Africa to siliciclastic sedimentation yet again. Cenomanian. The Cenomanian in Egypt was a time of major transgression (Fig. 5). Sedimentary records indicate that marine conditions has prevailed

in northern Egypt during Early Cenomanian, and reached the south during Late Cenomanian. The Cenomanian sequence north of the Western Desert exhibits a wide variation in thickness and lithology. Carbonates were dominant in the northeast, but changed westward into shale, and southward into sandy facies (Soliman and E1Badry, 1970). These deposits attain a maximum thickness in the subsurface in the depocentres of the Abu Gharadig (900 m), and Faiyum basins (1154 m). Such thickness, however, decrease to 140 metres over structural highs separating these basins (El Zarka, 1983; Awad, 1984). In surface exposures and subsurface sections the Cenomanian sequence forms a single clastic-carbonate cycle that suggests progressive increase of marine influence through time. The basal transgressive sandstone is coarse to fine grained calcareous and often glauconitic. In places, it is variegated, friable, cross-bedded and may contain fossils of Exogyra, Ostrea as well as vertebrate remains. It is often associated with carbonaceous, pyritic, fossiliferous shale. These sands and shales were deposited in estuarine and shallow marine environment (Allam, 1986). The clastic sequence grades upwards into oolitic, and bioclastic grainstone and mudstone sequence which is interbedded in places with anhydrite and chert (Soliman and E1 Badry, 1970). These suggests deposition in a shallow marine, near shore intertidal and supratidal setting. The Cenomanian in the Gulf of Suez basin, forms up to 300 m thick sequence at the centre of the present day Gulf, which coincides with a N W - S E trending embayment (Kostandi, 1959). In central

62 Sinai, the Cenomanian sequence consists of lagoonal to marginal marine shales and marls rich in oyster banks, that grade northwards in the subsurface into 326 metres of interbedded carbonates and shales. Further to the north of Sinai the sequence is dominated by dolomite and dolomitic limestone which was deposited over a very broad shelf of restricted circulation. These depositional areas change towards the Mediterranean offshore into deep marine shale of Albian to Santonian age (Jenkins, 1990). The Cenomanian sequences may exhibit a variable thicknesses through out Sinai due to subsequent tectonic movements and erosion during the Turonian. To the south of Egypt in Dakhla and Upper Nile basins, the Late Cenomanian marine sequence (Maghrabi Formation) rests on the underlying Albian-Cenomanian (Sabaya Formation) with an erosional surface marked by well developed paleosol (Hendriks, 1986). In Dakhla basin the sequence consists of massive claystone, alternations of mudstone, siltstone and fine-grained sandstone associated with occasional intraformational conglomerate, and cross-bedded sandstone filled channels. Sand, silt and shale are often arranged in fining-up, as well as, coarsening-up cycles. These exhibit bioturbation structures, and root casts, and are often associated with coal beds, vertebrate remains and brachiopods (Lingula sp.). Sedimentary structures of the sandstone filled channels indicate north to northeast sediment transport direction. The lateral and vertical arrangements of these facies suggest a prograding supratidal, intertidal and subtidal facies areas associated with an estuarine channel system (Hendriks, 1986). The sequence of Magharabi Formation also suggests that the Late Cenomanian transgression was followed by progradation and regression in southern Egypt, a possible prelude to the subsequent Turonian regression. Elsewhere, on the southern margin of the basin, tidal deposits are associated with small deltas, and muddy flood plains facies (Klitzsch, 1979). The Late Cretaceous transgression also reached the Upper Nile basin during the Late Cenomanian. In the vicinity of the city of Aswan the sequence is 30-40 m thick (Fig. 13). It pinches out towards the basin margin (Fig. 15), as it grades into fluvial deposits in the direction of basement highs (Van Houten et al., 1984; Ward and McDonald, 1979). The lower part of the sequence rests on the A1bian nonmarine sequence with a sharp to erosional basal contact, and forms succession of sandstone and mudstone exhibiting an upwards decrease in grain size. The sandstone is coarse to medium grained, planar cross-bedded and is extensively bioturbated. Palaeocurrent measurements suggests southeast and southwestern dispersal directions (Fig. 13B). The associated mudstone, on the other hand, is red to gray coloured and forms mottled paleosols containing fer-

A.S. EL HAWAT ruginous nodules. The upper portion of the sequence consists of a series of coarsening-up cycles of laminated claystone with plant remains, tippled silty clay and cross-bedded sandstone. Cycles are overlain by bioturbated ferruginous sandstone or oolitic chamosite and haematite. Locally, phosphate pebble conglomerate and Late Cretaceous shells of Inoceramus are found. Van Houten and others (1984) documented a reversal of palaeocurrent dispersal patterns in this part of the sequence, to northeast and northwest (Fig. 13). They concluded that following the initial transgression, marshes and coastal plain depositional conditions had prevailed. These were followed by progradation of oolitic ironstone-capped sand bodies which were developed in the marine embayment of Aswan during the peak of transgression (Fig. 9). The Cenomanian fossils and sediments indicate warm, humid to semi-humid climate, low tidal range and weak wave energy. These conditions were suitable for lateritic weathering and development of iron rich deposits (Hendriks, 1986).

Turonian-Coniacian. The Turonian-Coniacian sequence in Egypt is represented by a regressivetransgressive subcycle within the overwhelming Late Cretaceous transgression. The Turonian regression was a result of global tectonic events associated with volcanic activities and intrusions of alkaline volcanic rocks, the peak of which was centred at 90 Ma in Egypt (Meneisy, 1990). During this time, Abu Gharadig and Matruh basins were the main centres of deposition in northern Egypt. These broad depocentres were oriented in NE-SW direction in response to the progressive influence of the Laramide orogeny and the development of the Syrian Arc fold system (Fig. 16). Epeirogenic movements along N E SW trending Kattanyia-Mubarak high, separated Abu Gharadig basin from its eastern extension of Faiyum basin, where each of these basins received up to 1000 m and 880 m of Turonian-Coniacian sediments successively (Awad, 1984; E1 Zarka, 1983). Sedimentation in these depocentres was characterized by the occurrence of repeated cyclic succession of sandstone, carbonaceous shale, and fossiliferous carbonates. Regionally, northeastern Egypt including the Nile Delta region, has received a lesser amount of sediments in comparison to other depocentres. These relatively high palaeotopographic areas were sites of extensive carbonate accumulation including rudistid and coral reefs (Soliman and E1 Badry; fig. 10). The NW-SE trending embayment of the Gulf of Suez basin, on the other hand, was maintained during the Turonian. The Early Turonian open marine section of (Abu Qada Formation) extended along the same trend, but the Late Turonian (Wata Formation) transgressed further south over an irregular topography (E1-Shinnawy and Sultan, 1972). In Sinai the

SEDIMENTARY BASINS OF EGYPT: AN OVERVIEW OF DYNAMIC STRATIGRAPHY

63

MEDITERRANEAN SEA

~L

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,

,

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q

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Fig. 15. Palaeogeographyduring the Cenomaniantransgression (compiledafter Van Houten et al., 1984; Klitzsch and Wycisk, 1987). Turonian sequence consists of shale, marl and sandstone changing into uniform well bedded limestone and dolomite. These deposits represent inner shelf and restricted marine peritidal and sabkha sedimentation in central and northern Sinai. These deposits change into deep marine facies in the offshore area (Jenkins, 1990). The overlaying Coniacian depositional sequence, on the other hand, develop a well defined facies belts because of the ascent of structurally controlled topographic features after the Turonian tectonic events. The Coniacian sequence consists of fluvio-estuarine cross-bedded sandstone interbedded with variegated clays, and marls associated with oyster beds. These change northward into medium to coarse grained glauconitic and bioclastic limestone grading into cross-bedded oolitic shoal complex in central Sinai (Lewy, 1975). Further to the north these facies change into outer shelf, chalky limestone and marls.

In the intracratonic basins of southern Egypt, the Turonian (Taref Formation) is 110 to 200 m thick clastic-dominated sequence. It exhibits an erosion basal contact showing deep scours cutting into the underlying Cenomanian (Klitzsch et al., 1979; Ward and McDonald, 1979; Van Houten et al., 1984; Hendriks, 1986). Above the disconformity, the basal unit of the sequence consists of conglomeratic, feldspathic sandstone with quartz pebbles and intraformational mud clasts and oolitic ironstone (Van Houten et al., 1984). The sequence is followed up by a successive series of fining-up, cross-bedded medium to coarse-grained sandstone that exhibit north and northwest dispersal trends (Fig. 16). The sandstone grades up into argillaceous sandstone exhibiting occasional paleosols or bioturbation. The upper part of the sequence, however, consists of better sorted, finer grained sandstone, rippled siltstone and bioturbated mudstone (Fig. 13B). The textural and mineralogical

64

A.S. EL HAWAT

EROSION

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E

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Uplifted Folds

~

Palaeocurrents, Alluvial Deposits

Fig. 16. Palaeogeographyof Coniacian to Campanian (compiled after Klitzsch and Wycisk, 1987; Van Houten et al., 1984). immaturity of the rocks at the base of the sequence suggest sudden influx of clastic sediments as a result of uplifting of source area (Van Houten et al., 1984). The upwards increase in textural maturity and decrease of grain size suggest deposition of a prograding alluvial plain, delta and tidal flat deposits, which a resulted from a progressively waning clastic influx (Klitzsch et al., 1979; Ward and McDonald, 1979; Van Houten et al., 1984). The Turonian in the intracratonic basins is followed by a hiatus extending from the Coniacian to Early Campanian (Hermina, 1990, fig. 14.3). It is attributed to the uplifting of southern Egypt and development of the Syrian Arc fold system further to the north. Santonian. The Santonian time was the second phase of regressive events after the Turonian within the Late Cretaceous transgression. Unlike the Tur-

onian, tectonics and sedimentation during the Santonian was not associated with major magmatism, however, local volcanic olivine basaltic flows dated as early Campanian (78-84 Ma) were reported from Jabal Uweinat area (Meneisy, 1990). During this regressive phase the sea occupied only the deep basinal areas north of the Western Desert. Deposition in these basinal areas were dominated by fine-grained limestone, chalk and shale under restricted euxinic marine conditions. Because of the effect of the accelerated development of the Syrian Arc fold belt, the N E - S W structures in Bahariya and Siwa oasis became pronounced. Near structural highs the Santonian sequence is missing and the Campanian is often found unconformably overlaying the Cenomanian (Soliman and E1-Badry, 1970). On the Qattarah ridge north of Abu Gharadig basin, in the subsurface, the whole Mesozoic se-

SEDIMENTARY BASINS OF EGYPT: AN OVERVIEW OF DYNAMIC STRATIGRAPHY quence is missing due to erosion and non deposition, and the Cenozoic rocks rest unconformably on the Palaeozoic (Awad, 1984). The Santonian sequence however, is found in nearly all of the northern basins of the Western Desert. At the centre of the Abu Gharadig and Faiyum basins, a thick accumulation of the Santonian deposits rests conformably on the Turonian. These deposits consist of calcareous and carbonaceous pyritic shale that grades upwards and towards the basin margin into chalk and organic carbon-rich limestone. These facies suggest deposition in an outer neritic environment of restricted circulation which gave rise to euxinic conditions. According to Soliman and E1-Badry (1970, fig. 13), the area north of Dakhla and Upper Nile basins was the main depocentre during the Santonian, where siliciclastics were dominant. Meanwhile, carbonates dominated deposition in higher areas of the Nile Delta area and northwestern Egypt.

Campanian-Maastrichtian. During the Early Campanian Egypt was still under the influence of the Santonian regressive phase. Therefore, the Late Cretaceous transgression was not effectively resumed until the Middle Campanian. Egypt was then covered by an extensive shallow sea of partly restricted circulation due to the occurrence of a series of islands to the north. These islands constituted the NNE-SSW trending uplifted areas of the Syrian Arc fold system. Consequently, the prevailing oceanographic conditions led to the development of anoxic geochemical environment in which phosphatic sediments were deposited during Late Campanian. In Dakhla and Upper Nile basins the TuronianConiacian is overlain unconformably by the Middle Campanian (Quseir Formation or Mut Formation). This sequence forms a transgressive, fining-up depositional cycle consisting of varicoloured claystones, siltstones and occasionally conglomeratic very finegrained sandstone. Sandstones and mudstones are laminated, rippled, locally channelled and exhibit trough and herringbone cross-bedded structures. Ichnofossils include Diplocraterion sp. and Thalassinoides. Fossils include silicified wood, plant leafs, fish and other vertebrate remains that include dinosaurs. These facies were deposited in marginal marine, tidal flat and marsh conditions. These are overlain by bioturbated open marine, massive claystone, silty claystone and glauconitic, phosphatic sandstone of Late Campanian-Early Maastrichtian age (Hendriks et al., 1984; Hendriks and Luger, 1987). On the southern margin of the Upper Nile and Dakhla basins the sequence consists of tidal flat deposits that grade into a succession of fluviatile, fluviomarine, deltaic and open marine sandstone and shale facies of the lower part of Kiseiba Formation (Hendriks et al., 1984).

65

The Late Campanian phosphatic sequence (Duwi Formation of Youssef, 1957; Rakhlyat Formation of Hendriks and Luger, 1987) of south Egypt are located in a belt extending from the Red Sea to Dakhla basin (Fig. 17). Phosphatic beds are associated with organic rich shale and claystone that contain variable amounts of dispersed phosphatic pellets, nodules and fish remains. Erosional scours and small channels found in the shales are filled by glauconitic sandstone and phosphate. These shales also contain lenticular, low-angle cross-bedded and rippled phosphate bodies, large sand bars and dunes. On the Red Sea coast, where it was a palaeohigh, phosphate beds are associated with oyster reefal limestone that changes towards the centre of the Upper Nile basin into organic rich shale. It is thought that phosphate nodules and pellets were first originated as a result of precipitation of collophane from interstitial water in mud which was deposited in a quiet water under anoxic conditions. These pellets were later reworked and shaped into sand bodies by bottom currents and during storms, or during period of lowered sea level (Garrison et al., 1979; Glenn, 1979). North of the Western Desert and locally in the southern basins, the Late Campanian-Early Maastrichtian sequence is marked by an unconformity related to tectonic activities. These were associated with the reactivation of the Syrian Arc tectonics, and the updoming of the Red Sea high and the accompanied rejuvenation of the Precambrian fault system. Also, during this time, the connection between the Upper Nile and Dakhla basins was established as a result of tectonic inversion and subsidence of Kharga Uplift and marine transgression (Schandelmeier et al., 1987). These tectonic events were marked by the development of syndepositional mass flow deposits into the Upper Nile basin. Subaqueous gravity inducing slumping, sliding and intraformational conglomerate are commonly associated with laminated gray calcareous silty claystone and thinly bedded clay siltstone showing small-scale fining-up and coarsening-up cycles (Hendriks and Luger, 1987). The Maastrichtian sequence (Dakhla Fm.) is 150 metres thick sequence of shale which may extend in age to the Palaeocene. The Dakhla Formation consists of gray to white marls, laminated claystone with silty intercalations. Shales may contain layers of intraformational conglomerate with phosphatic nodules and vertebrate remains. Oysters and microfauna found in these shales indicate deposition in inner to middle shelf conditions. To the south of Dakhla and the Upper Nile basin, these shales change into deltaic and tidal flat estuarine clastics of Kiseiba Formation (Hendriks et al., 1984, 1987). To the north in the subsurface of the Western Desert, these facies change into chalk and chalky limestone

66

A.S. EL HAWAT

EROSION

[///////;] SHALLOW ~ OPENMARINE

~

PHOSPHATE DEPOSITS

UPLIFTED AREAS AND TEMPORARY EROSION Fig. 17. Palaeogeography of Maastrichtian to Lower Eocene (after Klitzsch and Wycisk, 1987).

with abundant chert bands, and occasional shale beds (Khoman Fm.). The thickness of these chalks may vary from 1644 metres at basin centre to 20-100 metres on structural highs due to syndepositional tectonic movements of fault blocks and active basinal subsidence (Hantar, 1990). These chalk facies extends eastwards into Sinai (Sudr Fm.).

Geological events and sedimentation The Mesozoic evolution and development of sedimentary basins of North Africa and Egypt, following the Hercynian event, was a consequence of the break up of Pangea since the Triassic time. Crustal movements associated with different phases of rifting and opening of the Atlantic ocean and the Tethys, and eustatic sea-level changes have influenced tectonics and sedimentation in the Egyptian basins. The Triassic is represented by a single transgressive depositional cycle restricted in its distribution to northern Egypt due to the residual influence of the Hercynian uplift in southern Egypt. The following

Jurassic sequence is also transgressive in nature as it is clastic-dominated at the base and is carbonate-dominated at the top. It reflects the accelerated rifting and drifting process leading to the opening of the Tethys. In northern Sinai and the Western Desert the Jurassic sequence is arranged in two major depositional cycles exhibiting an increased marine influence through time. Each cycle ensues with deltaic and near shore clastics, and grades up into shale and shallow marine limestones of Middle and Late Jurassic age. Evidence of the Jurassic transgression is not commonly found in the cratonic areas because of the residual effect of the Hercynian uplift (Fig. 8) of central and southern Egypt. Late Jurassic marine clastic sediments, however, were discovered in the subsurface in the Dakhla Basin restricted to Mesaha trough (Wycisk, 1987). The Late Jurassic-Early Cretaceous (150-130 Ma) boundary in Egypt was associated with alkaline volcanic activities that coincide with the initial opening of the Atlantic (Meneisy, 1990). This time was also associated with continued rifting along the

SEDIMENTARY BASINS OF EGYPT: AN OVERVIEW OF DYNAMIC STRATIGRAPHY

67

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Fig. 18. N o r t h - s o u t h s c h e m a t i c c r o s s - s e c t i o n o f J u r a s s i c - C r e t a c e o u s facies in E g y p t s h o w i n g L a t e Jurassic, A p t i a n and C e n o m a n i a n t r a n s g r e s s i o n s i n t e r r u p t i n g " N u b i a n " n o n m a r i n e s a n d s t o n e s e q u e n c e . L a r g e dots, n o n m a r i n e ; small dots, m a r g i n a l m a r i n e ; no pattern, o p e n marine; solid black, horsts. ( M o d i f i e d after Van H o u t e n , 1980.)

continental margin of North Africa and the initial collapse of Sirt rift basin in Libya. The following Cretaceous sequence in North Africa and Egypt forms a single depositional megacycle exhibiting a basal clastic-dominated cycle followed by a carbonate-dominated one at the top. The Early Cretaceous cycle was clastic-dominated because of the overwhelming influence of crustal tectonics, and that of the Late Cretaceous was carbonate-dominated because of the progressive and overpowering eustatic sea-level rise.

Early Cretaceous events: the end of the Nubian problem The Nubian Sandstone is a general but convenient term used to describe characteristic nonmarine to marginal marine sandstone deposits that accumulated on stable platforms and intracratonic basins north of the Afro-Arabian shield regardless of their age and location (Pomeyrol, 1968; Whiteman, 1970; Bonnefous, 1972). In spite of the controversy that was associated with the term since its introduction in 1837, geologists working in the region continued its use, until it was restricted to the Cretaceous. Recently, Klitzsch and others (1979) were able to subdivide the Jurassic-Cretaceous Nubian sequence exposed southwest of Egypt into formations based on the occurrence of transgressive marine interruptions in the sequence (Fig. 13), that are chronologically identical to global eustatic sea level rises of Vail and others (1977). Later, it was possible to correlate Late Jurassic-Early Cretaceous, Aptian and Late Cretaceous transgressions of the Nubian sequence throughout sedimentary basins of North Africa and north of the Arabian Craton (Van Houten,

1980; Van Houten et al., 1984). According to these authors some Nubian regressive sequences may contradict the pattern of global sea level changes, as their deposition was a result of local or regional tectonic uplifts of the craton, and not a results of eustatic sea level falls of global magnitude. Figure 18 illustrates the relationship between regressions and transgressions during the Mesozoic in Egypt. During Early Cretaceous, pre-Aptian time, sedimentary basins of Egypt received up to 1000 metres of clastic sediments from the Nubian craton to the south, the Arabian shield to the east, and from Uweinate-Sirt uplift and Cyrenaica platform to the west (E1-Shazly, 1977; Van Houten, 1980; Klitzsch, 1986). A continental and fluviomarine NeocomianBarremian sequence more than 900 m thick is found in the subsurface unconformably underlain by marine Jurassic in the sedimentary basins of the unstable pericratonic margin of northern Egypt. The intracratonic basins to the south, however, were filled by fluviatile sediments. The following Aptian transgression was a brief, widespread global event that sharply interrupted Early Cretaceous (Neocomian-Albian) regressive phase. It is represented north of latitude 29~ by the carbonate-dominated Alamain Formation, which changes to the south (to latitude 23~ into shallow and marginal marine clastics of Abu Ballas Formation. The upper part of the Early Cretaceous (Albian) is a regressive, fluviatile and deltaic sequence, that is spread out throughout North Africa and Egypt (Van Houten, 1980). Basaltic flows dated at 100 m.y. are found associated with these fluvial deposits (E1-Shazly, 1977). The Late Cretaceous in Egypt was initiated by an abrupt marine transgression that

68 corresponds with the Cenomanian global sea level rise (Vail et al., 1977). Facies and events of the Early Cretaceous sequence correlate well throughout North Africa and north of the Arabian shield (Van Houten, 1980; Van Houten et al., 1984). To the west of Egypt, in the southeast of Sirt basin of Libya, the fluvial-dominated Early Cretaceous Sarir (Faragh, Calanscio) Sandstone sequence is interrupted by clean sandstone units. These sands exhibit sedimentological and mineralogical attributes consistent with marine and marginal marine origin (E1-Hawat and E1 Worfalli, 1990; E1 Hawat, 1992; E1 Hawat et al., 1997). The clean sand facies change laterally to the east into shallow marine, lagoonal and lacustrine variegated shales of Aptian age (Viterbo, 1968). In the northwest-southeast trending troughs forming the deep centre of Sirt basin, on the other hand, the whole of the Early Cretaceous sequence changes into quartzitic marine sandstones rich in nannofossils (Coccolith Fm. of Bonnefous, 1972). Whereas, these major troughs have acted as passageways for the marine advance from the Tethys after the initial rifting and collapse of Sirt-Tibesti uplift, sedimentation in the marginal subbasins was fluvial-dominated during the Neocomian-Barremian and Albian times respectively. During the Aptian eustatic sea-level rise, however, these subbasins were flooded and influenced by marine water leading to the deposition of the middle sandstone and variegated shale units (El Hawat, 1992; E1 Hawat et al., 1997). This transgression was often interrupted by regressive events and siliciclastic influx generated from reactivation of basement blocks of the young rift. In general, depositional facies and events that resulted in the tripartite subdivision of the Early Cretaceous, Nubian sequence in Sirt basin are identical to those of the Dakhla basin. Late Cretaceous events

In contrast to the Early Cretaceous, the Late Cretaceous sequence in Egypt was characterized by an overall transgressive motif. This transgression was initiated by the Cenomanian global eustatic event, which was continued at an accelerated pace to peak during the Maastrichtian. However, this transgression was significantly interrupted by two tectonic events that left their record in the sequence. These are the Turonian and the Santonian tectonic events respectively. Meneisy (1990) noted a general increase of alkaline volcanic activity peaking during the Turonian (90 Ma). This activity was proceeded by the opening of the south Atlantic (Morgan, 1990), and was associated with a change in plate movement between Africa and Eurasia. Indeed, the Turonian time marks the change from extensional-sinistral plate move-

A.S. EL HAWAT ment which led to the opening of the Tethys into to compressional-dextral shear tectonics resulted in the subduction of the Tethys oceanic crust (Smith, 1971). In the intracratonic basins of Egypt this event led to a short epeirogenic tectonic activity on the craton and has led to reactivation of basement source areas, causing clastic influx and creating a short regressive pulse, in otherwise, transgressive conditions. The sequential waning of clastic influx in theses basins suggests resumption of marine transgression during the Coniacian time. The Turonian tectonics also was the cause for the development of the Syrian Arc system in the pericratonic area of northern Egypt and Sinai. This tectonic event is also recognized through out the North African basins. In offshore of Cyrenaica it was marked by a recognizable shallowing event in the deep marine Late Cretaceous sequence (El Hawat and Shelmani, 1993). In Sirt rift basin of central Libya, the event is marked by restriction of marine conditions in the marginal subbasins and deposition of a thick evaporite sequence. Also, the Cenomanian sequence in western Libya is terminated by a gypsiferous marl sequence representing a shallowing event after the initial transgression. Further to the west, in Tunisia and in the offshore area of Tripoli-Gabes basin the Turonian tectonics and associated changes in sea-floor topography have led to the development of anoxic conditions and deposition of deep marine carbonates with high organic carbon content. This anoxic facies are also reported to extend across the Mediterranean to Italy (Bishop, 1988). The second significant tectonic event have taken place during the Santonian. It was associated with the opening of the North Atlantic and represents a continuation to the Turonian event. In the southern basins of Egypt, the intra-Senonian unconformity forms a hiatus extending from the end of the Turonian to early Middle Campanian. However, the Santonian event was relatively weaker than that of the Turonian as no significant magmatism was reported. In north of the Western Desert basins the Santonian event was marked by a major regression, caused reactivation of structural highs of the Syrian Arc fold system and resulted in the deepening of depositional basins. In these basins marine conditions became restricted and fine grain limestone and chalks were then deposited in anoxic organic-matter-rich conditions. North of Cyrenaica this event was associated with shallowing, tectonic inversion of A1 Jabal al Akhdar trough and erosion. In A1 Jabal al Akhdar a major unconformity was developed, it extended westwards into Sirt basin and the Mediterranean offshore (El Hawat and Shelmani, 1993). By the beginning of the Middle Campanian transgression Egypt was covered by a broad, shallow and semi-restricted embayment which was partly barred

SEDIMENTARY BASINS OF EGYPT: AN OVERVIEW OF DYNAMIC STRATIGRAPHY

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@ . ~ CLASTIC WEDGES Fig. 19. Palaeogeography of Early Eocene time, northern Egypt (after Salem, 1976). from the Tethys to the north by Syrian Arc islands. This embayment offered the most favorable anoxic conditions for precipitation of collophane and the accumulation of phosphates in southern Egypt.

TERTIARY Sedimentation during the early Palaeogene continued to be influenced by compressive tectonics and dextral shear movements that dominated the Mesozoic in North Africa. Sedimentary basins of Egypt were subjected to the accelerated influence of these forces which continued to reshape the Syrian Arc fold system and rejuvenate basement fault blocks. The earlier Campanian-Maastrichtian transgression also, continued its inundation of these basins during the Palaeocene, but began to lose pace gradually until the withdrawal of the Tethys from southern Egypt at the end of the Early Eocene (Hendriks et al., 1987; Said, 1962). During this time, sedimentary basins in the unstable pericratonic margin became

narrow, elongate and with irregular bottom topography (Salem, 1976). These basins were shaped as relatively deep depocentres, surrounded by NE-SW trending islands or submarine highs, where thinner Tertiary and thicker Mesozoic sequence are found (Figs. 19 and 20). During the late Palaeogene there was a change in the tectonic style and nature of magmatism in Egypt. The Syrian Arc compressive tectonics that dominated the Late Cretaceous-Early Tertiary was changed into extensional tectonics related to the opening of the Gulf of Suez-Red Sea rift during the Late Tertiary (Morgan, 1990). The associated type of magmatism was also changed from alkaline to tholeiitic basaltic magma types respectively (Meneisy, 1990). This author (Meneisy, 1990) recognized three phases of magmatism and corresponding tectonics in Egypt. These are Late Eocene-Early Oligocene phase (40-t- 10 Ma) related to the Pyrenean orogeny, Late OligoceneEarly Miocene phase (24 4- 2 Ma) related to the initial rifting of the Red Sea and the Early-Middle Miocene phase (20, 18 and 15 Ma).

70

A.S. EL HAWAT 26

28

30

32

34 32

30

0 Sediment Thickness > 1.0 km / / / / ~20km

28

i

i

I

i

Fig. 20. Palaeocene-Eocene main depocentres in northern Egypt (after El Gezeery et al., 1972, 1975).

At the end of the Eocene and during the Oligocene these global tectonic activities were culminated by regression, siliciclastic sedimentation and volcanism throughout North Africa. In fact, the extent of the late Palaeogene and early Neogene regression and siliciclastic influx is not to be under estimated. During the Late Oligocene and Early Miocene time the southern Mediterranean area of Egypt, Libya and Tunisia, as well as, the offshore areas of Tripoli-Gabes and Sirt basins were transformed from carbonate-dominated to siliciclastic-dominated basins as the African shield began to rise (Benomran et al., 1987). This was accompanied by sinking of the Mediterranean and down faulting all along the North African continental margin in response to anticlockwise rotation of Arabia and the opening of the Red Sea and the Gulf of Suez (Sestini, 1984).

Palaeogene Palaeocene The Palaeocene sequence in the pericratonic basins of Egypt rests conformably on the Senonian carbonates but forms an unconformable relationship with the Late Cretaceous on the surrounding structural highs (El Zarka and Radwan, 1986). At the centre of Gindy basin, the Palaeocene sequence reaches a maximum of 283 m in thickness. It consists of light coloured, fine grained lime mudstone with chert nodules alternating with chalk and shale. Further to the south, the Late Cretaceous-Early Tertiary boundary in the intracratonic basins of

Egypt was not accompanied with intense tectonics. The Dakhla Shale Formation found in these basins appears to be lithologically continuous from the Late Cretaceous into the Palaeocene. However, the Cretaceous-Tertiary boundary is marked within the sequence by a thin unit of intraformational conglomerate rich in reworked Cretaceous fauna (Hendriks et al., 1987; Said, 1990). Away from the basin centre the Cretaceous-Tertiary hiatus is followed by glauconitic, conglomeratic material of Middle Danian age (Barthel and Hermann-Degen, 1981). The Palaeocene sequence in the intracratonic basins forms a tripartite subdivision consisting of Danian shale unit (upper part of Dakhla Shale Fm.), a middle limestone unit (Tarawan, Garra Fm. etc.) followed by an upper shale sequence of Landanian age (Esna Shale Fm.). Like the Dakhla shale, the Esna shales consist of varicoloured, marine, often euxinic marl and shale sequence with carbonate interbeds. The middle carbonate unit, on the other hand, consists of chalk, marly limestone. On the basinal margin, south and southeast of the Upper Nile basin these middle and outer shelf facies change into claystone, siltstone and sandstone alternating with coquinoid limestone and overlain by a sequence of massive, nodular fossiliferous limestone, claystone and marl. These facies successively represent inner shelf, lagoonal and middle shelf sedimentation (Hendriks et al., 1984). The boundary of the middle limestone unit with the underlying Dakhla shale forms a significant regional disconformity surface recognized through out

SEDIMENTARY BASINS OF EGYPT: AN OVERVIEW OF DYNAMIC STRATIGRAPHY southern and central Egypt and the Gulf of Suez and the Red Sea area. It corresponds with boundary between the planktonic foraminiferal zones P3 and P5, and is referred to as the "Velascoensis Event" (Strougo, 1986). The limestone above the boundary is extensively bioturbated, develop nodular texture and contains phosphate nodules, vertebrate remains, reworked coral heads and dwarfed fauna (Strougo, 1986; Hermina, 1990). These features may suggest slow rate of sedimentation, early lithification and hardground development. In other places, the basal part of the middle limestone unit consists of peritidal facies rich in sandy marls and vermetid debris (Barthel and Herrmann-Degen, 1981). In the Gulf of Suez-Red Sea area the middle unit consists of olistostromes of boulder size carbonate clasts floating in marly matrix and associated with coarse grained cross-bedded sandstone filled channels. Taken together these features points towards the occurrence of debris flow and mass movements. These were developed in response to syndepositional tectonism associated with faulting and increased of deformation that preceded Red Sea rifting (Strougo, 1986). Eocene

The depositional sequence of the Eocene epoch in Egypt exhibits a general shallowing upward succession of depositional units reflecting continuous and progressive uplift of the African craton as it was responding to the compressive tectonics between Africa and Eurasia. The Eocene megacycle consists of three progressively shallowing-up depositional cycles each of which reflect phases of transgression and sea level falls of global magnitude. In response to tectonics A1 Faiyum basin was subdivided by structural highs into northern and southern depocentres during the Eocene time (El Zarka, 1983). The southern depocentre, Gindy basin, attained a north-south orientation (Fig. 20), that forms a transition between the NE-SW basinal trends of the Western Desert such as Abu Gharadig and the NNW-SSE trend of the Gulf of Suez and the Red Sea high. West of the Western Desert, Siwa and Matruh basins were also trending in NNW-SSE directions in accordance to the structural trend of the Sirt basin of Libya. In the subsurface of the pericratonic basins the Early Eocene-Palaeocene attains a gradational boundary with no clear break in sedimentation. In Gindy basin the Early Eocene sequence constitutes up to 930 metres of alternating lime mudstone, chalk and thin shale and marl interbeds. Locally, the limestone is glauconitic and dolomitic, and the shales are pyritic and calcareous (El Zarka and Radwan, 1986). In the intracratonic basins to the south, the Early Eocene-Palaeocene boundary shows no physical stratigraphic manifestation as the palaeontological

71

boundary is located within Esna Shale Formation. These shales are overlain by the Thebes limestone which offers a recognizable physical stratigraphic Early Eocene boundary in the sequence. In the Thebes Formation in Upper Nile basin forms 300 metres thick shoaling-up cycle of rhythmically bedded lime mudstone and chalk with occasional chert nodules, grading upwards into oyster reefs and alviolinid rich sand shoals west of the basin. To the east, in the proximity of the Red Sea high, the sequence consists of limestone turbidites, intraformational conglomerates and rhythmically bedded cherty limestones, suggesting unstable depositional conditions related to reactivation of the structures along the high. At the basin's centre, the top of the Early Eocene sequence is marked by the development of nodular chalk hardgrounds. In the region of the Red Sea high this horizon is marked by subaerially exposed caliche and micro-karst surfaces (Snavely, 1979). In the Gulf of Suez area, the Early Eocene facies are similar to those found in the Upper Nile basin. However, the top of the sequence consists of reworked phosphatic deposits, which is equivalent to the hard ground and karstified horizon of the Upper Nile basin. The Early Eocene facies were deposited in deeply submerged stable basins throughout Egypt except in the vicinity of the structural highs. The extensive shallowing features at the top of the sequence, on the other hand, are consistent with the global regression which resulted from lowering of sea level at the end of the Early Eocene (Snavely, 1979; Abul Nasr and Thunell, 1987; Strougo et al., 1990). The terminal Ypresian regression have also resulted in the development of a pronounce disconformity surface within a deep marine sequence in AI Jabal al Akhdar (El Hawat, 1985); and it was also associated with restriction of marine water circulation and extensive evaporite precipitation in Sirt basin in Libya. This disconformity extends further to the west in the offshore area of Tripoli-Gabes basin and the onshore of Tunisia. It marks the top of the shoaling-up Ypresian nummulitic sequence and contribute to the development of Metlaoui Group oil reservoir (Bishop, 1988; Bernasconi et al., 1991) The sea regression at the end of Early Eocene from latitude 22 ~ 30'N to latitude 27~ in Upper Egypt, just south of the town of Asyut, led to the emergence of the cratonic areas of Dakhla and Upper Nile basins. During the Middle Eocene, the Lower Nile or Gindy basin was developed as the main depocentre between Asyut and Cairo. In Gindy Basin, the Middle Eocene forms a 717 m thick shoaling-up depositional cycle of bioclastic-algal limestone rich in large forms such as Nummulites gezahensis, alviolina, echinoids and oysters. This facies association was deposited in tidal fiat, bay, back reef shoals and shelf edge environments (Philobbos

72

A.S. EL HAWAT

and Keheila, 1979; Wahab and Khalifa, 1984). At the basin margin in Cairo area, the sequence is 200 m thick and consists of yellowish white, hard, poorly fossiliferous chalky limestone and dolomitic limestone, overlain by white, hard chalky limestone representing deposition in lagoonal, and shallow neritic conditions (Strougo et al., 1992). Further to the west and north of the Western Desert, the Middle Eocene consists of a thin clastic-dominated sequence of sand, silt and shale which was derived from the erosion of the Cretaceous structural highs (El Zarka and Radwan, 1986). Similar facies changes are found in the Gulf of Suez basin, where the Middle Eocene also, constitutes a shoaling-up nummulitic carbonate sequence. The sequence changes north of the basin into an increasingly clastic-dominated succession of shales, silts and interbedded limestones. The introduction of these clastics were also, attributed to the occurrence of Late Cretaceous structural highs source north of Sinai. The Middle Eocene sequence in the area is capped by a second reworked phosphatic 28 ~ I

bed (Abul Nasr and Thunell, 1987). The shallowing event of the Middle Eocene sequence which is recognized in different areas of Egypt is consistent with the global eustatic lowering of the sea level which marks the end of the Middle Eocene (Abul Nasr and Thunell, 1987; Strougo et al., 1992). A progradational, shoaling-up nummulitic Lutetian sequence overlain by peritidal facies and followed by a disconformity is exposed in Cyrenaica (El Hawat, 1986; E1 Hawat and Shelmani, 1993). The Lutetian sequence in the subsurface in Sirt basin consists of large-scale, progradational clinoform structures recognized in seismic lines. Continued lowering of the sea level during the Late Eocene resulted in the progressive emergence and erosion of the structural highs that provided clastic sediments to the basins of northern Egypt (Fig. 21). The Late Eocene in the Western Desert consists of a 220 m thick sequence of sandy limestone sandstone and shales associated locally with oyster banks. These represent deposition in a shallow, neritic, lagoonal and deltaic conditions. In the

30 o

I

I

I

32 ~ 32 o

Alexondrio /

7/I

/~//

,,z\

30 ~ -

/

'/// .30 ~

:r

.~.

iJj

:>8~ -

.28 o

I

!

3~o

28 ~ 0 L

50 I

I00 I

150 I

-

!

2 0 0 KM. I

LAND

~ I ~t

STRUCTuRAL HIGHS, ISLAND

DELTAIC SEDIMENTS

////! ///I

BASINAL

MUD

BASINAL

MARL

Fig. 21. Palaeogeography of Late Eocene time in northern Egypt (after Salem, 1976).

SEDIMENTARY BASINS OF EGYPT: AN OVERVIEW OF DYNAMIC STRATIGRAPHY subsurface, the Late Eocene facies are often undistinguished from the overlying Oligocene deposits. The Late Eocene siliciclastic deposits generally thicken northwards, but may also, thicken locally toward basin centres as they prograde away from the surrounding structural highs (El Shazly, 1977; Salem, 1976).

Oligocene The Tertiary marine regression and tectonic upheaval continued in Egypt during the Oligocene as the Arabian plate started to move in an anti-clockwise direction around a pivotal point in the Jordan (Fig. 22). This movement was initiated in response to resistance to differential drifting of the northwestern foreland of Africa in comparison to the northeast. Causing the movement of Sinai to the southwest, and the subsequent opening of the Gulf of Suez (Klitzsch 1986; Schandelmeier et al., 1987). These movements also led to the development of the NW-SE, ENEWSW fault systems which were associated with extensive volcanicity throughout Egypt. These were combined to contribute to the development and subsidence of the Gulf of Suez and the Nile Delta basins (El Shazly, 1977; Rizzini et al., 1978; Said, 1981). The Oligocene in Egypt was a time of uplift, regression, volcanicity and continental sedimentation

73

(Fig. 24). Outcrops southwest of Cairo consist of 250 m of sandstone and gravel with local limestone and shale interbeds of fluviomarine origin (Said, 1962). These facies change westwards into fluvial sands and gravels, and grade northwards into deltaic siltstone, shale and glauconitic sandstone. Locally, continental Oligocene facies are overlain by basaltic lava flows up to 250 m thick (El Zarka and Radwan, 1986). North of the Gulf of Suez basin, the Oligocene consists of fossiliferous foraminiferal marls. These change southwards into red beds, consisting of reddish, clayey and pebbly calcareous sandstones, associated locally with post rifting basaltic flows (Salem, 1976; Chowdhary and Taha, 1987).

Neogene Miocene The Early Miocene transgression was associated with deposition of siliciclastics (Moghra Formation) in the Western Desert (Said, 1962). In the subsurface the Miocene consists of 615 m thick coarsening-up deltaic sequence of sandstone, shale and limestone intercalations. This sequence pinches out eastwards towards the Nile Delta in the region of Tiba basin, and changes northwards into pro-delta silts and marine shale (Fig. 25). Following the initial Miocene

/ t

//

ARABIAN ) PLATE

,,,;/ T-EARLY

TERTIARY

I) Arabian Plate moves faster than Nubian Plate. 2) Development of NW-SE Riedel shear. 3) Development of NE-SW Syrian Arc folds.

"n"- OLIGOCENE- MIOCENE

"rrr- LATE

4) Anti- clockwise movement of Arabian Plate. 5) Opening of Gulf of Suez.

6) N.

MIOCENE- PLIOCENE 8= QUATERNARY

movement of Arabian Plate along Aqaba fault. 7) Opening of Red Sea graben.

Fig. 22. Sketch diagram of the structural developmentof Gulf of Suez-Red Sea-Gulf of Aqaba graben system (after Klitzsch, 1986).

.,,j

EOCENE AFTER FAULTING

Q L. EOCENE

~...~...........~. LATE CRETA. . . . ~ ... --_...--_--~! I I ~ valeocene . . . . I~ I I ~

~

I

/

Reworked lower Eocene breccia / a n d Conglomerate ln Igoonol deposits

f

~

Q .

(/., \ r~~~

../

END OF OLIGOCENE

Pre Miocene

j

ormlty

Fan Conglomerate ~ Lower- Middle Mlocene

/

Q

I

MIDDLE MIOCENE LOWER MIOCENE REER

/ / I

w/

/

/ ~ J

d

~

.

.

.

.

m

J

> f./3

Fig. 23. Rift margin block faulting development on the western side of Gulf of Suez (after Klitzsch, 1986).

t-n t--' :i: >

SEDIMENTARY BASINS OF EGYPT: AN OVERVIEW OF DYNAMIC STRATIGRAPHY 26 ~

?_8 =

I

I

,

,I

5 0o I

I

75

:52 ~ I

:54 ~

I

:52 ~

52 ~ M e d i t e r r a n e a n Sea

\

\

Shelf \

Carbonates ?

\ \ 30 ~

N

30 ~

28 ~

:)8~ 0

~

SLOPEFAN

I00

150

2C

Red

T 26 ~

50

28 ~

:50 ~

I!l /I II I]

PRODELTA

32 ~

[~'~'i'ii]

Sea

5 4~

sANos

LAND

DELTA

Fig. 24. Palaeogeography of the Oligocene, northern Egypt (after Salem, 1976).

transgression in the Delta basin, the Aquitanian was a time of low sea level and uplift throughout the delta as indicated by the low sedimentation rate of 60 to 70 m/Ma. East of the Delta a higher sedimentation rate was estimated (400-500 m/Ma) during the Burdigalian (Wray, 1985). In the Gulf of Suez basin the Oligocene rifting was followed by Early Miocene marine transgression over an irregular basin floor. The northern part of the basin centre was filled by 2200 m thick Globigerina marls that thins southwards to 200 m at the Gulf's entrance (Salem, 1976). These marls change laterally and upwards towards the rift's margin into calcareous sandstone, limestone and interbeds of marls and gypsiferous shales. The outer margin of the rift consists of crystalline basement half horsts and graben structures forming a series of platforms and basins. These are covered by Early Miocene sediment apron of mixed alluvial gravels conglomerate and sabkha evaporites that grade basinwards into platform carbonates and basin rim reefs (El Haddad et al., 1984; Coniglio et al., 1988; James et al., 1988). It was noted that the Early Miocene sequence was interrupted by tectonic movements, causing reactivation of faults and tilting of fault blocks within the basin (Fig. 23). These tectonics were associated with the development of an unconformity in the basinal sedimentary sequence, and the development of karstification, subaerial diagenesis of platform carbonates and synsedimentary collapse of basin rim reefs (James et al., 1988).

Salem (1976) suggested that the Middle Miocene transgression was associated with the eastwards shift of the Palaeonile from the Western Desert to its present day position (Figs. 25 and 26). The resulting cessation of siliciclastic influx into the Western Desert gave way to an increase in carbonate productivity and the deposition of Marmarica Formation that extended into Cyrenaica. During the Middle Miocene (Langhian), the Nile Delta was affected by tectonic instability and gravity faulting. This was followed during late Middle Miocene (Mid and Late Serravalian) by a pronounced sea level fall causing a depositional hiatus that ranges from 6 to 2 Ma in duration, and may extend to Early Tortonian (Wray, 1985; Harms and Wray, 1990). The Middle-Late Miocene delta cycle is a coarsening-up depositional sequence that thins gradually southwards in the direction of Cairo (Fig. 28). The early part of the cycle consists of more than 700 metres of green, gray claystone interbedded with fossiliferous marls and rare quartzose sandstone interbeds. Following an unconformity, the second part is of Late Tortonian-Early Messinian age. It consists of 1-3 km thick, northwest prograding clinoform sequence the deposition of which was taking place in association with rapid subsidence of the eastern part of the delta, which was influenced by an easterly longshore drift. During this phase the delta maintained a high sedimentation rate that reached 680 m/1 Ma (Wray, 1985). Sediments constituting this depositional phase consist of poorly

76

A.S. EL HAWAT 28 ~

26 ~ !

32 ~

30 ~

!

32 ~ i

3 4~ !

i

32 ~

\ \ \ \ \ 30 ~

\ 32 ~

\ \ \

e@

9

28 ~ 28 ~

I

I

26 ~

I

I

28 ~

I

30 ~ T

50

I00

I

I

150

I

32 ~

i% m ~

y&ll

V

34 ~

200KM.

I

LEGEND LAND

CARBONATES [VvV vV ~ E V A P O R I T E S

DELTA

ALLUVIAL

FAN

Fig. 25. Palaeogeography of Early Miocene time, northern Egypt (after Salem, 1976).

sorted conglomeratic sands, clays and calcareous sandstone representing a prograding fluvio-deltaic and coastal deltaic sedimentation and associated swamps and lagoons. Wray (1985) and Harms and Wray (1990) pointed out that during this phase of the delta history, a major integration of the Nile drainage system has taken place for the first time. North of the delta region, the Late Miocene cycle is capped by 40 m thick unit of Late Messinian anhydrite interbedded with thin clay layers (Fig. 27). In places, the top part of the delta sequence exhibits deeply incised channels and slump structures that produce an angular unconformable relationship with the overlaying Pliocene delta cycle (Rizzini et al., 1978). In the Western Desert, the Tortonian-Messinian sequence consists of 30-40 metres thick cross-bedded, well sorted, medium grained sands, clays and limestone, that grade up into gypsiferous clays and coarsely crystalline selenite (Omara and Sanad, 1975). The sequence exhibits a disconformable relationship with the Early and Middle Miocene below, and the overlaying Pliocene carbonates. In the Gulf of Suez basin the Middle Miocene consists of calcareous shale and marl interbedded with anhydrite, and grading upwards into massive anhydrite and rock salt and associated with gray shale interbeds. The sequence is overlain by Late

Miocene anhydrite with minor interbeds of sandstone and shale (Chowdhary and Taha, 1987). These evaporites attain a maximum thickness of 3540 metres at the southern end of the basin and decrease northwards to zero (Salem, 1976). Taken together with thickness variation of the Early Miocene open marine Globigerina marls, it was concluded that these evaporites were deposited in a barred basin which was opened to the Tethys from the north (Said, 1962; Salem, 1976, fig. 16). On the basinal margin, the Middle-Late Miocene sequence is similar to that of the Early Miocene. It consists of peritidal carbonates associated with basin-rim reefs and subtidal stromatolites on the gulf side. Away from the basin centre these sediments change into alluvial clastics near basement highs. This depositional suite of carbonates and clastics were covered by Late Miocene evaporites and was followed by evaporite solution collapse breccia that resulted from subaerial exposures and diagenesis (El Haddad et al., 1984; Coniglio et al., 1988; James et al., 1988). Pliocene Following the Messinian lowering of sea level during the Mediterranean salinity crisis (Hsti et al., 1973), and due to lowering of the base level of erosion, the River Nile drainage became deeply

SEDIMENTARY BASINS OF EGYPT: AN OVERVIEW OF DYNAMIC STRATIGRAPHY 26

28

v

v

30

!

!

MEDITERRANEAN

!

32 !

A

77 54 !

i

:52

SEA

Ale

\

\ \ 30

>, gg

SINAI

i i i

i

0 I

50 I

____J~_

I00 150 200 Km I

I

EASTERN DESERT

~(~./,~O,f

28

I

__._m,..--._...L

Delta fan

V~.~_~I

Shale

Alluvial fan

~

Calcareouss.s.

Slope fan

~

Carbonates

Land

~

Evaporites.

Fig. 26. Palaeogeography of Middle Miocene time, northern Egypt (after Salem, 1976).

entrenched as far south as the city of Aswan 1200 km inland (Ryan, 1978). The occurrence of the Pliocene (Plaisancian) marine and estuarine clays and sands in the vicinity of Aswan (Chumakov, 1968) was attributed to the opening of the Atlantic floodgate and refilling of the Mediterranean basin during the Early Pliocene transgression (Hsti et al., 1973). The Early Pliocene estuarine deposits were followed by Late Plio-Pleistocene fluviatile sequence. In the Nile Delta basin, an Early Pliocene transgressive sand unit (Abu Madi Formation) forms the base of the Plio-Pleistocene Nile Delta cycle (Fig. 27). It consists of thick- bedded, rippled, crossand plane-bedded, bioturbated sandstone, interbedded with clay and occasional conglomerate at the base. The following delta cycle is 1500 m thick coarsening-up sequence that consists of deep shelf and slope clays with few quartzose sand interbeds. These sand beds increase in thickness and frequency upwards, as they grade into 300 m thick, large scale prograding foresets of coarse- to medium-grained deltaic sand units. The following 700 m of the sequence is thick-bedded, conglomeratic, coarse- to medium-grained quartzose sandstone with reworked chert, quartzite and dolomite pebbles. The top of the sequence also contains coquina and peat deposits representing coastal and lagoonal sedimentation that

extended up to the Quaternary (Rizzini et al., 1978). West of the delta in the Faiyum area, the Pliocene consists of limestone, sandstone and bioclastic shoreline facies. These may change laterally in Wadi Natrun into an association of clay, limestone and sandstone with vertebrate remains indicating fluvio-estuarine and marginal marine conditions. Further to the west in the Western Desert and away from the influence of the Nile, the Pliocene sequence consists of 60 metres of pink oolitic limestone which is unconformably underlain by the Miocene (El Shazly, 1977). In the Gulf of Suez basin, the Late Miocene regression and erosion was followed by subsidence associated with the opening of the Red Sea to the Indian Ocean. The post-Miocene sequence in the subsurface is up to 950 m thick, coarse to mediumgrained sandstone interbedded with red brown claystone and siltstone. These contain minor intercalations of limestone and anhydrite, associated with fauna of Indo-Pacific affiliation (Chowdhary and Taha, 1987). On the basin margin, active extensional tectonics and the associated depositional suites that began in Early Miocene time continued until the present (Purser et al., 1987). It exhibits rapid lateral variation of carbonates (reefs), evaporites and siliciclastic sediments. This is a typical facies association

78

A.S. EL HAWAT

Rock Units AGE

Environments u

AVG.

Formation

Lithology

Thickness Metars

Holocene

Bilqas

50

a _u ~ a'6o .~._

"6 ~

E.~

to.

Neritic

~. J.

"-- 9149 o

o

o

o

o

o

o

Pleistocene

Mit Ghamr

700

9 ~

.

9

~176

.

.

.

~

~

~

.

Upper El Wastani

300

....

9

,,

9

,,

~

~

,

9

Middle Kafr al Sheik

U.I Z uJ c.) 0 _J 13.

1500

Lower Abu Madi

300 "

"'"'~

~

Rosetta

50

~

~

v Vv Vv

'OoOoO ~

"~

9 ~176

r ~

Messinian Qawasim

UJ Z UJ

9

9

~

....

0

Tortonian

Serravalian

Sidi _ .. Salim

>700

~

!

Langhian Fig. 27. Stratigraphic column of the Nile Delta (after Rizzini et al., 1978).

related to rift margin depositional climatic conditions (Fig. 29). Under sporadic wet periods causing flash coarse clastics from the hinterland.

set up in arid these conditions floods provided These were fol-

lowed by a rapid return to a siliciclastic-free, clear marine water leading to carbonate sedimentation and evaporite mineral precipitation in areas of restricted circulation (Purser et al., 1987).

SEDIMENTARY BASINS OF EGYPT: AN OVERVIEW OF DYNAMIC STRATIGRAPHY

79

N

S It

Pleistocene

~

~

~

t-z-o~

o".

o

.~- .-'~-, -"-~." ". ".'.'-'.

Holocene

~=

9

~

"'"

"-~----- :-~

."5"- . .

.

~

_ L ~ - - ~

.J'-~." ! . .---C<'.. " ~ . " ~ .

"- "

- . .:--~-~~'.

,-;~,~~

"~-<

9 ".

..-;-z.-_-f . : . . _ . .~...--"f ~," . ~ . . . L - . - . . - . . . - . - ~ . . . _ = _ " -- ; - .

I.

Mid-Up

---

"

-.

P,ioc n . ......

l-L..---:_.

i,.UW~I

-/ ./.~. ,,,I. I.-..../.. - . . . . . . . . .

9

n"

Pl,oeene

I/

/

.

.

7

_..--~

.

;/,~.~^

.

.

,-

.

. ~

.

.

.

.

.

.

. ...L ~..~....-

.~.~;,

Messln,an Messinian

.

.

.

.

Langhian Serravalian

_ / -

~

- ~

9

-~-

9

.~'....,;7...

-

9

9

;/..,/

: 1 ~ . ~ ~ j

.

.. . . . . .

9~

." - .:--"..~ ~ 9 ~--~_ ~_~

"

- - ~ 9- ' ~. q " : w "

~':.~

~.

Fig. 28. North-south sketch profile of the Neogene-Quaternary of the Nile Delta (after Rizzini et al., 1978).

Geological events and sedimentation N e o g e n e f a c i e s a n d e v e n t s in N o r t h A f r i c a

Facies changes are generally attributed to local variation of palaeogeographic settings, relative sea level change and/or to variability of clastic supply due to the prevailing local or regional tectonism. The influence of global geologic events such as eustasy, on the other hand, often extend to form a unifying common factor across these facies areas. Recognition of these events in the stratigraphic record becomes an essential element in establishing an actualistic stratigraphic subdivision of unfossiliferous depositional sequences, or those that contain no time diagnostic fossils. We know from the results of the Mesozoic study in Egypt how this technique was used with great effect in the subdivision of the Nubian Sandstone sequence. It has, also, led to the recognition of the Late Miocene, previously included in the Middle Miocene sequence in Libya and the Western Desert of Egypt (E1-Hawat and Salem, 1985, 1987; Youssef, 1988). The same principle is being applied here to correlate the Miocene throughout North Africa.

The Palaeonile Delta area was the main focus of siliciclastic sedimentation in northeast Africa during the Middle Miocene time. Away from this area, the influence of carbonate sedimentation increased gradually towards the east, in northern Sinai, and westwards into the Western Desert and northern Libya. In the latter areas, shallow marine Middle Miocene carbonates grade southwards into fluviatile and fluviomarine siliciclastics as in Sirt basin (Selley, 1968; EI-Hawat, 1980a), southeastern Cyrenaica (Bellini, 1969), and the Western Desert (Salem, 1976). High resolution biostratigraphic investigation of the Nile Delta indicated that a major sea level lowering has taken place in the Mediterranean, during the latter part of the Middle Miocene (Wray, 1985; Harms and Wray, 1990). It has resulted in the development of an unconformity in the delta sequence. It represents a hiatus of a minimum 2 Ma duration, that may extend locally to early Late Miocene, into the Tortonian. The influence of this eustatic event is recognized throughout the Mediterranean, as it was preceded by the separation of the Tethys from the Para-Tethys some 14 to 15 million years ago (Hsti et al., 1977). The Middle-Late Miocene unconformity extends

80

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from Sinai and the delta, to the Western Desert of Egypt. In Libya, it resulted in the development of the Middle-Late Miocene off-lap relationship reported from northern Cyrenaica and Sirt basin (E1-Hawat and Salem, 1985, 1987; EI-Hawat et al., 1985). The same unconformity was recognized further south in Sirt basin, and was used as a correlation datum for the study of the Middle Miocene (Selley, 1968; EI-Hawat, 1980a). Although the earlier part of the Late Miocene was a time of shoaling of the Betic Strait that connected the Mediterranean to the Atlantic Ocean. Open marine facies has prevailed in the Mediterranean during this time (Hsfi et al., 1977). Stratigraphic record in northeast Africa suggests that the Late TortonianEarly Messinian was a time of a limited transgression, which did not alter the net off-lap relationship with the Middle Miocene sequence. However, evidence of this transgressive event is demonstrated by the occurrence of local unconformable on-lap relations found between the Late Miocene and older formations on the marginal highs of western and eastern Libya (E1-Hawat et al., 1985) and the Western Desert of Egypt (Omara and Sanad, 1975). In Libya, depositional facies were dominated by shallow marine platform carbonates and reefs developed over topographically high areas. Topographic lower areas of Sirt basin, on the other hand, were filled by siliciclastics due to the influence of rivers. Eastwards in the Western Desert, the sequence consists of mixed siliciclastics and carbonates that grades further to the east into the clastic facies of the delta. In the latter, rapid subsidence rates and integration of the Palaeonile drainage system has contributed to

accumulation of up to 3 km thick delta clinoform sequence (Wray, 1985; Harms and Wray, 1990). The sudden onset of the Mediterranean salinity crisis during the Late Messinian was a result of the closure of its connection with the Atlantic Ocean (Hsfi et al., 1973, 1977). As a result, two types of facies associations characterize the Late Messinian sedimentation in North Africa. Their distribution was controlled by structures and geography. These are either evaporite-carbonate or evaporitesiliciclastic associations. The former is found on platforms and structurally high areas away from siliciclastic sources; and the latter is found in the structurally low areas which were influenced by river sources. In structurally high areas in Jabal Naffusa of western Libya, Sirt basin, northern Cyrenaica and the Western Desert of Egypt the Late Messinian consists of lenticular bodies of coarsely crystalline gypsum and gypsarenite, associated with limestone consisting of pelletal mudstone, algal stromatolites and oncolites (E1-Hawat, 1980b; Youssef, 1988). In the structurally low areas of Sirt basin, and off the Nile Delta, evaporites are associated with siliciclastics exhibiting evidence of subaerial exposure (De Heinzelin and E1-Arnauti, 1983; Rizzini et al., 1978). Due to continuous evaporation and consequent fall of the Mediterranean sea level, during the Late Messinian the resulting lowering of the base level of erosion caused deepening and lengthening of channels located in the low areas of Sirt basin and along the Palaeonile. In Sirt basin, the entire Miocene sequence was locally eroded by the post-Messinian Sahabi Channel (Barr and Walker, 1972), which also extended northwards in the Mediterranean offshore.

SEDIMENTARY BASINS OF EGYPT: AN OVERVIEW OF DYNAMIC STRATIGRAPHY In the Nile Delta this event is marked by an unconformity at the Miocene-Pliocene boundary (Harms and Wray, 1990). Erosion in the River Nile produced a deeply incised canyon that reached 2,500 m below sea level north of Cairo, and - 2 0 0 m in the area north of Aswan (Said, 1981). With the opening of the Atlantic floodgate at Gibraltar (Hsu et al., 1977), the Early Pliocene transgression transformed the Sahabi Channel and the Nile into impressive inland estuaries. Early Pliocene (Palisancian)estuarine fauna was discovered in an incised channel some 1,250 km from the present coast (Chumakov, 1972). The Sahabi Channel, set between the sea level and - 4 0 0 m; was filled by calcareous sandstone, sands and shales, and contains mixed marine and continental fauna of reworked Miocene and Pliocene age (De Heinzelin and E1-Arnauti, 1983, 1987). On the higher ground, the Early Pliocene consists of oolitic and bioclastic limestone; that rests unconformably underlain by Late Messinian evaporites.

SUMMARY AND COMMON THEMES

The Afro-Nubian craton consists of two parts separated by the River Nile. The western and oldest part of the craton is Late Archean nucleus surrounded by an Early to Middle Proterozoic metamorphic rocks. The former constituted granulite rocks of an early protocrust, and the latter consists of granitoid and polymetamorphic rock extending throughout northeast Africa. East of the Nile, the Nubian shield was developed as a result of crustal accretion during Late Proterozoic, Pan-African thermo-tectonic events. Evolution of the Nubian shield involved an early flysch sedimentation and crustal subduction associated with island arc volcanics and ophiolites. A phase of orogenesis and plutonic granitoid intrusions was followed by erosion and development of foreland basins. This in turn was followed by a final phase of late Orogenic event and post-Precambrian peneplanation. Basin development during the Phanerozoic was influenced by the Proterozoic crustal evolution and by the location of the North African plate between the relatively more rigid plates of northwest Africa and Arabia. Post-Pan-African crustal cooling, differential movements along old structural lineaments, drifting and collision with the northern continents, and development of hot spot anomalies beneath the craton were the main causes of tectonic movement, basin evolution and development in Egypt. Basin filling was often initiated by thermal crustal up-arching, fracturing and subsidence followed by marine transgression. These evolutionary stages were recorded in symmetrical or asymmetrical depositional cycles consisting of basal continental clastics grading into marginal marine sediments and open marine carbon-

81

ates and/or evaporites. Depending on their location in the craton and their tectonic history, the sedimentary depocentres of Egypt are either intracratonic, pericratonic, or rift basins. Sedimentation and basin filling during Early Palaeozoic was influenced by the regional northward tilt of the post-Pan-African plate, and the occurrence of the N N W - S S E trending horst-graben basement structures and mountains, continental glaciation and eustasy. These factors were combined to produce siliciclastic-dominated depositional regimes. Whereas continental and fluvial sedimentation were closely associated with epeirogenic tectonic movements and erosion of basement blocks. The Palaeozoic marine transgression during the Early Cambrian, Early Ordovician, Early Silurian and Early Carboniferous led to the reworking and widespread distribution of siliciclastic sediments over the craton. Dakhla, Siwa and Abu Gharadig basins were the main Palaeozoic depocentres in Egypt. Other depocentres have received a thinner sequence of Palaeozoic deposits. At the end of the Visean, Central and Upper Egypt was uplifted and erosion of the Early Palaeozoic took place. These tectonic movements have resulted from the collision of Gondwana with the northern continents on one hand, and because of the development of thermal anomalies beneath the plate, along Jabal Uweinat-Aswan trend on the other. These tectonic and palaeogeographic conditions have reversed the drainage flow direction of rivers southwards and restricted the Permo-Triassic sedimentation to northern Egypt. The inherited Permo-Triassic structural elements were terminated during the Mesozoic by rejuvenation of old Palaeozoic N N W - S S E structural trends and restoration of the northern tilt of the African plate. These were associated with the opening of the central Atlantic, extensional tectonism and sinistral strike-slip movement between Africa and Eurasia and the opening of the Tethys. Like Early Palaeozoic, the Mesozoic sedimentary sequence consists of a succession of depositional cycles reflecting tectonics and eustasy. Because of the above indicated constraints the Triassic and the Jurassic sequences were restricted in aerial distribution to northern Egypt. Sedimentation during Early Cretaceous in Egypt was essentially regressive in character due to the overwhelming influence of basement tectonism. The NeocomianBarremian and Albian sequences consist of fluvialand continental-dominated deposits, separated by the Aptian transgressive phase that covered most of southern Egypt. By contrast, the Late Cretaceous was a time of overwhelming transgression, which was associated with increased carbonate productivity and sedimentation. It started with the Cenomanian global sea level rise which reached its climax at the

82 end of the Cretaceous and Early Tertiary. However, the Late Cretaceous sea level rise was interrupted by two regressive phases associated with compressive epeirogenic and shear tectonic movements during the Turonian and Santonian times. These tectonic events were contemporaneous with the opening of the north Atlantic and the switch in the direction of plate movements along the Tethyan realm. The Tethyan tectonics were transformed into compressive and dextral shear movements leading to subduction of the Mediterranean oceanic crust and eventual uplift of the African craton. These events were associated with volcanic activities, development of unconformities, siliciclastic influx and restriction of marine circulation in sedimentary basins leading to evaporites deposition or anoxic conditions associated with precipitation of collophane. The most important tectonic feature in the pericratonic area, however, was the inversion of subsiding basins and the development of the N E - S W to E N E W S W trending Syrian Arc fold system. Transgression and the northwestern drift of Africa continued to influence sedimentation and tectonics during the Early Tertiary. The Early, Middle and Late Eocene sequences in Egypt are shoaling-up, transgressive-regressive cycles reflecting the global eustatic events. By the Late Eocene, the North African craton began to rise due to the accelerated compressive forces between Africa and Eurasia and the consumption of the Tethyan oceanic plate. At this time, differential movement between the Nubian and Arabian plates initiated N W - S E shearing along the Gulf of S u e z - R e d Sea high, which was culminated by extensional tectonics and the opening of the Gulf of Suez during the Oligocene. These tectonic events were the main cause for the retransformation of sedimentation throughout North Africa from carbonateinto siliciclastic-dominated depositional regimes. The Miocene started with marine transgression affecting the northern Pericratonic area and Gulf of Suez basin. This was followed at end of the Middle Miocene by a major eustatic fall in sea-level. Also, during the Miocene the River Nile and the delta assumed their present day position. The closure of the Mediterranean and the Messinian sea-level fall was marked by River Nile erosion and evaporite sedimentation north of the Delta and the Western Desert. Finally, the opening of the Red Sea to the Indian Ocean coincided with the flooding of the Mediterranean and drowning of the River Nile by marine water during the Pliocene.

ACKNOWLEDGEMENTS

This paper is dedicated to Prof. Soliman M. Soliman and Prof. Mourad I. Youssef of Ain Shams

A.S. EL HAWAT University, Cairo; Prof. William R. Dickinson formerly of Stanford University, California; Dr. Graham Evans, Prof. Richard C. Selley and Prof. Douglas J. Shearman of Imperial College, London, whose teaching and work was an inspiration to the author. I am grateful to Dr. R. A1-Khazmi, of Garyounis University, Libya, for reading drafts of this paper and for suggesting improvements to the text. Thanks are due to Ms. R. Banger of the University of Cambridge for her help during library research. The support provided by the Arabian Gulf Oil Co. (AGOCO) and GEOLIBYA, Benghazi, is greatly appreciated. This paper is a contribution of the IGCP Project No. 369 on the Comparative Evolution of Peri Tethyan Rift Basins.

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Paleontology and Geology of Sahabi. Alan R. Liss Inc., New York, N.Y., pp. 1-22. Druckman, Y., 1974. The stratigraphy of the Triassic sequence in southern Israel. Bull. Geol. Surv. Isr., 64: 1-92. E1 Baz, E, 1984. The Geology of Egypt: An Annotated Bibliography. E.J. Brill, Leiden, 778 pp. El Gaby, S., List, EK. and Tehrani, R., 1990. The basement complex of the Eastern Desert and Sinai. In: R. Said (Editor), The Geology of Egypt. A.A.Balkema, Rotterdam, pp. 175-200. E1 Gezeery, M.N., Farid, M. and Taher, M., 1975. Subsurface geological maps of northern Egypt, unpublished maps. General Petroleum Co., Cairo. El Gezeery, M.N., Mohsen, S.M. and Farid, M., 1972. Sedimentary basins of Egypt and their petroleum prospects. 8th Arab Petroleum Cong., Algiers, VIII, Paper 83 (B-3). El Gezeery, N. and Taha, I., 1971. Contribution to the stratigraphy, tectonics and oil shows in Abu Charadig Basin, Western Desert. Abstract, 9th Ann. Meet. Geol. Soc. Egypt, pp. 6-8. E1 Haddad, A., Aissaoui, D.M. and Soliman, M.A., 1984. Mixed carbonate-siliciclastic sedimentation on a Miocene fault block, Gulf of Suez, Egypt. Sediment. Geol., 37:185-202. E1 Hawat, A.S., 1980a. Carbonate-terrigenous cyclic sedimentation and palaeogeography of the Marada Formation (Middle Miocene), Sirte Basin. In: M.J. Salem and M.T. Busrewil (Editors), Geology of Libya, Vol. II. Academic Press, London, pp. 427--448. E1 Hawat, A.S., 1980b. Intertidal and storm sedimentation from Wadi al Qattarah Member, Ar-Rajmah Formation, Miocene, A1 Jabal al Akhdar. In: M.J. Salem and M.T. Busrewil (Editors), Geology of Libya. Academic Press, London, pp. 449-462. E1 Hawat, A.S., 1985. Submarine slope carbonate mass-movements in response to global lowering of sea-level: Apollonia Formation, Lower-Middle Eocene, AI Jabal al akhdar, NE Libya. Abstract, 6th. European I.A.S. Meet., Lleida, Spain, pp. 152-155. El Hawat, A.S., 1986. Fine-grained current drift carbonates and associated facies in a slope to shelf shoaling-up sequence: The Eocene, NE Libya. Abstract, European I.A.S. Meet., Krakow, Poland, pp. 208-210. EI-Hawat, A.S., 1992. The Nubian Sandstone sequence in Sirte basin, Libya: Sedimentary facies and events. In: A. Sadek (Editor), Geology of the Arab World, Vol 1, Cairo Univ., pp. 317-327. E1-Hawat, A.S. and El-Worfalli, H.O., 1990, The occurrence of marginal marine facies in the Sarir (Nubian) sandstone sequence and its regional stratigraphic significance, Sirte basin, Libya. Abstract, 15th Colloquium of African Geology, CIFEG, Univ. Nancy, p.84. El Hawat, A.S. and Salem, M.J., 1985. Stratigraphic reappraisal of Ar-Rajmah Formation, Miocene, AI Jabal al Akhdar, N.E. Libya: a case of field sedimentological approach. Abstract, 8th. Congr. Reg. Comm. Meditereranean Neogene Stratigraphy, pp. 206-208. El Hawat, A.S. and Salem, M.J., 1987. A case study of the stratigraphic subdivision of Ar-Rajmah Formation and its implication on the Miocene of Northern Libya. Ann. Inst. Geol. Publ. Hung., 70: 173-183. E1 Hawat, A.S., Salem, M.J. and Megerisi, M., 1985. Depositional sequences and history of Gulf of Sirte palaeoshoreline: Middle-Late Miocene northern Libya. Abstract, 8th Congr. Reg. Comm. Mediterranean Neogene Stratigraphy, pp. 209. E1 Hawat A.S. and Shelmani, M.A., 1993. Short notes and guidebook on the geology of A1 Jabal al Akhdar, Cyrenaica, NE Libya. Earth Science Society of Libya, Tripoli, p. 70. El Hawat, A.S., Missallati, A., Bezan, A. and Taleb, T., 1997. The Nubian Sandstone in Sirt Basin and its correlatives. In: M.J. Salem, A.J. Mouzughi and O.S. Hammuda (Editors), The Geology of Sirt Basin, Vol. I. Elsevier, Amsterdam.

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A.S. E L HAWAT Jenkins D., 1990. North and central Sinai. In: R. Said (Editor), The Geology of Egypt. A.A. Balkema, Rotterdam, pp. 361-380. Karcz, I. and Zak, I. 1968. Paleocrrents in the Triassic sandstone of Arayif-En-Naga, Sinai. Isr. J. Earth Sci. 17: 9-15. Keeley, M.L., 1989. The Paleozoic history of the Western Desert of Egypt. Basin Res., 2: 35-48. Khalil, N.A., Young, D. and Nairn, E.M. 1983. Paleozoic and Mesozoic stratigraphy and oil potential of Western Desert. Am. Assoc. Pet. Geol. Bull., 67: 495. Klerkx, J., 1980. Age and metamorphic evolution of the basement complex around Jabal A1 Uwaynat. In: M.J. Salem and M.T. Busrewil (Editors), The Geology of Libya, Vol. III, Academic Press, London, pp. 901-906. Klitzsch, E., 1970. Die Strukturgeschichte der Zentralsahara: Neuerkenntnisse zum Bau und zur Palaogeographie eines Tafellandes. Geol. Rundsch., 59(2): 459-527. Klitzsch, E., 1979. Zur Geologie Gilf Kebir Gebietes in der Ostsahara. Clausthaler Geol. Abh., 30:113-132. Klitzsch, E., 1983. Paleozoic formations and Carboniferous glaciation from Gilf Kebir-Abu Ras area in southwestern Egypt. J. Afr. Earth Sci., 1: 17-19. Klitzsch, E., 1984. Northwestern Sudan and bordering areas: Geological development since Cambrian times. Berl. Geowiss. Abh., A, 50: 23-45. Klitzsch, E., 1986. Plate tectonics and cratonal geology in northeast Africa (Egypt, Sudan). Geolog. Rundsch., 75: 755-768. Klitzsch, E., 1990a. Paleogeographical development and correlation of continental strata (former Nubian sandstone) in northeast Africa. J. Afr. Earth Sci., 10: 199-213. Klitzsch, E., 1990b. Paleozoic. In: R. Said (Editor), The Geology of Egypt. A.A. Balkema, Rotterdam, pp. 393-407. Klitzsch, E., Harms, J.G., Lejal-Nicol, A. and List, F.K., 1979. Major subdivisions and depositional environments of Nubian strata, SW Egypt. Am. Assoc. Pet. Geol. Bull., 63: 967-974. Klitzsch, E. and Lejal-Nicol. A., 1984. Flora and fauna from strata in southern Egypt and northern Sudan. Berl. Geowiss. Abh., A, 50: 47-79. Klitzsch, E. and Squyres, C.H., 1990. Paleozoic and Mesozoic geological History of northeastern Africa based upon new interpretation of Nubian strata. Am. Assoc. Pet. Geol. Bull., 74(4): 1203-1211. Klitzsch, E. and Wycisk, P. 1987. Geology of the sedimentary basins of northern Sudan and bordering areas. Berl. Geowiss. Abh., A, 75(1): 97.136. Kostandi, A.B., 1959. Facies maps for the study of the Paleozoic and Mesozoic sedimentary basins of the Egyptian region. 1st Arab Petroleum Congr., Cairo, 2: 54-62. Lewy, Z. 1975. The geological history of southern Israel and Sinai during the Coniacian. Isr. J. Earth Sci., 24: 19-43. Marzouk, I., 1974. Development of the ancient Nile Delta in space and time and its bearing on oil and gas potentialities. Abstract, 12th Annu. Meet. Geol. Soc. Egypt, Cairo, p. 5. Meneisy, M.Y., 1990. Vulcanicity. In: R. Said (Editor), The Geology of Egypt. A.A. Balkema, Rotterdam, pp. 157-184. McGarva, A.M., 1986. Sedimentology, petrology and diagenesis of the Zeitoun Formation (Devonian) of the Western Desert, Egypt. EGPC VIII Exploration Conf., Cairo. Morgan, P., 1990. Egypt in the framework of global tectonics. In: R. Said (Editor), The Geology of Egypt. A.A. Balkema, Rotterdam, pp. 91-111. Moustafa, A.R. and Khalil, M.H., 1995. Superposed deformation in the northern Suez rift, Egypt: relevance to hydrocarbons exploration. J. Pet. Geol., 18/3: 245-266. Neev, D., 1975. Tectonic evolution of the Middle East and the Levantine basin (easternmost Mediterranean). Geology, 3: 683-686.

S E D I M E N T A R Y BASINS OF EGYPT: AN O V E R V I E W OF DYNAMIC S T R A T I G R A P H Y Neev, D., 1977. The Pelusium Line w a major transcontinental shear. Tectonophysics, 38: 1-8. Omara, S.M. and Sanad, S., 1975. Rock stratigraphic and structural features of the area between Wadi el Natrun and the Maghara Depression (Western Desert, Egypt). Geol. Jahrb., 16: 45-63. Patton, T.L., Moustafa, A.R., Nelson, R.A. and Abdine, S.A., 1994. Tectonic evolution and structural setting of the Suez Rift. In: S. London (Editor), Interior Rift Basins. Am. Assoc. Pet. Geol. Mem. 59: 9-55. Philobbos, E.R. and Keheila, E.A., 1979. Depositional environments of the Middle Eocene in the area southeast of Minia, Egypt. Ann. Geol. Surv. Egypt, 9: 523-550. Philobbos, E.R. and Purser, B.H. (Editors), 1993. Geodynamics and sedimentation of the Red Sea - Gulf of Aden rift system. Geol. Soc. Egypt, Spec. Publ. No. 1,456 pp. Pomeyrol, R., 1968. "Nubian Sandstone". Am. Assoc. Pet. Geol. Bull., 52: 589-600. Purser, B.H., Soliman, M. and M'Rabet, A., 1987. Carbonate, evaporite, siliciclastic transitions in Quaternary rift sediments of north-western Red Sea. Sediment. Geol., 53: 247-267. Richter A. and Schandelmeier, H., 1990. Precambrian basement inliers of Western Desert geology, petrology and structural evolution. In: R. Said (Editor), The Geology of Egypt. A.A. Balkema, Rotterdam, pp. 185-200. Rizzini, A., Vezzani, E, Cococcetta, V. and Milad, G., 1978. Stratigraphy and sedimentation of a Neogene-Quaternary section in the Nile Delta area. Mar. Geol., 27: 327-348. Ryan, W.B.E, 1978. Messinian badland on the southeastern margin of the Mediterranean Sea. Mar. Geol., 27: 249-363. Said, R., 1962. The Geology of Egypt. Elsevier, Amsterdam, 377 PP. Said, R., 1971. Explanatory Notes to accompany the geological map of Egypt. Geol. Surv. Egypt, Pap. 56, 123 pp. Said, R., 1981. The Geological Evolution of the River Nile. Springer-Verlag, Berlin, 151 pp. Said, R. (Editor), 1990. The Geology of Egypt. A.A. Balkema, Rotterdam, 734 pp. Salem, R., 1976. Evolution of Eocene-Miocene sedimentary patterns in parts of northern Egypt. Am. Assoc. Pet. Geol. Bull., 60: 34-64. Schandelmeier, H., Klitzsch, E., Hendriks, F. and Wycisk, P., 1987. Structural development of north-east Africa since Precambrian times. Berl. Geowiss. Abh., A, 75.1: 5-24. Schurmann, H.M.E., 1971. Western Desert. In: Tectonics of Africa. UNESCO, pp. 423-427. Scott, R.W. and Govean, F.M., 1985. Early depositional history of rift-basin: Miocene in Western Sinai. Palaeogeol. Palaeoclimatol. Palaeoecol., 52: 143-158. Seilacher, A., 1983. Upper Paleozoic trace fossils from the Gilf Kebir-Abu Ras area in southwestern Egypt. J. Afr. Earth Sci., 1: 21-44. Seilacher, A., 1990. Paleozoic trace fossils. In: Said, R. (Editor), The Geology of Egypt. A.A. Balkema, Rotterdam, pp. 649-672. Selley, R.C., 1968. Near-shore marine and continental sediments of the Sirte Basin, Libya. Q. J. Geol. Soc. London, 124: 419-460. Sellwood, B.W. and Netherwood, R.E., 1984. Facies evolution in the Gulf of Suez area: sedimentary history as an indicator of rift initiation and development. Modem Geol., 9: 43-69. Sestini, G., 1984. Tectonic and sedimentary history of NE African margin (Egypt/Libya). In: J.E. Dixon and A.H.F. Robertson (Editors), The Geological Evolution of the Eastern Mediterranean. Blackwell Scientific Publ., Oxford, pp. 161-175. Shelmani, M., Thusu B. and Amauti, A., 1992. Subsurface oc-

85

curences of Middle and Upper Triassic sediments in Northeast Libya. In: A. Sadek (Editor), Geology of the Arab World, Vol. 2, Cairo Univ., pp. 233-240. Smith, A.G., 1971. Alpine deformation and the oceanic areas of the Tethys, Mediterranean and Atlantic. Bull. Geol. Soc. Am. 82: 2039-2070. Smith, A.G., Hurley, A.M. and Briden, J.C., 1981. Phanerozoic Palaeocontinental World Maps. Cambridge Univ. Press, 102 PP. Snavely, P.D., 1979. Regional depositional history of the Lower Eocene Thebes Formation, Eastern Desert, Egypt. Abstract, 5th Conf. Afr. Geol., Cairo, pp. 85-86. Soliman, S.M. and E1 Badry, O., 1970. Nature of Cretaceous sedimentation in Western Desert, Egypt. Am. Assoc. Pet. Geol. Bull., 54: 2349-2370. Soliman, S.M. and E1 Fetouh, M.A., 1970. Carboniferous of Egypt: isopach and lithofacies maps. Am. Assoc. Pet. Geol. Bull., 54:1918-1930. Strougo, A., 1986. The Velascoensis event: a significant episode of tectonic activity in the Egyptian Paleogene. Neues Jahrb. Geol. Palaeontol. Abh., 173(2): 253-269. Strougo, A., Bignot, G. and Abd-Allah, A.M., 1992. Biostratigraphy and paleoenvironmants of Middle Eocene benthic foraminiferal assemblages of north central Eastern Desert, Egypt. M.E.R.C. Ain Shams Univ., Earth Sci., Ser., 6: 1-12. Strougo, A., Bignot, G., Boukhary, M. and Blondeau, A., 1990. The Upper Libyan (possibly Ypresian) carbonate platform in the Nile Valley, Egypt: biostratigraphic problems and paleoenvironments. Rev. Micropaleontol., 33(1): 54-71. Thiebaud, C.E. and Robson, D.A., 1979. The geology of the area between Wadi Wardan and Wadi Ghardel, East clysmic rift, Sinai, Egypt. J. Pet. Geol. 1: 63-75. Vail, P.R., Michum, R.M. and Thompson, S., 1977. Seismic stratigraphy and global changes of sea level. In: C. Payton (Editor), Stratigraphic Interpretation of Seismic Data. Am. Assoc. Pet. Geol. Mem., 26: 83-97. Van Houten, F.B., 1980. Latest Jurassic-Early Cretaceous regressive facies, N.E. African craton. Am. Assoc. Pet. Geol. Bull., 64: 857-867. Van Houten, EB., Bhattacharyya, D.P. and Mansur, S.E.I., 1984. Cretaceous Nubia Formation and correlative deposits, Eastern Egypt: Major regressive-transgressive complex. Geol. Soc. Am. Bull., 95: 397-405. Viterbo,. I., 1968. Lower Cretaceous Charophyta from the subsurface "Nubian Complex" of the Sirte Basin, Libya. Proc. 3rd. Afr. Micropal. Coll., pp. 393--402. Wahab, S.A. and Khalifa, M.A.G., 1984. Sedimentology of the Middle Eocene Minia and Samalut Formations, West Beni Mazar, Western Desert, Egypt. J. Afr. Earth Sci., 21: 341-350. Ward, W.C. and McDonald, K.C., 1979. Nubia Formation of Central Eastern Desert, Egypt - - Major subdivisions and depositional setting. Am. Assoc. Pet. Geol. Bull., 63: 975-983. Weissbrod, T., 1969. The Paleozoic of Israel and adjacent countries. Bull. Geol. Surv. Israel, 47: 1-32. Whiteman, A.J., 1970. Nubian Group: origin and status. Am. Assoc. Pet. Geol. Bull., 54: 522-526. Wray, J.L., 1985. Miocene subsidence history and depositional facies, Nile Delta, Egypt. Abstract, 8th Congr. Reg. Comm. Mediterranean Neogene Stratigraphy, Budapest, pp. 618. Wycisk, P., 1987. Contributions to the subsurface geology of the southern Dakhla Basin and Misaha Trough (S. Egypt-N. Sudan). Berl. Geowiss. Abh., 75(A), 1: 137. Youssef, E.A.A., 1988. Sedimentological studies of Neogene evaporites in the northern Western Desert, Egypt. Sediment. Geol., 59:261-273. Youssef, M.I., 1957. Structural pattern of Egypt and its interpretation. Am. Assoc. Pet. Geol. Bull. 54: 601-614.

Chapter 5

The Iullemmeden Basin

R.T.J. MOODY

INTRODUCTION The Iullemmeden are a federation of Touareg peoples who occupy the central region of Niger (Dikouma, 1990). The name was first given to the basin by Radier (1953) who effectively redefined the area previously described by Furon (1935) as the Bassin de L'Oued Azazouk. The almost circular nature of the Mesozoic-Tertiary Iullemmeden Basin is emphasised by the outcrop of the Palaeozoic sediments to the north of the basin, in the Tim Merso]" Basin and Tin Serririne Syncline (Figs. 13). Their presence may indicate an earlier phase of deposition in a precursor basin, but the presence of Silurian graptolitic shales, similar to those throughout North Africa, suggest a broader geographic setting. According to Dikouma (1990) the borders of the Iullemmeden are clearly defined by the igneous masses of the A'fr, Hoggar and Adrar des Iforas to the north, the Bouclier BEnino-Nigerien to the south, the A'fr-Damagaram axis to the east and the Niger Fault to the west.

STRUCTURAL SETTING The Iullemmeden Basin has been described variously, as a cratonic basin (Betrand-Safarti et al., 1977); a graben developed along N W - S E basement lineaments (Wright et al., 1985); and as a half graben-synclinal structure, truncated to the southwest by NW-SE to N-S fault scarps by Guiraud et al. (1987). Windley (1978) commented on the strong structural and radiometric correlations that exist between the pre-2000 mya cratons of West Africa and South America. He also reviewed the origin of the Benue Trough, a failed arm of a Cretaceous triple junction, and noted the close association of the developing trough and the widespread regional doming south of

the embryonic Iullemmeden Basin. The origin of the Iullemmeden Basin may be linked with an elastic, or viscoelastic induced flexure of the lithosphere (Beaumont and Sweeney, 1978) resultant of the sediment and water load that collected in the Benue Trough and Bilma graben structures; the resultant surface depression exceeding the width of the initial graben fivefold. It would appear however, that the main area of the Iullemmeden Basin began to subside during the Permo-Triassic with pre-Upper Cenomanian, nonmarine sediments localised west of the Air Massif. The so-called "Bassin Continental Intercalcaire" of Greigert and Pougnet (1967) is the likely precursor to a broader, predominantly marine basin of the late Cretaceous-early Tertiary. The Benue Trough is an elongate structure which probably extends into the southern area of the Chad Basin (Matheis, 1976; Genik, 1993) (Fig. 1). There is little doubt that the structural style and depositional controls within the Benue Trough-Chad Basin areas are associated with rifting. Deposition within the Trough probably commenced during the Albian (Avbovbo et al., 1986). If so, then the opening of the Benue Trough postdates the early phases of Mesozoic terrestrial deposition in the Iullemmeden Basin defined by Greigert and Pougnet (1967). The major marine transgression at the end of the Cenomanian however, is common to both. The gentle, prolonged downwarping of the Iullemmeden Basin during the Upper CretaceousLower Tertiary is characterised by the widespread continuity of deposits. Doming is also evident to the north of the Iullemmeden, with the northern domes associated with major fault systems. Three main fault trends are recorded within the basin (Fig. 2). The most prominent trend NNE-SSW (Kogbe, 1981) and bound the Talach and Azaouak depressions and the In-Guezzam Horst. W S W - E N E trending faults are found mostly to the west of

African Basins. Sedimentary Basins of the World, 3 edited by R.C. Selley (Series Editor: K.J. Hs~i),pp. 89-103. 9 1997 Elsevier Science B.V., Amsterdam. All rights reserved.

90

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Fig. 1. Major sedimentary basins of west Africa (after Dikouma, 1990). the Air Massif. They may be reactivated along or parallel to the Guinea-Nubian lineament zone which cuts through the centre of the basin and marks the southern tip of the Air Massif. Faults with a N W - S E trend mark the western limit of the basin. The fault bounded Gao-Ansongo Trough, within the Sudanese Strait, is regarded as a N W - S E trending graben feature (Wright et al., 1985). Like the Benue Trough it is sediment-filled fault-bounded structure. It tilts toward the southwest and the Cretaceous infill attains a thickness of almost 4000 m (Guiraud et al., 1987). To the east the Iullemmeden and Chad basins (Fig. 1) are separated by the Tegama High. Both basins are complex or polycyclic structures, which result from the interplay, at intervals, of basement fractures. East of the Tegama High the Cretaceous-Tertiary sediments are associated with a network of grabens in

which the sediment fill exceed 10,000 metres (Genik, 1993) To the northeast the margins of the Air Massif are defined by the megashear zones. The N - S , NW, NE and E directed faults were active during the Palaeozoic and are linked with the alignment of the magmatic intrusions. They also influence the development of the basin. The Palaeozoic ring complexes of the Air are associated with the N80~ faulting and Precambrian wedges to the south of the Air are associated with the dextral shear zones which continue into the T~n~r6. An intra-Eocene compressive event is recorded in the Tilemsi-Timetrine area of northeastern Mali (Bellion and Guiraud, 1988). A Palaeocene-Eocene compressive phase probably affected other grabens to the east of the basin, as Avbovbo et al. (1986) provide evidence of the initiation of secondary struc-

THE IULLEMMEDEN BASIN

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.

.

Fig. 2. Simplified structural map of the Iullemmeden Basin. (after Dikouma, 1990). tures and the redistribution of stress/fracture fields in the Maiduguri area of northeastern Nigeria. An angular unconformity recorded in the Termit Basin between the Upper Cretaceous and Lower Eocene strata (Bellion et al., 1983) and evidence of slight folding and some faulting in the oolitic series (Lower Eocene) is an indication of compressive forces in eastern Niger. Some extension has occurred more recently in the Air with the development of basaltic lava flows and cones. Minor faulting controls the development of the Continental Terminal in the west and the direction of the east-west branch of the Gao-Ansongo Trough (Fig. 2).

STRATIGRAPHY AND SEDIMENTOLOGY Precambrian

The Precambrian basement rocks of the Iullemmeden Basin are well documented in the works of Greigert and Pougnet (1967) and Black et al. (1967).

They are divided into three units. In the Liptako region of western Niger, the basement is exposed at the surface forming a vast peneplain, 200-300 m above sea level. The geology is relatively straightforward with three geosynclinal units orientated NW-SE, separated by granitic intrusions. The upper part of the Precambrian in the Liptako Region is comprised of a sequence of greywackes, arkoses, cross-stratified quartzites and massively bedded conglomerates, characteristic of a molasse basin (Greigert and Pougnet, 1967). In the Niger Valley, the Lower Precambrian is restricted to the region of Gaya. There, gneisses of the Dahomeyian Series crop out in the fiver valley. They are overlain by quartzites and mica schists of the Atacorian and Buen Series of Middle and Upper Precambrian age. The Voltaien of western Niger consists of the Oti Shale Formation and the Bombouaka Sandstones. According to Greigert and Pougnet (1967), there is a direct lithostratigraphic association between the Voltaien and Gourma Infracambrian sediments.

92

R.T.J. MOODY

Fig. 3. Simplified geological map of the Iullemmeden Basin (after Dikouma, 1990).

Palaeozoic The lithostratigraphy of the Palaeozoic strata that crop out in the Tim Merso'f Basin on northern/ northeastern margin of the lullemmeden Basin is described by Greigert (1966). The Lower Palaeozoic, Cambro-Ordovician sediments are essentially marginal marine. They rest unconformably on the Precambrian basement and are grouped into the GrEs de Timesgueur, the Gr6s h Tigillites and the GrEs d'In Azaoua. They are overlain by the so-called Schistes ?~ graptolites which are the lateral equivalent of the Silurian Tanesouft Shales of North Africa. The overlying Devonian sediments are also predominantly sandstones with nonmarine deposits characteristic of the Middle and Upper Devonian (Greigert and Pougnet, 1967). A marine transgressive episode is recorded during the Vis6an and Namurian with a general regressive episode recorded during the Westphalian. Uplift and folding occurred during the Permian. Comparable deposits of Palaeozoic age exist in the neighbouring T~n6rb. du Tamesna and Djado basins.

Permo-Triassic In the area to the west of the Air, Permo-Triassic sediments are represented by a maximum 300 m

of sandstones. These are massively bedded, locally cross-stratified, with vertebrate remains and silicifled wood found toward the top of the sequence. The sandstones are referred to the Izegouandane Formation. Towards the north, the Permo-Triassic sediments rest unconformably on Westphalian nonmarine deposits. To the south they are overlain by coarse grained sandstones and conglomerates of the Alcarcesse Member.

Jurassic Cretaceous The post-Triassic, pre-Upper Cenomanian deposits of the Iullemmeden Basin are grouped under the descriptive term "Continental Intercalcaire" (Greigert, 1966) which is currently divided into the Agadez, Irhazer and Tegama Groups. The Agadez Group is divided into the Goufat Series and the Wagadi Series. The Goufat Series fines upwards from the Alcarcesse conglomerates through the Teloua Sandstone Formation into the Mousseden Formation which is comprised of shales with lenses of arkosic sandstones and grits. Large-scale dune bedding characterise the Teloua sandstones, with intercalated varicoloured shales. A total thickness of 250 m for the Goufat Series is recorded to the west of the Air. The base of the Wagadi Series is

THE IULLEMMEDEN BASIN marked by conglomerates, grits and sandstones of the Tchirezrine Formation. Together with the shales and intercalated sandstones of the overlying Abinky Formation, they constitute the Upper Jurassic component of the Agadez Group. The Wagadi Series is a fining upward sequence. The sandstones, usually red to grey, occur in stacked, planar cross-bedded sets which represent a series of migrating dune bedforms in a shallow braided river system. The extensive sheet geometry, and general coarse grain size, is related to a rapidly rising source area associated with an early rifting phase in the West African region. This sequence represents a steady fluviatile discharge, characteristic of a basin with variable margin topography. As the regional elevation decreased, the braided stream system was superseded by one of meandering streams with flood plain siltstones within fining-upwards cyclothems. Rootlet-beating palaeosols are common in the Abinky Formation. Lower Cretaceous

The Lower Cretaceous Assouas and Tchirezrine Formations (Dabla Series) are similar to the Abinky Formation in that lenses of feldspathic sandstones are found intercalated with red-brown shales. Both represent the persistence of a meandering river system across the eastern region of the basin. The Lower Cretaceous of the Central and Northeastern lullemmeden Basin is divisible into the Irhazer and Tegama Groups. The Neocomian aged Irhazer Group sediments are predominantly shales with localised bars of fine grained calcareous sandstones represented by the Azenak and Kouichi Members. Around In Gall, Irhazer Group sediments are approximately 380 m thick. In the centre of the basin the Tiouraren Formation (Moody and Sutcliffe, 1990), representative of the upper part of the group, consists of a cyclic series of evenly bedded and laterally extensive siltstones, marls and fresh water limestones. Each cycle, which can vary from 1-3 m in thickness, consists of a variegated red-purple-green siltstone capped by either a thin calcareous siltstone or calcarenite. Planar laminae and small scale fining-upward sequences are present throughout the formation. The thin capping units of each cyclothem (5-30 cm) are either carbonate-rich siltstones or ostracod packstones and grainstones. Algal or fungal stromatolites are developed within specific capping rocks; they are occasionally brecciated or silicified. Desiccation cracks and pedogenic features, such as isopachous phreatic cements, caliche glaebules and calcite-filled rhizoids are also observed. Occasional channel sands and gravels outcrop locally, mostly toward the base of the formation. The fine grained nature and extensive lateral distribution of the siltstones indicate a low energy fluviatile-

93 lacustrine depositional environment. The indication is of a denuded area with little topographic variation. The upper part of the Irhazer Formation was deposited in a meandering, shallow river system dominated by overbank silt deposition. Shallow lakes rich in ostracods, with soils developed in waterlogged areas as a result of temporary pulses. The shallow lakes were widespread, their waters replenished by overbank flooding from nearby river channels. These channels, as indicated by the dimensions of the preserved channel systems, were less than 2 m deep and 50 m wide. They are characterised by shallow, poorly developed, cross-sets and lag deposits with an abundance of transported, slightly abraded vertebrate material. The indication is of relatively sluggish water movement. Strong comparisons exist between the vertebrate faunas of the Irhazer deposits and the Apto-Albian Santa and Ilhas Formations of Brazil (Maisey, 1987; Carvalho, 1982). Very similar environmental conditions to those proposed for the Irhazer Formation (Moody and Sutcliffe, 1990) are recorded by Fastovsky (1987) and Wood et al. (1988) for the Upper Cretaceous formations of eastern Montana, North Dakota and Alberta. In the case of the Alberta, Judith River Formation, the bankfull depths are much greater than those found in the Irhazer Group. It may thus be inferred that the mean annual discharge and mean flood discharges were considerably lower than those of Alberta, although the taphonomic interpretation for the deposition of vertebrate fossils would be much the same. The outcrop of the Irhazer shales and of the overlying Tegama Group (Fig. 4) basal sandstones extends in a half circle from the Tin Merso'f area of north-central Niger southwards through In Gall and eastwards to the southern tip of the Air. The boundary between the two groups is marked by the Falaise de Tiguedi. Taquet (1976) divided the Tegama Group into eight members. These were established on outcrops described between Agadez in the northwest and Egaro/Termit to the southeast. Baudet et al. (1981) erected a revised lithostratigraphy for the central region of Niger in which the Tegama Group is divided into upper and lower series and nine formations. According to Taquet (1976), the sediments of the Tegama Group range from Barremian to Albian in age. The total thickness of the Tegama Group in central Niger is approximately 700 m (Baudet et al., 1981), whereas to the northeast, in the region of In Abangharit, the total thickness is less than 60 m. The Tegama Group also thins (50 m) northwards into the du Tamesna. To the west the various formations defined by Baudet et al. (1981) are less distinct. The lower part of the Tegama Group is called the Terzikasan Formation. It is, in part, equivalent to

94

R.T.J. MOODY MALl (MOODY& SUTCLIFFE,1991}

LITHOSTRATIGRAPHY CENTRAL IULLEMMEDEN BASIN (DIKOUMA, 1990)

NIGERIA (KOGBE19111)

Middle Niger

clay rich sandstones(CT3)

BIRNI N'KONI FORMATION

Gwandu Formation

Lignitic clay rich sandstones CT2)

EOCENE ADER DOUTCHI FORMATION

Tamaguilelt Formation

Ader Doutchi aiderolithic series (CT1)

9

Barmou Member Paper Shales GARADOUA SERIES

PALAEOCENE

Tamaake Member Linthia Limestone

'

Teberemt Formation 9

Kao Member Paper Shales

IN WAGAR FORMATIO

Saraliguedad Formation

Chair Keini Formation

Gamba Formation

Kalambeina Formation '

eL

0 0

Dange Formation Wurno Formation

:~ . 0 er

Dukamaje Formation

FARIN DOUTCHI FORMATION

MAASTRICHTIAN

ALANBANYA FORMATION SENONIAN (LOWER& UPPER)

,

UPPER TURONIAN

SENONIAN " ~ SERIES ~ 'WHITE ~

LIMESTONE'

"~

Do=chin 7 , n a ~ Formation ~

.~

.~

~

ii "!"

~

LOWER CENOMANIAN

i r <

.x

E

Taloka Sandstone Formation

I o

LOWER CENOMANIAN

TEGAMA GROUP

~>

m

Fig. 4. Lithostratigraphic correlation of the Upper Cretaceous-Tertiary deposits of the Iuilemmeden Basin (after Dikouma, 1990; Kogbe, 1981 and Moody and Sutcliffe, 1991).

the Tiguedi Series of Faure (1966), which comprises sandstones with interbedded shales, and crops out in the Damergou Hills, north of Zinder. The Terzikasan Formation of Baudet et al. (1981) is well exposed in the Falaise de Tiguedi near In Gall. The Falaise de Tiguedi is a well defined element of the landscape and separates the Pays de Tegama in the east, from the Adar Doutchi. The Terzikasan Formation is the lateral equivalent of Gad. 1 erected by Taquet (1976) for the Gadouafaoua area. The main outcrop of the Terzikasan Formation is composed of 20 m of crossstratified quartzitic and arkosic arenites. These are poorly sorted and predominantly fine grained. They are trough and planar cross-bedded with sets ranging from 7 cm-70 cm. The mean palaeocurrent direction is westwards with a low variance. The majority of sets show truncation at their upper boundaries; some exhibit water escape structures toward this boundary, where the oversteeping of the foresets is most common. The upper part of the Terzikasan Formation is characterised by the incoming of siltstone interbeds. Fining-upward cycles are common, each cycle having a cross-bedded sandstone at the base which grades into red to green siltstones. Silicified wood, including trunks 30-40 cm in diameter, are abundant in the siltstones. The lower parts of the Terzikasan Formation were deposited in a sandy braided fiver environment, with a steady discharge, producing a series of linguoid,

sinuous and straight crested bars that migrated downstream to give stacked planar cross-bed sets typical of the type model of Miall (1977). The character of the upper part of the Terzikasan Formation infers a change from braided to meandering fiver systems, which implies either a decrease in the elevation of the source area, or a rise in sea-level to alter the gradient of the fiver profile. The overlying Amezroun, Anyeli and Tin Sakan Formations (Baudet et al., 1981) are the lateral equivalents of Gad. 2 (Taquet, 1976). They are predominantly quartzitic and arkosic arenites with intercalated variegated shales. The overall pattern of sedimentation is a fining upward megasequence; a similar interpretation is likely for the overlying Guirmaga (Gad. 3) and Mohra (Gad. 4-5) formations. Each megasequence argues for a change from braided to meandering fiver systems. The combined thickness of the sequence is approximately 200 m. The overlying Tin Anasin Formation (Gad. 6) is dominated by medium to massively bedded sandstones with siltstones and shales towards the top of the sequence. The lower sandstone unit is over 100 m thick and is interpreted as a stacked series of sand bars. Massive sandstone and conglomerate units are also characteristic of the Ekismane (Gad. 7) and Borak (Gad. 8) Formations. The base of the Ekismane Formation is marked by the occurrence of some 60 m of cross-bedded quartzitic arenites with an abun-

THE IULLEMMEDEN BASIN dance of silicified wood. In many respects, this unit is similar to the Terzikasan deposits, indicating a braided stream environment. Shales and fine grained sandstones separate the basal unit from 60 m of massively bedded conglomerates. These conglomerates have a highly erosive base which truncates the underlying sequence, suggesting an increase of energy within the fluvial system. The truncation and the sudden influx of conglomeratic material indicate a reactivation of the source area. The upper 50 m of the Ekismane Formation consists predominantly of shales with sandstone channel deposits. The shales represent floodplain deposits and the sandstones relate to increased channel sinuosity and the return to a meandering river system. Shales with channel sandstones are characteristic of the overlying Borak Formation (Gad. 8), although the greater presence of clastic sediments infers a likely braided river influence. Similar sedimentary sequences in the Cenomanian En Nassame and Berere Members,which total 100 m in thickness, relate to the continual uplift of the region throughout the Lower Cretaceous. In the Southern Iullemmeden Basin, Lower Cretaceous deposits also crop out in the Sokoto region of Nigeria, commonly known as the Sokoto Basin. There the Continental Intercalaire is represented by the Jurassic-Lower Cretaceous Illo and Gundumi Formations (Kogbe, 1981). The Gundumi Formation comprises a sequence of arkosic, argillaceous grits, clays and coarse pebbly sands which are locally conglomeratic. The maximum thickness of this formation near Bakura, east of Sokoto, is approximately 320 m, but it thins rapidly away from the type section, with only the upper units represented toward the west. The contact between the Gundumi Formation and basement rocks is conglomeratic. The conglomerates cover a large area around Talata Mafara in the southeast of the basin and mark the advent of a braided stream system. The Illo Formation crops out in the southwest quadrant of the Sokoto Basin. They include continental, fluviatile and fluvio-lacustrine deposits, and are rarely thicker than 40 m. The type section is near Gore, approximately 10 km northeast of I11o. The basal unit is composed of poorly sorted grits and coarse sandstones. These are friable sediments with planar cross-beds and are associated with a braided fluviatile environment (Kogbe, 1981). Both the Gundumi and Illo Formations contain silicified wood belonging to various species of the genus Dadoxylon (Araucarioxylon). Other genera include

Dadoxylon (?Protopodocarpoxyolon), Embergerixylon and Mesembrioxylon (Kogbe and Lemoigne, 1972). A similar flora is recorded from the central Iullemmeden Basin and North Africa (Boureau, 1958) and from the Numidian Sandstone of northeast Niger (Djado Basin) by Greigert and Pougnet

95 (1967). The flora does not lend itself to zonal use and the general term Wealden is applied to the Lower Cretaceous deposits of the Sokoto Basin. To the northwest in the Gao-Ansongo Trough, the upper part of the "Continental Intercalaire" has been dated as Albian to Lower Cenomanian (Trofimov et al., 1969). The sporopollen obtained from the well Bourem-1 included an abundance of Perotriletes sp. and, more rarely, Anemiasporites sp., Leiotriletes sp., Classopollis sp., Steevesipollenites sp. and Tricolpopollenites sp. To the east of Bourem, the Ourtufoulut well (Merabet et al., 1971) approximately 300 km north of Niamey, cuts through a sequence of microconglomerates, argillaceous sandstones and shales of similar age to those from the Bourem well. These sediments are overlain by 27 m of shales, the flora of which is dated as Upper Cenomanian-Lower Turonian. The shales are the lateral equivalents of the Neolobites-bearing strata of the north-central quadrant of the Iullemmeden Basin. Guiraud et al. (1987) record the south-west-tilting Gao-Ansongo Trough contains 3500-4000 m of Cretaceous sediments. Marine sediments of the Upper CenomanianLower Turonian occur in the Tahabanat well, 100 km NNE of Ourtufoulut. The section recorded is 105 m thick and a diverse fauna of benthonic foraminifera, ostracods and bivalves marking the first of several Cretaceous-Tertiary transgressive episodes. The marine shales overlie sandstones and shales of the Continental Intercalaire. The thickest Lower Cretaceous sequence recorded in the central Iullemmeden area is approximately 1200 m whereas 3000 m of post-Aptian sandstones are recorded by Guiraud et al. (1987) in the Kafra Graben.

Upper Cretaceous The Upper Cretaceous deposits of the Iullemmeden can be divided effectively into four units of Cenomanian, Turonian, Coniacian-Santonian and Campanian-Maastrichtian age (Fig. 4). The Cenomanian sediments of the Iullemmeden Basin are referred to the En Nassame and Berere Members of the Talrass Formation (Baudet et al., 1981). A detailed account of the transition from nonmarine to marine Cenomanian sediments in the Tin Essako area is presented by Greigert (1966). The transition between the Lower Berere Member and the overlying Berere (Marine) Member marks the beginning of the first of several marine transgressions across West Africa. In the Iullemmeden Basin, the appearance of the belemnite Neolobites marks the presence of the Upper Cenomanian in the Iguellala Mountains. In the Tenere Desert, southeast of the

96 Air Massif, the first indication of a marine incursion, marking a connection via the Benue Trough, occurs 200 m below the Neolobites vibrayeanus lumachelle (Greigert, 1966). Upper Cenomanian shales, of the Berere (Marine) Member are exposed to the east of Tanout in the northeast region of the Damergou. The basal 15 m of the sequence is sand-rich with carbonaceous debris. These beds are overlain by a lumachelle with Neolobites vibrayeanus and Exogyra columba. A 10 m shale sequence with rare shelly interbeds separates the Upper Cenomanian lumachelle from a second, Lower Turonian, lumachelle with Nigericeras lamberti and Paravas-

coceras cauvini. The advance of the sea across West Africa during late Cenomanian times corresponds with a worldwide transgression (Moody and Sutcliffe, 1990). According to Reyment (1987a), Central Tethys invaded the Iberian Peninsula and spread across North Africa. Epicontinental extensions occurred across the Sahara and in the Middle East. At the same time, an invasion of the Benue Rift and the Gulf of Guinea stemmed from the South Atlantic. Reyment (1987a) draws comparison with the Greenhorn Cycle (T6) of the Western Interior of the United States and notes the presence of several faunal elements typical of North America in Nigerian-Niger passage beds of Cenomanian and Turonian age. The former is the equivalent of the more easterly Farak and Cheffadene formations of Greigert and Pougnet (1967). The Farak Formation is 200 m thick to the south of the Air. Traced southwards into the Fosse de Tefidet, the formation grades into marine sediments of Lower Cenomanian age. De Lapparent (1960) notes the presence of numerous dinosaurs from the Farak deposits. In the far eastern quadrant of the lullemmeden Basin, north of Bilma, sediments of similar age are referred to the Cheffadene Formation. The shales within the Cheffadene Formation commonly contain the imprints of leaves. A transitional series between the Lower and Upper Cenomanian exists in the region of In Abanghant-Tamaia (Tnr de Tamesna). Thin coal seams are evident in the sandstones, whereas freshwater bivalves such as Unio are abundant in the shales (Mongin, 1954). In the north-central region of the basin in the Iguellala Mountains, the late Cenomanian-early Turonian succession is 15 m thick and consists of shales with sand intercalations overlying the Continental Intercalaire. The shales contain a diverse fauna of bivalves and nautiloids. Neolobites vibrayeanus occurs approximately 3 m above the base of the succession, with Nigericeras sp. A mixed fauna of vascoceratid ammonites and irregular echinoids occur 1.7 m above the previous fossiliferous unit. Correlation between the Damergou and Iguellala Mountains is well established and the Exogyra olisiponensis

R.T.J. MOODY lumachelle is traced as far as Tenekert to the east of the Adrar des Iforas. The Daiet el Frass well in southern Algeria (Merabet et al., 1971) encountered Upper Cenomanian marine deposits between 780-803 m (BRT). The sediments are mostly limestones with Exogyra columba var. minor, E. olisiponensis and numerous other bivalves, gastropods and fragments of ammonites. This occurrence proves the southward incursion of Tethys, whereas the overlying Turonian shales with anhydrite are probable evidence of an early regressional phase. The presence of late Cenomanian-early Turonian littoral facies toward the Algerian border, and the persistence of braided stream environments along the southwest margin of the Iullemmeden Basin indicates the northern margin of the basin (Greigert, 1966). A local regression (late Cenomanian) as in southern Algeria, is marked by the appearance of shales with anhydrite interbeds and of massive, cross-bedded sandstone units in the north-central quadrant of the basin. In the centre of the basin around In Gall and Tahoua the Nigericeras shales and limestones (Lower Turonian) are succeeded by a monotonous sequence of shales with white limestone intercalations. The shales are variegated, being mostly red or violet with gypsum stringers. Approximately 150 m of such deposits are recorded from a well at Ibeceten, 50 km ENE of Tahoua. A sparse fauna of ostracods is recorded and the indication is of brackish water, deposition, possibly in a restricted coastal environment. A similar sequence (175 m thick) is encountered in the Tahabanat well at the eastern entry to the Gao-Ansongo Trough. The shales and limestones are Upper Turonian (Greigert, 1966) and elsewhere in the basin contain oysters, small bivalves and gastropods. In the Tamaia Massif, white limestones of the Upper Turonian yield natacid and strombid gastropods. The same fauna exists in the Tenekert region on the southeast flank of the Adrar des Iforas. The Upper Turonian sediments in the Iullemmeden Basin are collectively known as the Calcaires Blancs (lower part of the Doutchin Zana Group of Dikouma (1990), Fig. 4), and as the name suggests, are composed of white limestones. In contrast the white limestones are absent in the Damergou region to the southeast where the Nigericeras limestone is overlain by shales containing dinosaur bones. The shales with sandstone intercalations may attain a thickness of 80 m locally. The occurrence of the gastropod-rich limestones in the north central area of the basin marks the advent of the second marine transgression into the basin. This is well documented by Reyment (1987a, b) who considers the palaeoclimatic condi-

THE IULLEMMEDEN BASIN

97

Table 1. Various stratigraphic schemes for the Upper Maastrichtian

Upper Maastrichtian Campanian-Maastrichtian Lower Campanian

Jones (1948)

Parker (1965)

Radier (1957)

Upper sandstones and mudstones Mosasaurus shales Lower sandstones and mudstones

Wurno Formation Dukamaje Formation Taloka Formation

Termes III Termes II Termes I

From Moody and Sutcliffe (1990).

tions surrounding the isthmus to be similar to those of the Red Sea at the present. In contrast to the monotonous nature of the Turonian sediments, those of the early Coniacian are more varied. Limestones and sandy limestones are characteristic although sandy marls and shales are present locally. The main outcrop of the Coniacian parallels that of the Cenomanian-Turonian deposits and extends in an arc from the southern tip of the Adrar des Iforas west of Tahoua and southwestward to the Niger River. To the east of the Adrar des Iforas, a conglomerate composed of limestone clasts, containing benthonic foraminifera and bivalves, is overlain by a limestone with Chara. The indication is of a coastal depositional environment. A more diverse marine fauna is encountered to the east. In the region of In Gall and Tahoua, the Calcaires Blancs are overlain by a thin sequence of oyster-rich shales and lumachelles and black shales with fish, dinosaur, crocodile and turtle remains. The outcrop at Ibeceten, according to Greigert and Pougnet (1967) represents an intimate juxtaposition of marine and nonmarine (continental) deposits. The fish fauna includes the lungfish Polypterus and the freshwater fishes Lepisosteus and Amia. The sequence at Ibeceten is just 5 m thick whereas the marine sections to the north reach a maximum thickness of 35-40 m. In the Damergou Hills to the east, equivalent deposits to the shales at Ibeceten are overlain by approximately 20 m of shales with thin limestone partings. At the base, the shales contain abundant ostreid bivalves and gastropods whilst the top 15 m contains various fish and reptile bones. The base of the upper shales is marked by a microconglomerate rich in phosphate and glauconite. The whole of this sequence rests on a series of grits, sandstones and shales that are considered the equivalent of the Calcaires Blancs (Greigert, 1966). The Santonian (Upper Senonian) is marked by a regressive episode throughout the Iullemmeden Basin. In the Gao-Ansongo Trough, the Senonian is represented by 20-60 m of sandstones and siltstones with plant debris (Merabet et al., 1971). On the southern margin of the basin, the Santonian? is represented by a sequence of feldspathic grits and sandstones with intercalated shales. These are, in

part, the lateral equivalent of the Upper Sandstones (Jones, 1948) that border the Adrar des Iforas and were referred by Greigert (1966) to the Continental Hamadian (Fig. 4). To the south of Zinder near the Niger-Nigerian border the Continental Hamadian is approximately 130 m thick. It consists of grits and sandstones which are locally cross-bedded and argillaceous. To the west around the town of Maradi, conglomerates and arkoses with thin shale horizons pass progressively eastwards into a finer grained sequence of shales and sandstones (Greigert, 1966). The Continental Hamadian is believed to be the lateral equivalent of the upper part of the Illo Group in the Sokoto Basin. The Campanian-Maastrichtian depositional history of the Iullemmeden Basin is marked by an initial continental-lacustrine phase followed by a major transgression and the subsequent return of nonmarine conditions in the Upper Maastrichtian. Perhaps the most clearly documented stratigraphy is that of the Sokoto region (Table 1). The top of the Lower Sandstones and Mudstones are dated as earliest Palaeocene locally, whereas the Mosasaurus Shales are referred to either the Laffiteina or Libycoceras ismaeli zones of the Maastrichtian. Parker (1965) assigned these sediments to the Taloka, Dukamaje and Wurno formations which were subsequently placed within the Rima Group (Kogbe, 1979, 1981) (Fig. 4). The Taloka Formation consists of four major lithofacies, the base of the formation being characterised by 6 m of red-purple, argillaceous siltstones (Kogbe, 1981). These are overlain by approximately 5 m of whitish grey siltstones with carbonaceous stringers, 18 m of light brown, friable siltstones with claystone/carbonaceous shale intercalations and 13 m of white, poorly consolidated siltstone with a distinctive Callianassa-burrowed horizon at the top. The type section at Taloka is 60 m thick. The Dukamaje Formation consists predominantly of shales with subsidiary limestones and mudstones. The base is marked by a ferruginous bone-bed, overlain by laminated and gypsiferous shales. Locally, to the south of Dukamaje, two bone beds occur, their presence being associated with strand-line deposition (Kogbe, 1974). A thin white limestone towards the top of the formation has yielded bi-

98 valves, gastropods, corals and echinoids. The Upper Campanian-Lower Maastrichtian ammonite, Libycoceras afikpoensis (Reyment) is also recorded from the Dukamaje Formation (Kogbe and Wozny, 1980). L. afikpoensis is also recorded from the Nkporo Formation of the Calabar region in the eastern Niger Delta (Kogbe, 1981; Zaborski, 1982, 1984) and related species occur throughout Niger (Greigert, 1966). The Nkporo Formation is dated as Upper CampanianLower Maastrichtian (Zaborski, 1982, 1984). At Lokpauku, approximately 130 km northeast of Calabar, the Nkporo Formation is overlain by the Mamu Formation which is of Lower Maastrichtian age. Juvenile specimens of L. afikpoensis have been recorded from the Gombe region of northeast Nigeria. The preferred migratory route between the south Atlantic and the Iullemmeden Basin is via the Niger and Sokoto Embayments (Kogbe, 1980, 1981), although Reyment (1987b) argues that the epicontinental seaway was displaced westward during the uppermost Maastrichtian. This conclusion supports the work of Zaborski (1982, 1984) and indicates that the Benue Trough was the mostly likely migratory route during late Campanian-early Maastrichtian times. The Niger Embayment-Sokoto Embayment link was accessed only during the uppermost Maastrichtian-Palaeocene. Adjetuni and Kogbe (1986) propose a compromise solution by postulating the flooding of both areas during the Maastrichtian, leaving the Jos Plateau region as an island. The Dukamaje Formation has yielded abundant invertebrates and vertebrates (Moody and Sutcliffe, 1990). The thickness of shales increases into the centre of the Iullemmeden Basin, from 0 m, 2 km south of the Niger-Nigeria border, to over 40 m near Kao. The Upper Maastrichtian in the Sokoto Basin is represented by the Wurno Formation. This comprises approximately 20 m of massive siltstones, variegated siltstones and silt-rich clays. The general character of the Wurno Formation and the abundance of burrow systems, flaser bedding and wavy bedding (Kogbe, 1981) indicate a tidal flat environment along a NESW trending shoreline. There is little doubt that the three formations described by Parker (1965) and placed by Kogbe (1979, 1981) within the Rima Group, are the lateral equivalents of Termes I-III of Radier (1957), and Unit A of Alam (1986). In the central Iullemmeden however, the total thickness of these Upper Campanian-Maastrichtian deposits exceeds 250 m, compared with approximately 90 m in the Dukamaje district of Nigeria. Moody and Sutcliffe (1990) use the term Igdaman Group for the central Niger deposits, inferring that

R.T.J. MOODY the lack of significant lateral variation basin-wide allows the formation names applied by Kogbe (1979, 1981) to be used in Niger. In a detailed description of the Upper Campanian-Maastrichtian of the Gao region, Moody and Sutcliffe (1990, 1995) record a gradual increase in water depth from lagoon to deltaic-shallow marine conditions. Cross-stratification in the lower units of the "Mosasaurus Shales" yields southeast to east palaeocurrent directions with tidal channels cutting down through coastal lagoon and tidal flat environments. In the Ourtufoulut well, at the entrance to the Gao-Ansongo Trough, 150 m of Cretaceous strata consists of basal grits and sandstones, Mosasaurus Shales and Upper Sandstones (Merabet et al., 1971). In the east of the basin, the CampanianMaastrichtian deposits are either absent or represented by a reduced sequence of continental deposits as in the Damergou region. The Upper Maastrichtian regressive event in the Iullemmeden Basin is represented by the Upper Sandstones and Mudstones, which are marginal marine in character.

Tertiary The established lithostratigraphy for the Palaeocene of the lullemmeden Basin is based on the work of Kogbe (1972, 1981) and Dikouma (1990) (Fig. 4). The Sokoto Group consists of three formations, the type sections of which all occur in the northwest region of Nigeria. The associated sediments are marine or coastal. Radier (1957) described the Gao-Ansongo Trough succession and Greigert (1966) documents the outcrops in the main area of the Iullemmeden Basin. Termes V-VII of Radier (1957) were hitherto used in the description of these sedimentary sequences. Sufficient variation exists between the Sokoto and Iullemmeden basins to warrant the use of different formation names for the two areas. The differences are related to changes in the environment of deposition and in the increased thickness of sediments in the main basin. The lateral changes that exist between Niger and Mali are also reflected in the use of a different nomenclature. In the Sokoto Basin, the Sokoto Group (Kogbe, 1981) is unconformable on those of the Rima Group. The basal sediments of the former are referred to the Dange Formation (Parker, 1965). Locally, the boundary between the Palaeocene and Maastrichtian is erosional with a thin conglomerate containing coprolites, mollusc debris and gypsum. The formation is comprised of blue-grey shales with thin limestones, the maximum thickness being approximately 50 m. An abundant vertebrate fauna has been described by White (1934) and Halstead (1973, 1975). Borehole samples from the Sokoto region

THE IULLEMMEDEN BASIN have yielded a rich and diverse fauna of benthonic foraminifera. These are Dano-Montian in age and are referred to the Laffitteina bibensis zone (Krasheninnikov and Trofimov, 1969). The Dange Formation is overlain by the white limestones and shales of the Kalambaina Formation. The maximum thickness of this formation is 20 m. Both limestones and shales are rich in invertebrates with corals, bivalves, gastropods, echinoids and nautiloids associated with a diverse fauna of benthonic foraminifera and ostracods. The echinoids include Schizaster (Linthia) sudanensis, Echinopsis cf. friryi and Plesiolampas cf. saharae; the nautiloids are Deltoidonautilus molli and Cimonia sudanensis. Kogbe (1981) records the planktonic foraminifera Globigerina triloculinoides which, if correct (as there are few planktonics recorded elsewhere in the basin), would indicate a P1-P4 Palaeocene age (Keller, 1988). Kogbe (1981) however, claims the presence of age-diagnostic benthonic foraminifera and correlates the Kalambaina limestones with the Upper Palaeocene limestones of Mali and Niger. The Kalambaina Formation (Parker, 1965) is overlain in the Sokoto Basin by the grey laminated shales of the Gamba Formation. The shales are essentially devoid of macrofauna but contain a diverse microfauna of benthonic foraminifera and ostracods. Thin phosphatic microconglomeratic intercalations yield fish teeth and scales and the moulds of small bivalves. According to Kogbe (1981), the Gamba Shales are the lateral equivalents of the Series Argileuse of Greigert (1966). Petters (1977) refers the Kalambaina Formation to the Operculinoides bermudezi zone of Krasheninnikov and Trofimov (1969) and thus by inference, places the Gamba Formation in the Lockhartia haemei zone (Palaeocene) of the same authors. The maximum thickness of the Gamba Formation is 10 m and is capped by a Primary Ferruginous Oolite which Kogbe (1981) dates as Upper Palaeocene. The depositional history of the Maastrichtian and Palaeocene deposits of the Sokoto Basin are detailed by Alam (1986). Four major depositional units, A D, are defined which correspond to five formations studied by Parker (1965). The Maastrichtian Rima Group sediments are referred to Unit A and are interpreted as rhythmic distal wadi plain deposits. The suggestion is of sporadic sediment supply associated with mild rain storms. The rhythmic pattern is supposedly large scale, the increase in sand and in bed thickness associated with the changing direction of stream channels. Unit B corresponds to the Dange Formation of Parker (1965). This shale-rich sequence is regarded by Alam (1986) as a mudrich, sabkha deposit related to the infiltration and evaporation of water in ephemeral streams with low gradients. The environment of deposition is thought

99 to be similar to the Permian Red Cove Formation of Texas (Handford and Fredricks, 1980; Handford and Bassett, 1985). The presence of marine invertebrates and vertebrates in these shales (Moody and Sutcliffe, 1990) would argue for stagnant offshore or lagoonal conditions. Unit C of Alam (1986) corresponds to the Kalambaina Formation and is interpreted as coastal mud fiat deposits. This is based on a high mud content and the interpretation of wispy laminae as possible algal mats. The presence of abundant echinoids, corals and larger foraminifera would appear to contradict this interpretation. Alam (1986) does not refer to either the Gamba Formation or the overlying oolitic limestones. Unit D corresponds to the Gwandu Formation, which is suggested to be the product of a fluviatile environment. In the central region of the Iullemmeden Basin and along the western flank of the Adrar des Iforas, sediments of equivalent age to the Sokoto Group crop out over hundreds of square kilometres. They are either horizontal or dip slightly with excellent sections revealed in deeply eroded wadis and gullies. The Continental Hamadian defined by Greigert (1966) and described above is evidence of continued continental deposition on the margins of the basin. Boudouresque et al. (1982) date the Continental Hamadian as Upper Maastrichtian-Lower Danian from the sporopollen-dinoflagellate assemblage of

Periretisyncolpites giganteus, P. magnosagenatus, Buttinia andreevi, Ceratiopsis granulostriatum, Longapertiles microbaculatus, Retidisporites nigeriensis, Striamonocolpites anastomosus and S. microcanilis. The flora is that of a dense tropical forest with a hot, humid climate. Boudouresque et al. (1982) refer to a marine influence within the flora and state that this agrees with the palaeogeographic interpretation of Petters (1977), who argues for an inland sea connected to North Africa rather than the southern Atlantic. The most complete sections of Palaeocene strata in the Iullemmeden Basin crop out around the towns of Tahoua and Kao. The maximum exposure in outcrop is approximately 80 m, but in the Ekkineouane well, to the northwest of Tahoua, the combined thickness of the Lower and Upper Palaeocene strata is 110 m. Three lithological units can be recognised within this well, which correspond to the Dange, Kalambaina and Gamba formations of the Sokoto district. Greigert (1966) records thicknesses of 40 m, 44 m and 36 m, respectively, for the three units within the Ekkineouane well. The lowest shales of unit 1 are of Maastrichtian age and the K - T boundary appears conformable. Shales of the lowest Palaeocene outcrop at Wajee to the east of Tahoua. They are blue grey with subsidiary siltstones and white limestones, the sequence being approximately 20 m thick. The

100 basal beds yield the bones of dyrosaurid crocodiles whereas Schizaster (Linthia) sudanensis is found in the top 15 m. The Wajee Shale Formation is the lateral equivalent of the Dange shales. The white limestones that overlie the Wajee Shale Formation are the lateral equivalent of the Kalambaina Formation (Fig. 4). They are well exposed at Mont Igdaman, 4 km to the northeast of Kao, and in the scarplands around Tahoua. The contact with the Wajee Shale Formation is gradational, the lower limestones being argillaceous. A more diverse fauna of bivalves, gastropods, nautiloids and echinoids is characteristic of the lower argillaceous limestones. The presence of Schizaster (Linthia) sudanensis and Plesiolampas saharae provides direct correlation with the Kalambaina Formation. In the Iullemmeden Basin however, faunal content and petrological characteristics allow for subdivision of the limestones into 3 biolithological units. The lower argillaceous muddy limestones with Schizaster are 3-4 m thick. They are overlain locally by nodular well cemented limestones with oysters, high spired gastropods and echinoid debris. The thickness of this unit is variable but exceeds 35 m as in the Tahoua and Afsaranta water wells. The diagnostic fossil of this unit is the benthonic foraminifera Operculinoides sp. Marls and argillaceous, thinly bedded limestones can separate the Operculinoides limestones from the overlying carbonates of the Lockhartia haemei zone. The latter includes the "Calcaire blanc dur" a hard, cemented limestone, which usually exhibits well developed karstification. The upper limestones also vary in thickness but rarely exceed 2-3 m. According to Greigert (1966), the Upper Palaeocene had the most significant of the five transgressions that entered the Iullemmeden Basin. Faunas include large clams, corals, bryozoa and larger foraminifera recorded from the Operculinoides limestones to the north of Tahoua. Reefal build-ups, although limited around Tahoua, are not recorded elsewhere in the basin. Moody and Sutcliffe (1990) use the term Garadoua Formation for the Palaeocene limestones of Niger. At Wajee, 25 km east of Tahoua, the boundary between the limestones and the overlying Schistes Papyreuses (Greigert, 1966) is marked by a root bed. The shales are blue grey in colour but when weathered turn yellow brown and exfoliate. They mark the base of a regressive episode. Intercalations of phosphatic conglomeratic or thin calcareous mudstones with phosphate pellets occur at the base of the paper shales. These shales are usually unfossiliferous but the phosphatic levels can be rich in opaline fish bones. Small, thin shelled bivalves, which are monospecific and abundant, crowd the upper surfaces of the thin limestones. Single, vertical burrows, probably after annelids, indicate a

R.T.J. MOODY coastal mudflat or lagoon sabkha environment. High humidity may have been one reason for the lack of gypsum within the system. The equivalent of the Gamba Shales in the Iullemmeden Basin and the Tilemsi Valley are the In Jinjira and Tamaguilelt formations (Sutcliffe, pers. comm.). The successions in these two areas are quite distinct from those of Nigeria with a maximum thickness of 27-30 m recorded from the Tamaguilelt area, north of Gao. At In Jinjira, NNE of Tahoua, the equivalent of the Primary Ferruginous Oolite (Kogbe, 1981) is 80 cm thick and overlies a rooted and karstic surface. The oolitic ironstone/mudstone is unfossiliferous but is bioturbated, probably by organisms producing ophiomorphid-type burrows. Phosphate nodules are abundant at the base of the oolitic unit. Two metres of blue grey shale and a 50 cm bed of large phosphate nodules complete the sequence at In Jinjira. The age of these post-Gamba Formation sediments is doubtful, but phosphates of Eocene age occur above the paper shales and oolites in the Tilemsi Valley. A ferruginous oolite in the Tilemsi Valley is well exposed over a wide area. Locally it is thoroughly bioturbated, individual burrows housing the well preserved remains of callianassids shrimps. The dimensions of the burrows varies throughout the region and this may be evidence of a variation in water turbulence. At Tamaguilelt, the oolite rests on a thin oncolitic horizon. The indication is of a shoreline succeeded by a shallow open water regime. Abundant remains of bony fish, turtles and crocodiles are recorded from the oolite. They include a number of pycnodont species such as Pycnodus bowerbankii, a form previously known from the London Clay of Sheppey, SE England. In contrast to the central Iullemmeden Basin and the Sokoto region, the evidence from the Tilemsi region indicates that the oolites become diachronous towards the east and are indicative of a small-scale incursion of the last Saharan sea during the Ypresian. At Tamaguilelt and Samit, the oolite may rest on the equivalent of the Kalambaina limestones. It is in turn overlain by 4-8 m of paper shales with sandstone, phosphate and muddy limestone intercalations. According to Alam (1986), the shales would be formed in a mud-rich sabkha, the oolite limestones being deposited on a coastal bar. Phosphates are common throughout the Tilemsi area. At Tamaguilelt, 5 m of pelleted phosphates overlie the first paper shale unit. The phosphates are essentially composed of reworked coprolites and faecal pellets, the larger droppings exceeding 10 cm in length. Abundant remains of catfish, pycnodonts, turtles, crocodiles and sea-snakes are associated with the phosphates. The majority are disarticulated and have been transported. Lung fish and

THE IULLEMMEDEN BASIN more rarely the bones of the mammal Moeritherium are found, which indicate a near-shore environment. Few sedimentary structures characterise these phosphates but an overall coarsening-upwards sequence is recorded. The thickness of the phosphate unit can show considerable lateral variation within a few kilometres and they may represent broad channel-fill deposits. Paper shales or bauxitic variegated red/green shales may overlie the main phosphate units. Thin pelleted phosphatic sands occur as local intercalations. The total thickness of shales and mudstones overlying the main phosphate in the Tilemsi Region are approximately 10 m thick. Complex ophiomorphid burrow systems occur in the thicker siltstone/mudstone units. However, the gradual incoming of brown mottled palaeosols mark the end of the widespread influence of marine or inland sea environments within the Iullemmeden Basin.

Eocene, Continental Terminal The Continental Terminal (Kilian, 1931) succeeds the Palaeocene-Ypresian deposits in the western area of the Iullemmeden Basin. It is, in part, the lateral equivalent of the Gwandu Formation of the Sokoto region which Kogbe (1981) claims is of Eocene-Miocene age (Fig. 4). The sediments are terrestrial in origin and are supposedly the product of erosion related to regional stress patterns. The Gwandu deposits cover an area of 22,000 km 2 and the type section is 20 m thick (Kogbe, 1981). Kogbe details the sporopollen from this formation and notes the predominance of angiosperms within the flora. The palynomorphs include mangrove, palm and exotics which do not occur naturally in Africa at the present. In the central and western Iullemmeden Basin, the Continental Terminal is subdivided into 3 depositional series by Greigert (1966). These series are termed CT1, CT2 and CT3. A maximum thickness of 580 m is postulated. Lang et al. (1986) sought to clarify the concept and stratigraphy of the Continental Terminal. Based on the subdivisions established by Greigert (1966) these authors describe the Ader Doutchi Series as a sequence of marine to marginal littoral deposits, which are the equivalent of CT1 (Greigert, 1966), and dated as Middle Eocene by Boudouresque et al. (1982). In accordance with these authors Lang et al. (1986) chose to recognise the Ader Doutchi Series as a distinct lithostratigraphic unit, separated from the Continental Terminal by a major stratigraphic break. The age of the series would confirm the proposed correlation with the Gwandu Formation of Nigeria and indicates that the uppermost units in the Tilemsi Valley are of similar age. The Ader Doutchi Series are

101 also known as the S6ries Siderolithique (Greigert, 1966). The constituent lithologies are ferruginous mudstones and sandstones with oolitic horizons. These are described as subautochtonous by Lang et al. (1986), who suggest that the initial phase of deposition was in an organic-rich lacustrine environment. Boudouresque et al. (1982) support a coastal/ continental depositional environment with flora consisting of Echiperiporites icacinoides, Cicatricosis-

porites dorogensis, Chenolophonidites costatus, Striatapollis bellus and Bombacacidites sp., suggesting the proximity of savannah conditions. The stratigraphic break that separates the Ader Doutchi Series from the overlying Continental Terminal is the equivalent of the Oligocene in chronostratigraphic terms. According to Lang et al. (1986), the Continental Terminal is the equivalent of units CT2 and CT3 of Greigert (1966). In terms of lithofacies, it is divided into the S6rie Argilo-Sableuse Lignites (CT2) and the Gres Argileux du Moyen Niger (CT3). They are supposedly Mio-Pliocene in age, but are definitely pre-Quaternary. A maximum thickness of 480 m is recorded by Lang et al. (1986). The depositional break between the Ader Doutchi Series and the Continental Terminal is the result of regional uplift which is most prevalent in the eastern quadrant of the basin.

Quaternary A brief synopsis of the Quaternary of the Iullemmeden Basin is presented by Greigert and Pougnet (1967). The present day outcrop of post Continental Terminal deposits is patchy and a broad division into alluvial and dune sediments is recognised. Dubois and Lang (1981) provide a detailed comparison of Upper Quaternary sediments of the Sahel area (Niger-Tchad) with the emphasis placed on palaeoclimatology. It is interesting that the period of maximum humidity in the Niger-Tchadian Sahel area was between 9500-6500 years BP, which corresponds with a high in terms of global rainfall. It is possible that the characteristic karstification of the uppermost limestones of the Palaeocene is a result of this climatic episode.

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102 Beaumont, C. and Sweeney, J.F., 1978. Graben generation of major sedimentary basins. Tectonophysics, 50: 19-23. Bellion, Y. and Guiraud, R., 1988. D6formations d'origine compressive d'~ge intra-6oc~ne ?a Ouest de l'Adrar des Iforas (Mali). C.R. Acad. Sci. Paris, S6r. 11,307: 529-532. Bellion, Y., Benkhelil, J., Faure, H., Guiraud, R., Le Theoff, B. and Ousmane, B., 1983. Le bassin du Nig6r oriental: Observations structurales et 6volution g6odynamique. In: Bassins s6dimentaires en Afrique. La Soci6t6 G6ologique de France. Marseille, 57(2): 57 (r6sum6). Bellion, Y., Saint-Marc, P. and Damotte, R., 1989. Contribution ?a la connaissance des marins au passage Cr6tac6-Tertiare dans la vall6e du Tilemsi (Nord-Mali). J. Afr. Earth Sci., 9: 187194. Bertrand-Sarfati, J.D., Fabre, J. and Moussine Pouchkine, A., 1977. G6odynamique des aires s6dimentaires cratoniques: quelques examples sahariens. Bull. Cent. Rech. Explor. Prod. Elf Aquitaine, 1( 1): 217-231. Black, R., Jaujon, M. and Pellaton, C., 1967, Notice explicative, de la carte g6ologique de l'M'r. Bur. Rech. G6ol. Min., 6d. Direction des Mines et de la G6ologie du Niger, 57 pp. Boudouresque, L., Dubois, D., Lang, J. and Trichet, J., 1982. Contribution ?a la stratigraphie et ?a la pal6og6ographie de la bordure occidentale du bassin des Iullemmeden au Cr6tac6 Superieur et au C6nozoique (Niger et Mali, Afrique de l'Ouest). Bull. Soc. G6ol. France, 7(24) 4: 685-695. Boureau, E., 1958. Pal6obotanique Africaine. Evolution des riores disparues de i'Afrique nord-equatoriale. Bull. Sci., France, 2(!)" 1-64. Carvalho, M.S., 1982. O Genero Mawsonia na Ictiofaunula do Cretaceo do Estado da Bahia. An. Acad. Brasil. Cienc., 54: 519-539. De Lapparent, A.F., 1960. Les Dinosauriens du "Continental Intercalcaire" du Sahara central. M6m. Soc. G6ol. France, N.S., 88(A): 56 pp. Dikouma, M.S., 1990. Fluctuations du Niveau Marin au Maastrichtien et au Paleocene dans le Bassin Intracratonique des lullemmeden (Ader-Doutchi, Niger). Th6se Inter-universitaire, Dijon-Niamey, 231 pp. Dubois, D. and Lang, J., 1981. I~tude lithostratigraphique et g6omorphologique du Continental terminal et du C6nozoique inf6rieur dans le bassin des Iullemmeden (Niger). Bull. Inst. Fondamental d'Afrique Noire, 43, S6r. A, 1-2: 1-42. Fastovsky, D.E., 1987. Paleoenvironment of vertebrate-bearing strata during the Cretaceous-Paleogene Transition, Eastern Montana and Western Dakota. Palaios Soc. Econ. Palaeontol. Mineral., 2: 282-295. Faure, H., 1966. Reconnaisance g6ologique des formations s6dimentaires post-pai6ozo'/ques du Niger oriental. M6m. Bur. Rech. G6ol. Min., Paris, 4 7 : 6 2 0 pp. (Publ. Direction des Mines et G6ologie du Niger. No 1). Furon, R., 1935. Le Cr6tac6 et la Tertiaire du Sahara Soudanais. Arch. Mus6e Natl. Histoire Naturelle, Paris. S6r. 6, 12: 1-96. Genik, C.J., 1993. Petroleum geology of Cretaceous-Tertiary rift basins in Niger, Chad, and Central African Republic. Am. Assoc. Pet. Geol. Bull., 77(8): 1405-1434. Greigert, J., 1966. Description des Formations Cr6tac6 et Tertiaires du Basin des Iullemmeden. Mem. Bur. Rech. G6ol. Min., Paris, 2:234 pp. Greigert, J. and Pougnet R., 1967. Essai de description des formations g6ologiques de la Republique du Niger. Mem. Bur. Rech. G6ol. Min., Paris, 48:236 pp. Guiraud, R., Bellion, Y., Benkhelil, J. and Moreau, C. 1987. Post-Hercynian tectonics in Northern and Western Africa. Geol. J., Thematic Issue, 22: 433-466. Halstead, L.B., 1973. Hunting prehistoric reptiles in Nigeria. Nigerian Field, 38: 4-14. Halstead, L.B., 1975. Sokotosuchus ianwilsoni n.g., n.s.p. A new

R.T.J. M O O D Y teleosaur crocodile from the Upper Cretaceous of Nigeria. Niger J. Min. Geol., 11: 101-103. Handford, C.R. and Fredricks, P.E., 1980. Facies patterns and depositional history of a Permian sabkha complex. Red Cove Formation. Texas Panhandle. Bur. Econ. Geol., Austin, Texas, Geol. Circ., 80(9): 38 pp. Handford, C.R. and Bassett, R.L., 1985. Permian facies sequences and evaporite depositional styles, Texas Panhandle. In: C.R Handford, R.G. Loucks and G.R. Davies (Editors), Depositional and Diagenetic Spectra of Evaporites. Society of Economic Paleontologists and Mineralogists, Core Workshop, 3: 210-238. Jones, B., 1948. Sedimentary rocks of Sokoto Province. Bull. Geol. Surv. Nigeria, 18:75 pp. Keller, G., 1988. Extinction, survivorship and evolution of planktonic foraminifera across the Cretaceous/Teriary boundary at le Kef, Tunisia. Mar. Micropalaeontol., 13: 239-263. Kilian, C., 1931. Des principaux complexes continentaux du Sahara central. C.R. Somm. Soc. G6ol. France, pp. 109-111. Kogbe, C.A., 1974. Paleo-ecologic significance of vertebrate fossils in the Dukamaje and Dange Formations (Maastrichan and Paleocene) of northwestern Nigeria. J. Min. Geol. (Nigeria), 8( 1): 49-55. Kogbe, C.A., 1979. Geology of the South Eastern (Sokoto) Sector of the Iullemmeden Basin. Bull. Dep. Geol., A.B.U. Zaria, Nigeria, 2(1): 420 pp. Kogbe, C.A., 1980. The Trans-Saharan Seaway during the Cretaceous. In: M.J. Salem and M.T. Busrewii (Editors), The Geology of Libya, Vol. I. Academic Press, London, pp. 93-96. Kogbe, C.A., 1981. Cretaceous and Tertiary of the Iullemmeden Basin in Nigeria (West Africa). Cretaceous Res., 2: 129-186. Kogbe, C. and Lemoigne Y., 1972. Bois en structures conserv6es de Mesozo'/que C6nozo'l'que de l'Afrique Occidentale (Nigeria and Niger). Rev. Esp. Micropaleontol., Ser. 73: 505-521. Kogbe, C.A. and Me'hes, K., 1986. Micropaleontology and Biostratigraphy of coastal basins of West Africa. J. Afr. Earth Sci., 5( 1): 100 pp. Kogbe, C.A. and Wozny, E., 1980. Biostratigraphy of Maastrichtian and Paleocene Formations of northwestern Nigeria. Savana, Ahmadu Bello University, Zaria, 9(1): 35-43. Krasheninnikov, V.A. and Trofimov, D.M., 1969. Stravitei'nyy analyz bentosnykh foraminifer datscko paleotsenovykh othozheniy Mali, ablastitetsia i sever-zapoadnoy Europy. Akad. Nauk SSSR Vopr. Mikropaleontol., 12: 108-144. Lang, J., Kogbe, C.A., Aiidou, S., Alzouma, K., Dubois, D., Houesson, A. and Trichet, J., 1986. Le sid6rolithique du Tertiaire ouest-africain et le concept de Continental terminal. Bull. Soc. G6ol. France, 11(4): 605-622. Maisey, J.G., 1987. Coelacanths from the Lower Cretaceous of Brazil. Am. Mus. Novitates, New York, 2866:1-30 Matheis, G., 1976. Short review of the geology of the Chad Basin in Nigeria. In: C.A. Kogbe (Editor), Geology of Nigeria. Elizabethan Publishing, Surulere, pp. 289-294. Merabet,O., Klotchko, V., Timonine, L. and lvannikov, A., 1971. Pr6cision sur l'age de la transgression du C6nomano-Turonien inferieur au Sahara central. Publ. Serv. G6ol. Algerie, 41: 217223. Miall, A.D., 1977. A review of the braided-river depositional environment. Earth-Sci. Rev., 13: 1-62. Mongin, D., 1954. Sur divers lamellibranches d'eau douce r6colt6s dans le "Continental Intercalcaire" du Sahara. C.R. Acad. Sci. France, 239(13): 771-773. Moody, R.T.J. and Sutcliffe, P.J.C., 1990. Cretaceous-Tertiary crossroads of migration in the Sahel. Geol. Today, 6(1): 1923. Moody, R.T.J. and Sutcliffe, P.J.C., 1991. The Cretaceous deposits of the Iullemmeden Basin of Niger, central West Africa. Cretaceous Res., 12:137-157.

T H E I U L L E M M E D E N BASIN Moody, R.T.J. and Sutcliffe, P.J.C., 1995. The sedimentology and palaeontology of the Upper Cretaceous-Tertiary deposits of central West Africa. In: W.A.S. Sargeant (Editor), Vertebrate Fossils and the Evolution of Scientific Concepts. Gordon and Breach, Amsterdam, pp. 539-554. Parker, D.H., 1965. Sheets 2, 3, 7, 8. Geological Map Series of Nigeria, 1:100,000. Geological Survey, Nigeria. Petters, S.W., 1977. Ancient seaway across the Sahara. Nigerian Field, 42: 22-30. Radier, H., 1953. Contribution h l'6tude stratigraphique et structurale du d6troit soudanais. Bull. Soc. G6ol France, (6) 3: 677-695. Radier, H., 1957. Contribution ~ l'6tude g6ologique du Soudan oriental (A.O.F.). Bull. Dir. f6d. Mines G6ol. Afr. occid, fr. Dakar, 26: 1-556. Reyment, R.A., 1987a. Transgressional maxima of the Late Cretaceous in the Western Tethys. In: Proceedings of the International Symposium on Shallow Tethys (Wagga Wagga). A.A. Balkema, Rotterdam, 2: 303-308. Reyment, R.A., 1987b. Spanish and North African region of Tethys in the Late Cretaceous. Proceedings of the International Symposium on Shallow Tethys (Wagga Wagga). A.A. Balkema, Rotterdam, 2:309-317. Taquet, P., 1976. G6ologie et PalEontologie du gisements de

103 Gadouafa (Aptian) du Niger. Cah. Pal6ontol., Paris, 191 pp. Trofimov, D.M., Aristova, K.Ye. and Petrosyants, M.A., 1969. Verkhnemolovyje i paleogenovyje otlozheniya grabena Gao (Yuzhnaya Sakhara). lzv. Vyssh. Uchebn. Zavedenij., Geol. Razvedka, 5: 62-67. White, E.I., 1934. Fossil fishes of Sokoto Province. Geol. Surv. Nigeria Bull., 14: 1-78. Windley, B.F., 1978. The Evolving Continents. Wiley and Sons, Chichester, 385 pp. Wood, J.M., Thomas, R.G. and Visser, J., 1988. Fluvial processes and vertebrate taphonomy: The Upper Cretaceous Judith River Formation, South-central Dinosaur Provincial Park, Alberta, Canada. Palaeogeogr., Palaeoclimatol., Palaeocol., 66: 127143. Wright, J.B., Hastings, D.A., Jones, W.B. and Williams, H.R., 1985. Geology and Mineral Resources of West Africa. Allen and Unwin, London, 187 pp. Zaborski, P.M.P., 1982. Campanian and Maastrichtian sphenodiscid ammonites from southern Nigeria. Bull. Br. Mus. (Natural History), Geol., 36(4): 303-332. Zaborski, P.M.P., 1984. Upper Cretaceous ammonites from the Calabar region, south-east Nigeria. Bull. Br. Mus. (Natural History), Geol., 39: 1-72.

Chapter 6

Rift Basins of the Sudan

RAMSIS B. SALAMA

The Sudanese rift structures form intracontinental basins, bordered on all sides by anorogenic terrain. These basins are seen as the result of a multistructural system of rifts which appear to have been activated several times since the Palaeozoic. Improved methods of dating and extensive geophysical surveys have led to a better understanding of the mobile belts and the Basement rocks in the general environs of the basins, and of the elevated blocks and highs surrounding the sedimentary basins. Detailed mapping and extensive drilling in the sedimentary basins led to a better understanding of the Phanerozoic. The rapid rate of uplift and subsidence have assisted in the rapid accumulation and filling of the basins with unconsolidated sediments, ranging from a few hundred metres to some thousand metres. The rift structures contain sediments of several age groups, origin and mode of deposition. Some of them are as old as the Palaeozoic passing through to Mesozoic, Tertiary and Quaternary (Bahr El Arab rift and Blue Nile rift). The thickness of the Tertiary sediments exceeds 15 km in Bahr El Arab rift. Palaeozoic sediments were found to occupy N E - S W grabens. Jurassic? and Cretaceous sediments were deposited in N W - S E troughs. Tertiary sediments filled grabens that range in thicknesses up to 15 km. Highly saline ground-water bodies which coincide with hydrogeologically closed basins occupy the flowing end of each of the rift basins. These saline bodies have been interpreted as buried saline lakes, sabkhas or playas. Alternating wet and dry periods eventually filled up the basins with intercalating layers of fresh and saline layers. The filling up of the basins led to the interconnection of the hydrological regime and the development of the integrated River Nile. Rift structures in Sudan form the major ground-water basins, outside these basins ground water is found only in small quantities and is of poor quality. The rift basins also act as reservoirs for hydrocarbon accumulation; all the known oil fields of inland Sudan are located in these basins.

INTRODUCTION

The concepts of plate tectonics have revolutionised the interpretation of the geological history in the Sudan. Previously, the geological map of Sudan was dominated by four main units, the PrecambrianCambrian Basement complex rocks which covers 50% of the map area, Palaeozoic continental deposits scattered on the northwestern corner, Mesozoic Nubian sandstone formation outcropping in the northern part of Sudan and covering about 30% of the total area, and Tertiary Umm Ruwaba deposits covering the central and southern parts. The sedimentary deposits were thought to have filled shallow synclinal basins. Due to the small number of geologist employed by the Geological survey, the vast area of the country and the belief that the mineral resources of the country are not worth exploring, the pace of geological mapping and exploration was moving at a very slow rate. It was through the search for ground water that most of the recent geological information was compiled.

A systematic pattern of structurally controlled emerged from the analysis of the massive amount of information collated from the detailed geological, hydrogeological and geophysical work carried out during the period from 1966 to 1982. By 1975 most of the basins were known, the extensive geophysical work revealed the presence of more than 5 kilometres of sediments in the northern part of Bahr El Arab rift. This was the first indication that these basins might be rift structures and triggered the hunt for oil exploration. In this chapter I will concentrate on the Sudanese Rift basins. First, I will discuss the geological structures during the Precambrian-Cambrian, Palaeozoic, Mesozoic and Tertiary. A detailed description of the tectonic activity in Sudan and adjoining African countries will be given. Next, I will discuss the sedimentological processes which led to the filling up of the rift basins during the last stages of rifting and the Sag period. Also discussed, is the chemical pattern of the buried saline ground-water zones and the possible scenarios for the reconstruction of the buried

African Basins. Sedimentary Basins of the World, 3 edited by R.C. Selley (Series Editor: K.J. Hsti), pp. 105-149. 9 1997 Elsevier Science B.V., Amsterdam. All rights reserved.

106 saline lakes and their relation to the evolution of the River Nile and the ground-water and petroleum resources of these basins. I have relied on field data collected during my work in Sudan from 19661981, and the ideas I put forward during my study period in the University of New South Wales, where I was able to collate an extensive amount of information and put together the theories which were partly published in several papers and in my unpublished Ph.D thesis. The review covers selected work related to the recent concepts which are relevant to the development of the Sudanese Rift System. The outlines given are not meant to be comprehensive or detailed.

GEOLOGICAL SEQUENCE Precambrian-Cambrian The concepts of the Sudanese Basement Complex evolution has changed radically since Whiteman (1971). Vail (1976) has suggested that the Kibaran belt swings clockwise from Uganda through the west of Sudan as the Zalingei folded belt (Fig. 1). In addition, Vail (1982) has shown that the NE trending grain of the Basement is intersected by N-S and N W SE trending fractures parallel to the Red Sea coast, with meridional fractures probably related to the Red Sea Rift zone. Almond (1982) summarised the eastern Africa mobile belts and correlated them with the Sudanese Basement rocks, Almond (1982) postulated the extension of the Archaean craton and the western margin of the Mozambique Belt north of the Ugandan border as far north to the Bayuda Desert. Krrner et al. (1987) suggested that the East African Mozambique belt extends from Tanzania, Kenya and northeast Uganda through the southern Sudan and western Ethiopia into the region of the River Nile. They also suggested that the gneisses of the River Nile at Sabloka represent the infrastructure of the ancient African continental margin onto which the juvenile arc assemblage of the Arabian-Nubian shield was accreted during intense horizontal shortening and crustal interstacking of a major collision event. This is in agreement with the model proposed by Fleck et al. (1980) in which the newly created crust of proto-Arabia eventually collided with the proto-African continent to the west. On both sides of the suture, collision resulted in thickening of the crust, transcurrent faulting on a major scale, and widespread magmatism expressed in numerous late to post-tectonic igneous complexes with ages between 660 my and 600 my. Fleck and his coworkers drew a parallel between this sequence and the collision of India with Asia. They also correlated the collision in Arabia with the "Pan African Event" recognised in over two-thirds of the African conti-

R.B. SALAMA nent by a concentration of radiometric dates in the range of 650-450 my. E1 Ageed and E1 Rabaa (1981) in his correlation of the Pan African lithostructural belts in the northeast Nuba mountains implied a two fold subdivision of the basement into a lower cratonized gneissic group, which may represent an older ensialic basement, possibly of earlier Precambrian age, overlain by an upper mobile geosynclinal metasedimentary and metavolcanic assemblage with late- and postorogenic igneous phases and epeirogenic molasse facies sediments. Regional correlation with African orogenic belts suggests that the gneissic group may be part of an older, deeply eroded basement succession (probably Mozambiquian), whereas the geosynclinal sequence may be part of an aulacogenic PanAfrican "basement". Ahmed (1982) in his study of the Precambrian lineaments in NE Sudan, concluded that, besides the curvilinear features, a strong pattern of Precambrian structures trending mainly in N-S, NNW-ENE, and less distinct, poorly exposed E - W and NW-SE systems were developed. As a result of this, the region was dissected into a mosaic of large basement blocks surrounded by tectonic zones of faulting and strong shearing. Faults along the lineament zones seem to have been rejuvenated several times during the subsequent tectonic events which were active since post-Cretaceous until recent (Gass, 1955; Kabesh, 1962; Qureshi and Sadig, 1967; Ruxton, 1965; Vail, 1978; Whiteman, 1971). In Sudan, large transcurrent faults include the north-west trending Aswa fault (Almond, 1982) and its possible extension northward into Sudan (Salama, 1985a). Palaeozoic In northwestern Sudan and adjacent Chad, Libya and Egypt, Palaeozoic deposits of marine and continental origin are recorded (Whiteman, 1971; Vail, 1978) but in Sudan little detailed work has yet been completed (Klitzsche, 1983). In central Sudan the Nawa formation, which is reported from boreholes only, was assigned a probable Palaeozoic age (Andrew, 1948). E1 Rabaa (1976) recorded unmetamorphosed sedimentary sequence of pre-Mesozoic age in two localities one along the River Nile in northern Sudan, and the other at Khor Abu Habil in western Sudan. E1 Rabaa (1976) considered the deposition of Abu Habil sediments to be structurally controlled along major troughs produced by intensive faulting. The rocks are considered to represent late or postorogenic molasse facies and probably constitute a part of a major tectonic pattern at the late stages of geosynclinal activity in the late Precambrian history of north Africa. Salama (1985a) suggested that the Palaeozoic rocks may be more widespread than once thought,

RIFT BASINS OF T H E SUDAN

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and proposed four areas of Palaeozoic(?) deposits (Table 1 and Fig. 2): (1) Wadi E1 Kuu Greywacke: Consisting of arenaceous layers with Basement Complex fragments underlain by limestone and sandstone.

(2) Baggara Greywacke: formed of limestone and arkosic layers with Basement Complex fragments lithologically similar to Abu Habil sediments. (3) Abu Habil Greywacke: Sediments beneath the Tertiary deposits of Umm Ruwaba basin form the

108

R.B. SALAMA

Fig. 2. Geological map of Sudan, showing major structural trends, Palaeozoic grabens, Mesozoic depressions and volcanic fields. Paleozoic(?): PI = Abu Habil greywacke; P2 = Wadi El Kuu greywacke; P3 = Baggara greywacke; P4 = Blue Nile mudstone. Mesozoic: M I = River Nile sandstone; M2 -- western Sudan mudstone; M3 -- Gedaref sandstone. Cainozoic volcanics; V I -- J. Marra; V2 = Tagabo; V3 = Meidob; V4 = El Melik; V5 = Bauyda; V = Red Sea Hills.

RIFT BASINS OF THE SUDAN

109

Table 1 (?)Palaeozoic sediments in Sudan Proposed unit

Area

Type section

Abu Habil greywackes

Locality type area in Abu Habil and extends below Umm Ruwaba and Abu Rukba (12 37 N, 30 38 E)

5. volcanic tuff 4. pyroclasts 3. conglomerates 2. sandstone 1. limestone

Wadi El Kuu greywackes

Sag E1 Naam area in Darfur Province

5. arenaceous deposits 4. limestone 3. volcanics 2. mudstone 1. arkosic layer (well records in Sag E1 Naam)

Baggara greywackes

Between Abu Materig and Ramies (south Darfur)

3. mudstone 2. arkosic deposits 1. limestone (well records between Abu Materig and Ramies)

Blue Nile mudstones

South of Khartoum Northern Gezira

4. brown mudstone 3. volcanics (veneer) 2. clay 1. limestone and volcanics (Umm Udam borehole)

After Salama (1985a).

northern extension of the Abu Habil type formation which is formed of siltstone, limestone, sandstone, conglomerate and volcanics. The presence of stromatolite in the limestone indicate deposition in shallow, near shore, fresh or marine environment. As no evidence is available that the sea ever transgressed as far as this central part of Sudan, it is assumed that the environment of deposition of these Palaeozoic deposits was shallow fresh water. (4) Blue Nile Mudstone: Formed mainly of micaceous sediments, blackish clays, thin layers of limestone and volcanic tuffs. Mesozoic Following the Pan-African Event, the Nubian Shield remained remarkably stable for more than four hundred million years until towards the end of the Mesozoic era, there began the long succession of gentle warping and sporadic volcanic outbreaks which preceded and followed the opening of the Red Sea oceanic rift (Salama, 1985a). The formation of Sudan swells, which was followed by the volcanicity and rifting in the central part of Sudan, are related to the same forces causing the opening of the Red Sea. During the Mesozoic, the northern and eastern parts of the continent acted as depositional basins for sediments from the higher cratons and deposited in fans, deltas and lakes. In Sudan, the deposits crop out north of a line extending NE from Jebel Marra in the west, passing through northern Darfur, northern Kordofan, Khartoum and the Northern Province. The sediments form outcrops lying unconformably over the more elevated basement rocks. South of this line the sediments are reported from well logs, in faulted

blocks which strike N W - S E (Salama, 1985a; Schull, 1988). These sediments are included under the general name, Nubian Sandstone Formation. Vail (1974) mentioned that the formation is poorly known in Sudan. There is also international controversy over its name, age, correlation, origin, stratigraphy and distribution. Salama (1985a) in agreement with Weissbord (1970) that the term Nubian was ill-defined from the beginning, and in support of Pomeyrol's plea (1968) to abandon the term "Nubian Sandstone" (endorsed by Klitzsche, 1983), proposed the classification shown in Table 2 and Fig. 2 for the continental Mesozoic sediments which crop out in the northern part of Sudan. The River Nile Sandstone outcropping along the Nile from Khartoum to Shendi to Dongola. The Central Sudan Mudstone between E1 Nahud, J. Hilla, Umm Kedadda, Abyad, Meidob and Tagabo. Gedaref Sandstone in Gedaref area. The classification is based on palaeotransport directions (which indicate two major depositional depressions extending in a N W - S E direction (Kheiralla, 1966; McKee, 1962; Vail, 1974)), the areal extent of the deposits and on lithology. In the subsurface the thick sequence of Cretaceous sediments fill deep troughs more than 6000 m in thickness (Schull, 1988). The CretaceousPaleocene sediments reflect two cycles of deposition, each represented by a coarsening upward sequence (Fig. 3). The first cycle is represented by the Sharaf, Abu Gabra and Bentiu formations. The second cycle is represented by the Cretaceous Darfur Group and the Paleocene Amal Formation. The sediments have been classified to the following (Schull, 1988):

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Proteacidites Sigalii Traorites 37

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RIFT BASINS OF THE SUDAN

111

Table 2 Mesozoic continental sediments in Central Sudan Proposed unit

Area

Type section

River Nile sandstone

Along the Nile from Khartoum-Shendi-Dongola

5. mudstone 4. quartzose sandstone 3. Merkhiat sandstone 2. intraformational conglomerates 1. pebble conglomerate (Kheiralla, 1966)

Central Sudan mudstones

Between El Nahud, J. Hilla, Umm Kedadda Abyad, Meidob and Tagabo

4. intercalations of sandstone and mudstone 3. mudstone 2. silty or clayey sandstone 1. sandstone and conglomerates (Karkanis, 1965)

Gedaref sandstone (Whiteman, 1971; Gedaref formation)

To include the sandstone below the surface in Gedaref area

3. mudstone 2. sandstone 1. conglomerate (Whiteman, 1971)

After Salama (1985a). (a) Sharaf and Abu Gabra formations. The early graben-fill clastics are first cycle sediments derived from the gneissic basement complex. During the early phases of rifting, Neocomian and Barremian claystones, siltstones, and fine grained sandstones of the Sharaf Formation were deposited in fluvialfloodplain and lacustrine environments. Toward the basin edges and in the areas of major sediment influx these sediments graded to coarse alluvial clastics. The Aptian-early Albian Abu Gabra Formations represents the period of greatest lacustrine development. Several thousand meters of organic rich lacustrine claystones and shales were deposited with interbedded fine grained sands and silts. The nature of the deposit was probably the result of humid climate and the lack of external drainage, indicating that the basins were tectonically silled. The Abu Gabra Formation is estimated to be about 1800 m thick. In the northwestern Muglad block organic sands were deposited in a lacustrine deltaic environment. These form the primary source rock of oil in the interior basins. (b) Bentiu Formation. During the late AlbianCenomanian, the predominantly sand sequence Bentiu Formation was deposited. The alluvial and fluvial-flood plain environments expanded, probably due to change from internal to external drainage. These thick sandstone sediments were deposited in braided and meandering streams. This unit which is up to 1500 m thick, shows good reservoir quality. Sandstone of the Bentiu formation are the primary reservoirs of Babanusa Block. (c) Darfur Group. The Turonian late Senonian period was characterised by a cycle of fine to coarse grained deposition. The lower portion of the group is characterised by the predominance of claystone, shale, and siltstone. These initial deposits followed the second rifting phase. The excellent regional

correlation of this unit verifies the strong tectonic influence on sedimentation. Flood plain and lacustrine deposits were widespread. Interbedded with the floodplain and lacustrine claystones, shales and siltstones are several fluvial/deltaic channel sands generally 3-20 m thick. These reservoirs are significant reservoirs in the Unity area. The Cretaceous ended with the deposition of increasingly coarser grained sediments, reflected in the higher sand percentage of the Gazal and Baraka formations. These units were deposited in sand-rich fluvial and alluvial fan environments which prograded from the basin margins. The Gazal formation is also an important oil reservoir unit in the Unity field. The Darfur group is up to 1800 m thick.

RIFT STRUCTURES AND RIFTING PHASES

In a detailed study of the rift structures in Sudan, Salama (1985a) deduced that Cainozoic up-doming, volcanicity and tensional stress associated with the movement of the African plate created rift structures which were formed by successive block faulting along palaeotrends. These were followed by subsidence and linear uplift to create the biggest rift structures in Africa; the Sudanese Rift System. This extends from the western boundaries of Sudan to the eastern borders with Ethiopia and includes: (a) Bahr E1 Arab rift, (b) White Nile rift, (c) Blue Nile rift, (d) River Atbara rift, and (e) Wadi E1 Kuu rift (Fig. 4). Domal uplift, midplate volcanism, fracturing and faulting are the major forces leading to the formation of rift structures (Gass et al., 1978). The Sudanese volcanic belts J. Marra, J. Meidob, J. Tagabo, Baiuyda and the Red Sea volcanics are all located on crystalline basement which had been uplifted by the Darfur dome (Gass et al., 1978;

112

R.B. SALAMA I

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Francis et al., 1973; Vail 1972; Salama, 1985a, b). The Darfur dome is mainly parallel to the Zalingei fold belt of Vail (1978), which indicates that

the friction at the contact zones between the mobile belts can generate enough heat to raise the temperature. The high temperature causes the thin-

RIFT BASINS OF THE SUDAN ning of the lithosphere accompanied by volcanism. This causes the development of low density regions which leads to the initiation of the uplift. The contact points between the mobile belts are zones of weakness along which magma may flow up and increase the rate of upwelling. This upwelling is enough to cause tensional forces which caused fracturing and faulting perpendicular to the direction of stress, creating the extensive northwest faulting system. At the same time the spreading of the Red Sea produced another tensional field which caused the N - N E fracturing and faulting patterns. The intersecting normal faulting systems led to the formation of blocks, which were continuously subsiding. The subsidence was accompanied by uplift on the flanking sides. This in turn led to the creation of another compressional field, which in turn created its own fracturing system (Salama 1985a, b). The extension of the Mozambique belts northward into Sudan (Vail, 1976; Almond, 1982; Krtiner et al., 1987) as the Butana belt which is mainly of the green-schist type indicating that this area was exposed to shallow levels of cratonic erosion. While the Darfur dome, with its gneissic-granitic exposures, i.e in Malha crater, and the shallow depth to the granitic zone (Salama, 1985a) indicates exposures of deeper levels of the basement, which have been caused by the domal uplift along the Zalingei fold belt. The separation of Arabia from the African plate along the Red Sea with its rotation poles in North Africa (Girdler and Daracot, 1972) created northeasterly tensional stress fields and north-westerly structures in Sudan. In contrast, in the southern part of the continent rotation about poles in southeast Africa caused north-easterly structural patterns. The basic theory for the formation of the Sudanese Rift System, is that the anticlockwise rotation of Africa during the last 70 million years (Girdler, 1968) led to the opening of the southern Sudanese Rift System. The clockwise rotation which started in the last ten million years created the northern Sudanese Rift System. The two systems are separated by the Darfur dome, which is a major uplift which started at the end of the Mesozoic and possibly continued during the Tertiary. This is shown by the fact that the Darfur dome zone is not covered with any Tertiary sediments, which fill up very deep rift grabens south and north of this dome (Salama, 1985a). The common feature of the east Africa Rift structure is that rifting occurs along the main uplifted zone. This is not the case in the Sudanese Rift System, where the combined stress forces led to the formation of the rift perpendicular to the direction of those two forces away from the elevated block. The nondevelopment of rift structures along the elevated zones can be due to either that the elevated zone has not developed

113 enough to lead to rifting in its top part or that the elevated domes were highly granitized (Malha Crater). Based on the length, depth and width of the Sudanese Rift System, it is clear that the forces which caused the rift are more prominent in the western side and seems to fade away towards the east. This indicates that the stress in the Sudanese Rift System increases from east to west. Browne et al. (1985) in their description and comparison of the White Nile rift, concluded that the Sudanese rift developed earlier than the main rifting in East Africa and had a trend that was nearly perpendicular to the eastern and western rifts. The mechanism of rifting also differed, with extension and crustal subsidence in the Sudan, and uplift and little extension in East Africa. They also concluded that a structural lineament, which they assumed to be the extension of the Central African shear zone, appeared to have acted as a structural barrier to the development of deep Cretaceous-Tertiary sedimentary basins in the northern Sudan. Salama (1985b) disagreed with this last conclusion on the ground that Bahr E1 Arab rift is not terminated by this lineament and the presence of Wadi E1 Kuu rift northward was considered an extension of the Aswa shear zone of Hepworth and Macdonald (1966). The same thing applies to the Blue Nile rift which seems to extend structurally in a northwest direction beyond the proposed line. The lineament which seems to define the northern limit of the southern rift grabens is the Darfur dome. Browne and Fairhead (1983) suggested three periods of rifting which have occurred in response to crustal extension, this provided the isostatic mechanisms of subsidence which was accomplished by normal faulting parallel and subparallel to the basinal axes and margins. Girdler (1983) proposed three stages for the development of the eastern Africa Rift structures; 44 to 38 million year, 16 to 11 million year, and 5 to 0 million year. Baker (1965), from his work in Kenya, showed that the dates for the development of the rift are: (a) Late Pleistocene which is still going on (0-1 million years), (b) 4 - 2 million years, and (c) 12-18 million years. Salama (1985b) has shown that the filling of the Sudanese rift basins would require periods ranging from 11 million years in the east (Atbara rift) to 33 million years in the south (Bahr El Arab rift), with the older rift basins in the west and the younger in the east. This makes these comparable with the first and second stages of development of the eastern African rift. Schull (1988) suggested from the oil drilling penetrations that the initial strongest rifting phase have begun in the Jurassic(?)-Early Cretaceous (130-160 Ma) and lasted until near the end of the Albian. The termination of the initial rifting is stereographically

114

R.B. SALAMA

marked by basin wide deposition of Bentiu Formation thick sandstone deposits. The second rifting phase occurred during the Turonian-late Senonian characterised by the wide spread deposition of lacustrine and floodplain claystones and siltstones. This rifting phase was accompanied by minor volcanism which occurred 82 Ma (4-8 my) dated from a dolerite sill in the northwest Muglad basin. The end of this phase is marked by the deposition of an increasingly sand-rich sequence that ended with the with the deposition of thick Palaeocene sandstone Amal Formation. The final rifting phase began in the late Eocene-Oligocene and is characterised by thick lacustrine and floodplain claystone and siltstones. Late Eocene basalt flows are recorded in Melut block near Ethiopia. The rifting was followed by an intracratonic sag phase during the middle Miocene of very gentle subsidence accompanied by little or no faulting. Limited outcrops of volcanic rock in the area southeast of Muglad dated at 5.0 Ma and 2.7 Ma indicate that minor volcanism occurred locally. This is associated with the extensive volcanism of Jebel Marra, Tagabo and Meidob along the Zalingei Folded Belt (Salama, 1985a, b).

THE SUDANESERIFT SYSTEM In the following discussion the terms graben, trough and basin will refer to first, second and third order divisions within the mentioned rift. Bahr El Arab rift

Bahr E1 Arab rift comprises two major structures, the Baggara graben and the Sudd graben (Fig. 4) (Salama, 1985a, b). Baggara graben covers the area between the Nuba mountains in the east and the Central African Republic in the west. In the north it is defined by the faulted Mesozoic deposits south of the Darfur dome. North of this line the Mesozoic deposits crop out in the form of a chain of low lying, fiat topped hills, covered by lateritic deposits, and extends for a distance of more than 200 km. South of the faulted zone the Mesozoic sediments are found below the surface and are known from borehole records and remnants of sandstone and laterites cropping out along Bahr E1 Arab (Salama, 1985a). The Sudd graben is bounded in the north by the Nuba mountains and the Akoke ridge in the northeast (Geophysics and Strojoexport, 1977), in

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Fig. 5. Section across Bahr El Arab rift, showing the Abu Gabra structural trend in Muglad basin (A) and Unity structural trend in southern Muglad Basin (B). (Modifiedfrom T.J. Schull, 1988. Reprinted by permission of the American Association of Petroleum Geologists.)

RIFT BASINS OF THE SUDAN

115

the west and southwest by the uplifted basement ridge which marks the Nile-Congo divide, and in the east by the basic volcanics extensions of the East African rift. The basins within the Babanusa trough are defined by extensive faulting system extending in NW, NE and EW (Fig. 5a and b). The grabens and horsts indicate a step-like subsidence of separate blocks. The intensity of the faulting and

the subsidence increase southward, where it attains a depth of more than five kilometres at the Unity oil field, and an estimated eleven kilometres south of Bantiu oil field (Fig. 6a and b) (Anon, 1981a, b, 1982; Schull, 1988). The Kurkur hills which is a Basement uplift in the northern margin of the Baggara graben defines the northern margin of Abu Gabra trough. Its western boundary is represented by a major NW fault system extending southward along the SE sharp bend of Bahr E1 Arab. It extends southeast defining the western margin of the Sudd graben and connecting with the northerly extension of the Aswa line. There are some indications that this fault system is active in the Sewar trough. Near Sewar ("Lat. 11 10 N, Longt. 24 35 E"), there is a zone of hot ground water, extending along the fault line, as confirmed by boreholes lying along this line (Table 7) (Salama, 1985a). White Nile rift

Fig. 6. (A) Depressions and elevations in Babanusa graben, showing the extensive faulting system. (Modified from Geophysics and Strojoexport, 1976.)

The White Nile rift (Fig. 4) is formed by the junction of two major grabens; the Umm Ruwaba graben extending in a NW direction and the White Nile graben extending in a N to NW direction. It is bounded in the north by E - W faults and N E SW fault systems which bring it into contact with Mesozoic sandstones with a southerly down throw of 400 m forming Bara trough. At Bara further south, another fault system is known from geophysics with a northerly downthrow of 600 m (Ali, 1978). An uplifted basement block extending in a N W - S E direction divides the graben into two troughs; the Bara trough in the east and Umm-Ruwaba Renk trough in the Southeast. The Umm Ruwaba trough is characterised by a series of horsts and grabens formed by two sets of fault systems E - W and N W - S E (Fig. 7) (R.E.G.W.A., 1979). The basement rocks and the overlying Mesozoic deposits are block faulted upwards at Rabak whereas the Mesozoic sediments are downthrown more than 100 m at Kosti. A series of grabens and horsts formed by N W SE fault systems, forming the Renk, Wankir, Wadi M

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116

R.B. SALAMA

Fig. 7. Geophysical section in Abu Habil trough showing the step faulting pattern. (Modified from R.E.G.W.A., 1979.)

Adar, Gereid and Akoraweng troughs of the White Nile graben which extends along the White Nile from Renk south to near Kodok (Geophysics and Strojoexport, 1977; Anon, 1981a, b, 1982) Blue Nile rift

The Blue Nile rift (Figs. 4 and 8) extends from south of the Sabaloka gorge "The Sixth Cataract" southeast to the Sudan borders with Ethiopia, following the Blue Nile, River Rahad and River Dinder. It is formed of three main grabens, from north to south, Khartoum, Wad Medani and Singa. Its southern limit is the elevated Basement block extending in a N W - S E direction with various hills cropping out along this boundary line, i.e.J. Dali, J. Mazmum, J. Moya, J. Doud and J. Biuyt. The northern margin is another elevated block of basement rocks extending in a N W - S E direction east of River Rahad, i.e. J. Fau. A series of fault systems striking N W - S E parallel to the Blue Nile, River Er Rahad and River Dinder having westerly downthrows with southerly increase in depth. River Atbara rift

The River Atbara rift (Fig. 9) extends from Atbara town at the junction of the fiver Atbara with the River Nile and extends southward to the Ethiopian border, coveting El Gash and river Atbara tributaries southward. The River Atbara seems to follow a fault line extending from Atbara to Qoz Regeb in the south forming the Atbara trough. Wadi E1Makabrab, which is a seasonal stream parallel to the River Atbara is also following a system of fault lines extending N W - S E

to Wadi E1 Hawad with easterly downthrows of 100 and 200 m, respectively. Along the River Atbara the fault is further downthrown another 100 m, with thick unconsolidated Tertiary sediments filling the trough. Aroma graben which occupies the recent Gash delta is triangular in shape and extends from Kassala town along the Gash river NW to near Hadaliya. J. Kassala is the northern uplifted block of this trough. In Aroma the down faulted Tertiary and Mesozoic deposits are downthrown more than 300 m to the north (Saeed, 1974). Wadi El Kuu rift

Wadi El Kuu rift (Fig. 10) is formed by two grabens, Sag E1 Naam NW trending graben and the Wadi E1 Kej E - W graben. Shagra trough which forms the northern part of Sag E1 Naam, is a fault bounded trough (Mohamed, 1975). The Mesozoic deposits cropping out in the southern side are downthrown more than 300 m in the trough. Lava flows ranging in thickness from 20-100 m seems to have invaded the fissured sandstone. Sag E1 Naam trough which is a fault bounded trough (Medani and Vail, 1974) with sediments of more than 2000 m. The eastern margin is the Summeiat uplifted block, which separates the trough from the Central Sudan Mesozoic sediments. It is connected to Abu Gabra trough of Bahr E1 Arab rift through Wadi El Kuu faulted zone. Wadi E1 Kej graben is formed of a series of horsts and grabens, i.e. Shangil Tobaya and Wadi E1 Kej trough: A possible northerly extension of Wadi E1 Kuu rift, is another fault bounded trough, separated from Shagra trough by the Mellit volcanic field;

RIFT BASINS OF THE SUDAN

117

Fig. 8. Schematic diagram showing the rift system in Bahr E1 Arab rift, White Nile rift and Blue Nile rift.

Wadi E1 Sherak and Sanya Hayei. This seems to be the nearest fault system to the Darfur dome.

SEDIMENTARY SEQUENCE OF THE SAG PERIOD

The previously described rift structures in Sudan formed the main depositional basins in which sed-

iments from the high uplifted zones accumulated. The Sudanese rift basins are mainly intracontinental basins, bordered on all sides by anorogenic terrain. The rapid rate of uplift and subsidence have assisted in the rapid accumulation and filling of the basins with unconsolidated sediments, ranging from a few hundred metres to some thousand metres. The rift structures contain sediments of several age groups,

118

R.B. SALAMA H

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origin and mode of deposition. Some of them are as old as the Palaeozoic passing through to Mesozoic, Tertiary and Quaternary (Bahr E1 Arab rift and Blue Nile rift). The thickness of the Tertiary sediments exceeds 10 km (Anon, 1981, 1982; Schull, 1988) in Bahr E1 Arab rift. In the following discussion, each rift basin will be discussed separately and the drainage system, types and possible modes of sedimentation of the top 500 to 1000 m which was deposited during the Sag Period will be discussed. Deltas and river fans

The rift basins in Sudan were filled with fluvial sediments mostly in the form of fans (Fig. 11). Several workers have described the fans in these basins. Williams and Adamson (1982) described the Gezira (the area between the Blue and White Nile) as a complex low angle alluvial fan of a type common in other semi arid areas of the world. They recognised three upward fining alluvial cycles. Williams and Williams (1980) in their synthesis of Nile evolution showed that extensive alluvial fans

were built up in southern and central Sudan at the outlet of major rivers issuing from the highlands of Ethiopia. Adamson and Williams (1980), showed that the surface of Bahr E1 Arab rift is covered by active and abandoned alluvial fans, by swamp deposits, by active distributary streams and by prior stream channels. They described a series of low angle alluvial fans and trough sediments in the boundary of the Nuba mountains and the Babanusa trough. Adamson et al. (1982) showed the presence of palaeochannels of the Blue Nile, Rahad and Dinder. These can be seen in Fig. 12. The palaeochannels of the Blue Nile show very clearly two branches. One is parallel to the recent Blue Nile and terminates in a fan form in the Soba area, while the other moves west towards the recent White Nile and some of its branches move southwest, towards Ed Dueim. Marsail Salama (pers. commun.), in a detailed interpretation of ERTS imagery for Khor Abu Habil, showed the presence of palaeochannels extending from Umm Ruwaba eastwards with main tributaries towards the northeast to Kosti and southeast to Keri Kera. She also showed that Abu Habil fan deposits

RIFT BASINS OF T H E SUDAN

119

Fig. 11. Central Sudan Alluvial fans, deltas and swamps. Map showing the alluvial fans of Gash and Atabara in the Atabara rift system; The Blue Nile fans in the Blue Nile rift system; The Abu Habil fan and Mashar marshes in the White Nile rift; the Sudd in Bahr El Arab rift. Blue Nile fan from S.I.K.R. (1987), Adamson (1982); Abu Habil fan from Gunn (1982), ERTS imagery (1972). All other fans from topographical maps and air photos. (After Salama, 1987.)

120

R.B. SALAMA I

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are mainly clays and very fine silts. Similar deposits are known in most of the internal drainage systems of today's major wadies, such as the River Gash, Wadi E1 Kuu, Wadi Nyala, Wadi Bulbul and Wadi Shellango (Salama, 1985b, 1987). The variables that affect sediment yield of a catchment area are relief, runoff, precipitation distribution and vegetative cover. Interchanging the individual extremes of these four variables will affect the sediment yield from one extreme to the other. Postulating any of those variables into the geological past can produce different end products. Taking this into consideration, a reconstruction of the palaeochannels of the Sudanese Rift System has been made in order to assist in determining the modes of deposition of the large amount of unconsolidated sediments within the rift structures that appears to have taken place

during the Tertiary period. No attempt has been made to reconstruct the palaeoflow characteristics of those channels. The existing drainage system has been used as a replica for the palaeochannel system to divide the basin into catchment areas and for the calculation of the volumes of sediments and rates of denudation.

Bahr E! Arab rift palaeoriver system As outlined in Fig. 4, the Bahr E1 Arab rift basin is a longitudinal basin extending in a N W - S E direction, it is nearly 1150 km in length, 600 km in width in its northern part and about 330 km in width in its southern part. It is surrounded by the high uplifted blocks; the Darfur dome in the north, the Nuba mountains in the east, the Congo-Nile

RIFT BASINS OF THE SUDAN [

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Fig. 13. Catchment areas and sedimentation troughs of Bahr El Arab rift. divide in the south and west, and the J. Marra massif in the NW. The high lands would be very much similar to the highlands outlined by Holmes (1965). Faure (1975) calculated the rate of uplift of the East African highs to be about 0.1 mm yr -~ in the mid Tertiary. Using this figure to calculate rates of uplifts in the Sudanese blocks, gives the following heights; 2500 m for the Congo-Nile uplift, 3800 m for the Nuba mountains, 1500 m for the Darfur dome and 2000 m for J. Marra massive. The highs were eroded, as shown by palaeochannels, depositing sediments in Bahr E1 Arab troughs. The recent studies have shown that all the existing wadies, streams and rivers are structurally controlled. It also showed that these rivers occupied their present day courses since the rifting started. It is therefore logical to use the recent fiver system as a replica to represent the old palaeochannels in the following discussion (Fig. 13).

Wadi El Kuu This wadi drains the eastern sides of J. Marra, southern parts of Tagabo, and flows southward to join Abu Gabra trough near Muhagriya, draining an area of 100,800 km 2. This river moved large amounts of sediments from the slopes of J. Marra to be deposited first along W. El Kuu in the Sag El Naam trough (Table 3), where most of the coarse sediments were deposited. After some filling of the basin, the medium and fine sediments were deposited in the Abu Gabra trough. This is indicated by the conglomeratic sediments deposited in Sag E1 Naam, whereas the fine sediments and clays were deposited in the central part of Abu Gabra trough. Characteristics of the sedimentological sequences from the boreholes in Abu Gabra trough, suggests that the area between Ghazalla Gawazt, Daein, Dar Es Salam and Gellabi is a distal fan of Wadi E1 Kuu (Table 3).

122

R.B. SALAMA

Table 3 Bahr El Arab sediments Locality I-

Sediments

Mode of deposition

Source

Distal fan

J. Marra

Distal fan Conglomeratic

J. Marra J. Marra

Distal fan

J. Marra

Conglomeratic alluvium Distal fan

J. Marra J. Marra

Intermediate fan Lake or marginal lake deposits

Nile/Congo divide Nile/Congo divide

Intermediate or distal fan

Nile/Congo divide

Wadi El Kuu sediments

Dar Es Salam El Gellabi Abu Gabra Sag El Naam

Clays, dark greyish to green, with small basement fragments, rounded iron concretions, volcanic chips and kanker nodules Clays and sands, dark grey clays, veneers of fine whitish sand Coarse sand, gravels and boulders

H - Bulbul, Ibra and Nyala sediments

Buram Buram Bulbul

Dark green gravelly clays, with felspathic fragments, rounded iron concretions, volcanic chips and kanker nodules Well-rounded gravels, pebbles and conglomerates Clays, silts and sand

III - Bahr El Arab sediments

Gileizan Gileizan

Gravelly clays, pebbles and sands 10 m of gravelly carbonate deposits

I V - Bahr El Ghazal sediments

Abyei

Clays, with few rounded gravelly particles and small rounded iron concretions

V - Sobat River sediments

Malakal

Clear washed sands with laminated clays; greyish and greenish clays; fine sands, clays and silts

Lake deposits

V I - Wadi El Ghalla sediments

Babanusa Targalla

Reddish sandy clays, with gravel Green to grey clays

Intermediate fan Distal fan

Nuba Mountain Nuba Mountain

After Salama (1987).

Wadi Ibra, Wadi Bulbul and Wadi Nyala W. lbra and W. Bulbul drain the southern and southeastern parts of J. Marra, flowing southwards and joins Bahr E1 Arab, whereas W. Nyala terminates 20 km southeast of Nyala town in a delta formation (Salama, 1971). The thickness of the sediments in the Qoz Dango trough indicates that the three wadies were active for a long period. The sediments are formed of a mixture of distal fans, channel deposits, shallow lacustrine and floodplain deposits.

area of the two rivers is larger than the catchment area of the Blue Nile. Before erosion this area would have had the same elevation as the Abyssinia plateau. If the amount of rainfall would have been the same, it follows that the amount of discharge and sediment load would be approximately similar to today's rates of discharge of the Blue Nile and the sedimentary load would be the same. The sediments are mainly alluvial and fluvial flood plains and lacustrine deposits.

Bahr E! Arab palaeoriver system Bahr E1 Arab drains 60,800 km 2 along the western borders of Sudan and Central African Republic. This area is the NW uplifted dome which formes the Nile Congo divide. It is estimated that this uplift was 2-3 km higher than today's elevation. A detailed description of Bahr E1 Arab sediments is given in Table 3. It is formed of alluvial, intermediate and distal fans, marginal and lake deposits and shallow lacustrine deposits. In Gileizan area, thick carbonate deposits were recorded (Salama, 1985a) from shallow wells and drilled boreholes.

Sobat River This river drained an area of 86,400 km 2 on the western side of the Abyssinian heights. About 86,400 km 3 of sediments are estimated to have been deposited by the Sobat river in the eastern Bahr E1 Zaraf trough. The remainder of the sediments contributed to the filling up of the White Nile trough of the White Nile rift.

Bahr E! Ghazal and Bahr E! Jebel Both these rivers drain an area of 340,800 km 2 (not including the Victoria Nile catchment, which is considered not to have been connected to Sudan prior to 12,000 yr BP). The combined catchment

Wadi El Ghalla Wadi E1 Ghalla drained the western and southern parts of the Nuba mountains, which are mainly metamorphic and sedimentary rocks of Palaeozoic and Mesozoic age. This may explain the arenaceous type of fluvial sediments comprising well rounded sands and gravels mixed with lateritic fragments, sandstone and mudstone clastics, high percentage of iron and the absence of volcanic fragments.

RIFT BASINS OF THE SUDAN

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White Nile rift palaeoriver systems The White Nile rift basin is longitudinal in shape and extends in a N W - S E direction. It is about 520 km in length and 400 km in width at Umm Ruwaba trough and 360 km in width in the White Nile trough (Fig. 14). The Nuba mountain northern watershed area deposited most of its sediments in the northern part of the basin through Khor Abu Habil, while the southern part was filled from the eastern highs through Khor Yabus and Khor Adar.

Khor Abu Habil Khor Abu Habil drains the northern area of the Nuba mountains, and flows in a NE direction towards the White Nile at Kosti, where a palaeodelta is known to exist (Gunn, 1982). All the sediments derived from the Nuba mountains were deposited in the Umm Ruwaba trough.

The fine deltaic deposits of Khor Abu Habil were deposited in the area between 15 km north of Tendelti and Tendelti. Coarse sediments were deposited in the deeper parts of the troughs, and are encountered at various depth intervals from the surface to 600 m (Salama, 1985a, b). The sediments deposited by Khor Abu Habil are characterised by the absence of volcanic clasts and the high percentage of coarse material in the margin of the fault zones at a depth of 200-300 m. The high percentage of fines and clays in the fan area and the presence of salt layers (evaporites) in the fan delta area. There are several published descriptions of the sediment types in the Umm Ruwaba graben. Most workers suggested that the Umm Ruwaba sediments are fluviatile and lacustrine (Andrew, 1948; Rhodis et al., 1964; Shafie, 1975; Salama and Salama, 1974). Whiteman (1971) suggested that the sediments were laid down in a series of land deltas.

124 Salama (1985a) showed that the Umm Ruwaba sediments have multiple origins with layers of very fine sand (possibly aeolian), and layers of gypsum and evaporites reported in several boreholes.

Khor Yabus and Khor Adar These wadies, which drain the western side of the Ethiopian high lands, cover the area between the Blue Nile and the Sobat fiver catchments that is almost 100,000 km 2. The sediments varied from alluvial, distal fan, swamps and lake deposits. Blue Nile rift palaeoriver system The Blue Nile rift basin is a longitudinal basin extending in a N W - S E direction, it is about 350 km in length and ranges in width from 80 km in the northern part to 180 km in the southern part. The Blue Nile drains the central and northern part of the high Abyssinian plateau (Fig. 15). It is joined in Sudan by the Rahad and Dinder tributaries. The sediments in the Blue Nile rift, have been studied by different workers, e.g. Andrew (1948), Kheiralla (1966), Williams (1966), Whiteman ( 1971), Williams and Adamson (1973), Williams and Adamson (1980), E1 Boushi and Abdel Salam (1982), Adamson et al. (1982) and Salama (1985b). Williams and Adamson (1980) gave the following description: "With few exceptions, the late Pleistocene and Holocene sediments bordering the present Blue Nile between Sennar and Khartoum represent an upward-fining fluviatile sequence. The Holocene alluvium consists of dark, alkaline cracking clays rich in sub fossil shell fragments; and the late Pleistocene sediments generally comprise current-bedded fine, medium and coarse sand, with extension outcrops of massive calcium carbonate in bank sections between Hasaheisa and Khartoum, and more localised pockets of water transported volcanic ash". Abdel Salam (1966) recognised three subdivisions which were not consistent with lithological logs of more than 500 wells, studied in the Gezira area (Salama, 1985b) and can be seen in Fig. 16a and b. The deposits were formed of layers of clay, sand clay, clayey sand, sand and gravel, which can not be separated into divisions. A feature that is characteristic of alluvial fans and deltas. The clays are alkaline, dark in colour, and low in organic matter (Ruxton, 1956). They contain a high proportion of unweathered minerals such calcic plagioclase, titan-augite, hornblende, and brown biotite (Andrew, 1948). The sand is composed of angular to subangular quartz grains, with calcium carbonate nodules and mica. The gravels are predominantly quartz along with fragments of metamorphic and igneous rocks, basalts and agates as well as kanker nodules.

R.B. SALAMA

River Atbara rift palaeoriver system The River Atbara rift is formed of two main basins the Atbara and the Gash Delta basin separated by a ridge. The Atbara is a longitudinal basin extending 180 km in a N W - S E direction and 7 0 100 km wide. The triangular shaped Gash delta basin is 600 k m 2 in area. Two main rivers supplied sediment to the basins; the River Atbara and the River Gash flowing from northwestern parts of the Abyssinian plateau. They are both seasonal streams, having their annual floods during the rainy season from June to September. The sedimentary column indicates high energy deposition by the Atbara at its initial stages. The alluvium varies in thickness from 50 to 200 meters, comprising gravels, sands and clays with the arenaceous materials forming more than 50%. The sediments that fill the Gash delta are mainly alluvial deposits ranging in thickness from 50 to 100 m. They are intercalated beds of unconsolidated coarse to fine grained gravel, silts and clay with the fine materials, silt and clay, increasing downstream.

Sediment loads, estimated volumes and source scenarios As mentioned earlier the variables that affect sediment yield of a catchment area are relief, runoff, precipitation distribution and vegetative cover. Interchanging the individual extremes of these four variables will affect the sediment yield from one extreme to the other. Postulating any of those variables into the geological past can produce different end products. Nevertheless, I will try here to reconstruct the volume of sediments which have been massed by these palaeoriver systems. The sediment load of the river Atbara was estimated (Hurst and Phillips, 1931) to be about 14 million tonnes (m.t.) annually, which gives a denudation rate of 0.04 millimetre per year (mm yr -~ ). The Blue Nile and its tributaries are characterised by a very high suspended sediment load during the flood season ranging from 4000 mg 1-l in August, to 100 mg 1-l during the recession period (El Badri, 1972). The sediment load of the Blue Nile was estimated to be 41 m.t. annually (Hurst and Philips, 1931), giving a denudation rate of 0.05 mm yr -1 , while the present rate of erosion from the Blue Nile and Atbara was estimated to be about 120 and 24 cubic metre per kilometre per year (m 3 k m -2 yr -l ) respectively (El Badri, 1972). McDougall et al. (1975) used comparatively lower figures ranging from 6 to 15 m 3 km -2 yr -l in their calculation of the denudation rates of the Trap series in Ethiopia, to calculate the volume of sediments deposited in the Nile Delta. McDougall et al.

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R.B. SALAMA

ily account for the volume of sediments estimated by McDougall et al. to have been deposited in the Nile cone. This would change completely the hypothesis put forward by several other authors postulating that I

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the sediments in the Nile delta are of Abyssinian origin. An average erosion rate of 26 m 3 km -2 yr -l was used in this study for the calculation of the sediment yield, taking into account the effects of desertification either man made or natural, on today's rate of erosion, and the erosion rates of the Blue Nile and Atbara during the last 70 years. It is assumed that the rate of erosion during the Tertiary assuming this to be constant, is of the same order as the erosion rate during the last 70 years. The high erosion rates are caused by cyclic dry and wet periods during the Tertiary, as compared to man made desertification in recent times. In the White Nile rift, the sediments from Khor Abu Habil, Yabus and Adar (Table 4), are not sufficiently large to account for the sediments in the White Nile trough. Additional sources are required to fill up the basins. The Sobat River is a possible alternative due to its proximity, although the presence of Akobo ridge makes it a remote possibility. On the other hand a shortage of sediments in Bahr E1 Arab rift (Table 5) would occur if the sediments of the Sobat River are not included. The other alternative would be the White Nile sources in equatorial lakes. However, as mentioned previously, the N i l e Congo divide would be of sufficient elevation to obstruct water moving northward from the equatorial lakes" furthermore Doornkamp and Temple (1966) suggested that the equatorial lakes were discharging in a westward direction. The only other alternative source would be a contribution from the Blue Nile if the gradients so allowed. It was found by a study of the gradients from the Blue Nile and White Nile, that the Abu Habil trough of the White Nile could feasibly be a sedimentary trough for sediments from the Blue Nile river (Fig. 17a and b). This would account for the additional sediments which McDougall et al. (1975), included in the Nile cone, but which Said ( 1981 ) excluded.

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RIFT BASINS OF THE SUDAN

129

Table 4 Catchment areas, volume of sediments, rate of deposition and period of deposition in White Nile rift River or Wadi

1 Catchment area (km 2)

2 Sediments (kin 3)

3 Period of deposition (my)

4 Rate of deposition (ram yr - l )

80800 12960 86400

23000 57600

12.0 17.0

0.04 0.06

80800 12960 46400 40000

23000 57600

12.0 17.0

0.04 0.06

White Nile rift (alternative A): 1. Abu Habil 2. Wadi Adar and Yabus 3. Unknown (probably the Blue Nile)

While Nile rift (alternative B): 1. Abu Habil 2. Wadi Adar and Yabus 3. Sobat 4. Blue Nile (!) (2) (3) (4) (5)

Rate of denudation 26 m 3 km -2 yr -I. Column I direct measurement of catchment areas from topographic sheets. Column 2 calculated from average thickness of Tertiary deposits. Column 3 calculated by multiplying column 1 by denudation rates and dividing the result by column 2. Column 4 calculated by dividing the annual rate of deposition by the sedimentary basin area.

Table 5 Catchment areas, volume of sediments, rate of deposition and period of deposition in Bahr El Arab rift River or Wadi

1 Catchment area (km 2)

2 Sediments (km 3)

3 Period of deposition (my)

4 Rate of deposition (mm yr -I )

1. Wadi El Kuu 2. Ibra and Buibul 3. Bahr El Arab 4. Bahr El Ghazal 5. Sobat River 6. Wadi El Ghalla

100800 56000 60800 340800 256400 83200

92000 32000 28000 296000 151200 81600

35 21 18 33 22 38

0.14 0.09 0.17 0.15 0.14 0.10

(I) Rate of denudation 26 m 3 km -2 yr -l. (2) Column I direct measurement of catchment areas from topographic sheets. (3) Column 2 calculated from average thickness of Tertiary deposits. (4) Column 3 calculated by multiplying column 1 by denudation rates and dividing the result by column 2. (5) Column 4 calculated by dividing the annual rate of deposition by the sedimentary basin area. After Salama (1987).

An alternative scenario, would be the contribution of sediments from both the Sobat and Blue Nile rivers as shown in Table 5. As concluded from the previous section, following the cessation of subsidence and infilling of the grabens and troughs of the Blue Nile rift during the Tertiary and possibly part of the Quaternary epochs, sedimentation proceeded to move southward towards the White Nile rift, to fill the Abu Habil trough. Similarly after the sediments of the river Atbara during the Tertiary and possibly part of the Quaternary epochs filled the Atbara graben, the flow started moving northward to Egypt.

N A T U R A L R E S O U R C E S OF THE RIFT BASINS

The rift structures of Sudan form the major ground-water basins, outside these basins ground

water is found only in small quantities and is of poor quality. The rift structures also forms the reservoirs for hydrocarbon accumulation; all the known oil fields of inland Sudan are located in these basins. The hydrological and hydrogeological closed basins of these rift structures led to the formation of saline lakes in the distal end of each of these structures. The formation of the River Nile is also closely linked with the formation of the Sudanese Rift System. Ground-water resources of the rift basins

Salama (1976), made the first attempt to divide the Sudan into ground-water basins. The division of the ground-water basins was based on the geological formation which formed the basins, and his classification, which depended upon the geological information available at that time, has changed significantly since. In this study, the basins previously

130

R.B. SALAMA

suggested by Salama (1976), will be modified to include the new structural patterns of the rift system discussed above.

Types of aquifers and aquifer characteristics Bahr El Arab rift ground-water basin. This basin (Fig. 18) extends over the Bahr El Arab rift and includes the Baggara and Sudd subbasins (Salama, 1976). The basin covers an area estimated to be about 377,000 km 2, which is 18% of the area of Sudan, and can be considered as the largest single basin in Sudan. Baggara subbasin extends over the Baggara graben of Bahr E1 Arab rift (Hunting Technical Services and Macdonald, 1976; Tohami, 1978; Salama, 1971, 1976). Ground water in this subbasin is usually found in saturated strata, which include sediments of the Tertiary and Mesozoic age. Most of the wells drilled within the graben intersected one or more of three types of aquifers. (1) Unconfined aquifer which covers most of the basin area, especially in the marginal areas where the recharge takes place from the wadies and through the contact zone between the faulted bed rocks. Several hand dug well fields are known to exist, and in those areas the water-bearing formation ranges from 10-30 m below the surface (i.e. Muhagriya; in the northern part at the delta of Wadi E1 Kuu, Tiwal in the west at Wadi Ibra, Aradeiba in the east at Wadi E1 Ghalla). From the records of percussion boring which was used extensively in the past, nearly all the logs of the drilled wells indicate some type of unconfined zone, which in most cases were not highly productive. The central part of the Baggara basin at Abu Gabra trough seems to be the area where this unconfined zone is not widely developed. (2) The second aquifer is semi-confined to confined aquifer mainly found in Tertiary sediments. It is characterised by thick deposits of clays, fine sandy clay and clayey sand. In the areas where the fine deposits are found above the coarser material, usually the water is under semiconfined conditions. (3) The third aquifer is a confined aquifer found mainly in the Mesozoic sediments at the boundaries of the graben. The confining layers are either mudstone layers within the Mesozoic deposits, or unconsolidated fine deposits of the Tertiary.

Several regional water level maps were produced for the basin (Salama, 1976; Hunting, 1976; E1 Tohami, 1978). The general trend of the ground-water movement is from the north, east and west towards the central part of Abu Gabra trough, where it forms a closed trough, which is continuous within the Sudd graben. Hydraulic gradients were calculated to be 1.2 x 10 -3 and 7.5 x 10 -3 for the eastern and western parts, respectively. Transmissivity ranges from 250 to 750 m 2 day -~, calculated for a saturated thickness of 50 m from the aquifer (Salama, 1976) (correcting for well losses and partial penetration). The analysis of pumping test data from 171 boreholes in the Baggara basin gave the values shown in Table 6. Storativity values ranged from 10 - 3 - 1 0 -4 for the second and third aquifers to 10-2-10 -3 for the unconfined alluvial aquifers (Salama, 1976). No long-term regional water level trends has been noticed in the basin, which indicates that the basin is in a steady state condition. Since the abstraction rates are very small compared to the immense amount of water stored in the aquifer; and no fluctuations were noted for seasonal recharge, the present day abstraction rates would not greatly affect the general water level trend.

Sudd subbasin. The Sudd subbasin extends over the Sudd graben of the Bahr El Arab rift (Salama and Salama, 1974; Salama, 1976; Geophysics and Strojoexport, 1977). The water is usually found in the sandy layer of the Tertiary sediments, yet due to the marked lithological variations and the high Table 7 Temperatures, pH and Eh of boreholes in Sewar and Qoz Dango troughs Location

Temperature pH

Eh

Trough

Gileizan Buram Ramies El Sonta Sewar El Fardos K. Abu Salama Kulkul

26.0 28.0 35.0 40.0 60.0 36.0 36.0 36.0

20.0 40.0 40.0 30.0 80.0 10.0 35.0 10.0

Qoz Dango Qoz Dango Sewar Sewar Sewar Sewar Sewar Sewar

7.1 7.4 5.9 6.4 5.5 6.2 6.5 7.2

Table 6 Transmissivity values in the Baggara aquifer complex in m2 day-l

Lower quartile Median Upper quartile Highest value

171 boreholes in Baggara basin

60 boreholes in lower aquifer

87 boreholes in upper aquifer

6.95 25.90 89.20 1205.00

10.37 106.00 503.00 1205.00

6.71 23.10 51.70 293.00

After Hunting Technical Services and Macdonald (1976).

RIFT BASINS OF THE SUDAN

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132 percentage of fines and clays, the aquifer is not homogenous. The Tertiary aquifer can be considered as consisting of small different water-bearing bodies (aquifers), with highly variable hydrogeological properties, separated by other water bodies (aquitards), which play an important role in the hydrogeological properties of the aquifer (Salama and Salama, 1974). Due to the presence of thick clay deposits in the northern part, south of Bahr E1 Arab some wells drilled by spudding to depths of 300 m, failed to reach a water-bearing stratum. Some hand dug wells in the vicinity of Bahr E1 Arab also proved to be dry. This rules out Bahr E1 Arab as a potential recharge source, at least in the southern part of Abu Gabra trough. Transmissivity was calculated on the basis of the lithology of sediments from the bore-logs, as compared with other similar areas in the Baggara basin. It was found to range from 25-50 m 2 day -l. Correcting for partial penetration and well losses T values ranging from 100-500 m 2 day -l, for 50 m of saturated aquifer thickness. These values are low, compared with the values of Transmissivity in the Baggara basin, but are acceptable taking into consideration the high percentage of fines and clays in the Sudd basin. Blue Nile rift basin This basin covers an area of 76,000 km 2. The basin has been divided into three subbasins, roughly coinciding with the structural divisions of the Blue Nile rift, these subbasins are: Khartoum, Gezira and Singa (Salama, 1976) (Fig. 18). Khartoum subbasin. Three main aquifer are identified in Khartoum basin (Andrew, 1948; Abdel Salam, 1966; E1 Boushi and Whiteman, 1968; E1 Boushi, 1972; Kheiralla, 1966; Saeed, 1974, Maimberg and Abdel Shafie, 1975; S.G.E.P., 1979). (1) Upper semi-confined Gezira aquifer; this is found mainly along the White Nile, where it is capped by a thin clayey layer, the aquifer is extensively utilised for irrigation purposes from hand dug wells. (2) Lower semi-confined Gezira aquifer; this covers the area between the White and Blue Nile, the northern part of Khartoum north to Sabaloka (Six Cataract), on the left side of the White Nile and the right side of the Blue Nile. This aquifer receives direct recharge from the Nile and is considered as the most highly utilised aquifer in Sudan. More than 500 wells tap this aquifer. The upper 10-15 m are fine sands and clays, while in the central part, south of the green belt the thickness increases. Water is usually found in a layer of sand and gravel ranging in thickness from 3 to 10 meters.

R.B. SALAMA (3) Khartoum Mesozoic semi-confined aquifer; the Mesozoic deposits forming the aquifer, crop out in the area west of the White Nile, and east of Khartoum North. It is also found below the surface in all the other areas of the basin. Ground-water maps prepared for the basin shows the following characteristics (Abdel Salam, 1966; Kheiralla, 1966; Salama, 1976; Saeed, 1974; S.G.E.P., 1979): (a) Ground water moves away from the two rivers. (b) There are two ground-water troughs, one in the northeast of the Blue Nile and the other in the central area between the Blue Nile and White Nile. (c) The hydraulic gradient varies considerably. In the area adjacent to the River Nile, it ranges from 0.002-0.005 close to the Nile and 0.001-0.0009 at a distance of more than 5 km away from the Nile (S.G.E.P., 1979). (d) The influence of the Nile water levels on the ground-water levels fluctuations has been related to distance away from the Nile (Saeed, 1974, 1978). This was found to be 5 m at a distance of less than 200 m; below 1 m at a distance of less than 2 km; and no effect in observation wells more than 2 km away from the river. But this is not true over the whole area, in some places the effect does not occur beyond one kilometre, especially near the Nile at Kosti and Rebek where the effect is minimal. The fluctuations of the ground-water levels resuiting from high rates of abstraction are noticed in the irrigation areas, where levels are known to have dropped more than 10 m in the upper semi-confined aquifer, at the northern part of Khartoum Province. Transmissivity ranged from 38.4 m 2 day-~ to 5950 m 2 day -l (Saeed, 1974). Salama (1976) from the analysis of more than 500 pumping test data in the area, showed that the transmissivity ranged from 500 m 2 day -1 to 2000 m 2 day -~. With transmissivity increasing with depth S.G.E.P. (1979). Storativity values range from 10-2 to 10 -3 for the upper semi confined aquifer, 10 -3 to 10 -4 for the lower semi confined aquifer, and 10 -3 to 10-5 for the confined aquifer. Gezira subbasin. The Gezira subbasin covers the Wad Medani graben of the Blue Nile rift (Abdel Salam, 1966; E1 Boushi and Whiteman, 1968; E1 Boushi, 1972; E1 Boushi and Abdel Salam, 1982). The water-bearing strata in the Gezira are similar to those of Khartoum basin. The ground water occurs under semi artesian and leaky conditions, as a result of impervious layers of clays in the Gezira formation (El Boushi, 1972). The upper Tertiary semi-confined aquifer have transmissivity values ranging from 100-500 m 2 day -l, and storativity of 10 -2 to 10 -3. The lower Tertiary semiconfined aquifer transmissivity values of 500-1500 m 2 day -I

RIFT BASINS OF THE SUDAN and storativity of 10 -2 to 10 -3. The deep confined aquifer T of 300-2000 m 2 day -~ and storativity 10 -3 to 10 -5 . Ground-water level fluctuations ranged from three meters near the Blue Nile to nearly steady conditions near the central part. Singa subbasin. This subbasin covers the Singa graben of the Blue Nile rift. Detailed hydrogeological study carried out in the area show that these aquifers are similar to the types found in Khartoum basin. The ground water occurs under semiconfined and leaky conditions. The general direction of the ground water is away from the rivers towards the central part of the aquifer, with a general ground-water component from south to northwest. Aquifer characteristics showed marked similarity to the Gezira basin. White Nile rift basin This ground-water basin extends over the White Nile rift, covering an area of 100,800 km 2. It includes Umm Ruwaba basin and the northern part of Sudd basin (Salama, 1976). The basin is formed of three subbasins; Bara, Umm Ruwaba and the White Nile. Bara basin. Bara basin includes Bara trough of the White Nile rift. It is fault bounded on the north, east and west, and it is separated from the Umm Ruwaba subbasin in the southeast by Umm Dam ridge (Hunting Technical Services, 1970); extending in a NW-SE direction (Rhodis et al., 1963; Hunting Technical Services, 1970; Mabrook, 1972; Maimberg and Abdel Shafie, 1975; E1 Boushi et al., 1975; All, 1978). All the strata below the unconfined water table are water saturated. Those include the sediments of the Tertiary and Mesozoic deposits. The types of aquifer recognised in Bara basin are: (a) Unconfined aquifer: 20-30 years ago, most of the villages in the Bara basin have a system of open wells varying in depth from a few meters to about 20 meters, some of them would go dry during summer, or the yield would decrease significantly. These wells were abandoned after the drilling of the deep wells and the construction of a permanent water point. All these wells tap the unconfined aquifer, which seems to spread over the whole of the Bara basin. This unconfined aquifer is utilised for irrigation in the eastern part of the basin, in Bara and in the Kheiran district. The aquifer is formed of a few metres of gravels and sands, in some places the gravel is intercalated with clayey sand and sandy clay. (b) Tertiary semi-confined aquifer: The semi-confined aquifer lies below the unconfined aquifer and is separated from it by a layer of clays and fine

133 sandy clays and clayey sand, of variable thickness, but it reaches its maximum thickness in the southern part of the basin, where a thickness of more than 500 m is known, i.e Bara. Usually the water is under semi-confined to confined conditions. (c) Mesozoic confined aquifer: In the northern part of the basin, where the Mesozoic faulted blocks are overlain by the thick Tertiary deposits, the water is found under artesian conditions (Umm Balgei flowing well; E1 Boushi et al., 1975). Transmissivity values ranged from 100-500 m 2 day -~ and storage coefficient of 10-3-10 -4 with delayed yield effect. Umm Ruwaba basin. Umm Ruwaba basin covers the Umm Ruwaba graben of the White Nile rift (R.E.G.W.A., 1979; T.N.O., 1979). The basin is fault bounded and extends in a SE direction. The water-bearing formations in this subbasin are similar to those of the Bara basin, except for the fact that the Mesozoic aquifer is semi confined and not confined as in the Bara basin. The Mesozoic aquifer is restricted to Kosti area only, and unexpectedly, it is dry even near the White Nile and with thicknesses exceeding 200 m. Transmissivity values for the Mesozoic sediments are almost twice that for the Tertiary values. T ranges from a low of 12.5 m 2 day -l to 120 m 2 day-i for the Tertiary deposits. Water level fluctuations ranged from two meters near the Nile to about 50 cm two kilometres away (R.E.G.W.A., 1979). White Nile basin. This basin extends over the White Nile graben, it was previously considered as part of the Sudd basin (Salama and Salama, 1974; Salama, 1976; Geophysics and Strojoexport, 1977). The water-bearing formations are similar to those of the Umm Ruwaba basin, with one major difference; the Mesozoic sediments although encountered in few wells (Salama and Salama, 1974) does not seem to have the same aquifer properties previously known in the other subbasins. River Atbara rift basin This rift basin is divided into two subbasins; the River Atbara and the River Gash subbasins. The River Atbara subbasin extends north from the Abu Haraf water divide to the Atbara River, and covers an area of 23,896 square kilometres. Two types of aquifer are recognised: (1) Semi-confined river alluvium deposits extending along the River Atbara from Atbara town, south to Qoz Regeb. The top layer is usually clayey sand and sandy clay, followed downwards by gravels and sand. The water is usually under semiconfined conditions. (2) Semiconfined Mesozoic deposits; occurring in almost all

134 the other areas of the subbasin. It attains its maximum thickness at W. El Makabrab. Several observation wells along the river Atbara, show seasonal fluctuations of about 10 m. The semi confined aquifer has a T value of 100-1000 m 2 day-1 and storativity 10 -2-10 -4. The River Gash subbasin extends over the alluvial deposits of the river Gash in Kassala town, and extends downstream to cover the Gash delta to the north of Wagara. Previously the alluvial river deposits were considered as one aquifer only, which is separated sometimes by aquitard layers (Saeed, 1969; E1 Amin, 1979). Further detailed work (T.N.O., 1982), showed that the deposits form two aquifers; Upper and Lower aquifer, with a top clay layer and a continuous aquitard layer which separates the two aquifer layers. The depth to the aquitard layer varies from 6 m to about 30 m below ground surface, with an average depth of 12.5 m. Where the upper aquifer is not developed, the aquitard and the top layer may form one unit. The general direction of the ground-water flow is following the Gash river downstream towards the delta. The aquifer exhibit marked seasonal fluctuations (El Amin, 1979), the magnitude of which depends on the distance of the observation well from the Gash river and or whether the observation well is in direct hydraulic connection with the recharging river. Ground-water recovery at the banks of the Gash river amounts to about 10-12 m, this decreases rapidly to about 3 m one kilometre away, and to only 1 m two kilometres away (T.N.O., 1982). Transmissivity values ranged from 100-690 m 2 day -! (El Amin, 1979), which is low in comparison to other alluvial aquifers (Salama, 1971). The storativity ranged from 10 -2 to 10 -3.

The evolution of the ground-water flow systems Based on hydrogeological and hydrochemical data from the Sudd basin the general direction of ground-water movement is from the basin boundaries towards the central part (Salama and Salama, 1974). The water levels of the Sudd basin forms a closed trough, isolevel 300 m is the lowest level in the trough and the lowest reduced water level in all the rift basins. In Bara basin, detailed water level maps (Rhodis et al., 1964; Hunting, 1970; Salama, 1976; Ali, 1978) shows that ground-water movement is from the west to the southeast with a hydraulic gradient of 0.00053 (Salama, 1976). In Umm Ruwaba basin ground-water movement is from the east to the west in the northern area, and from the west to the SE in the southern part. It also shows two closed ground-water troughs at the

R.B. SALAMA northeastern and southem ends, the northern one at 350 m a.m.s.1, and the southern one at 330 m a.m.s.1 (Salama, 1976). In the White Nile basin the ground-water flow is completely reversed, in contrast to the direction of ftow in the northern area, the general direction of ground-water flow is from the west to the east with a major trough south of Kosti and another trough in W. Adar. The general ground-water movement in River Atabara basin is from the SW to the NE. There is a big trough at the northeastern part of the basin (Salama, 1976). A large ground-water trough exists in the central part of the Gezira basin (Abdel Salam, 1966; Salama, 1976). Ground-water flow in large sedimentary basins is controlled by four main processes (Verweij, 1993): (1) sedimentation in a subsiding sedimentary basin; (2) introduction of heat into a basin; (3) tectonic processes acting on a basin, and (4) infiltration of meteoric water in a sub-aerial basin. These four processes were at one stage or another working separately or together in the rift basins of Sudan. The rifting phases which formed the Sudanese Rift System caused the formation of deep basins. Due to the continuous subsidence in the basins, together with continuous recharge of meteoric water which was taking place at the aquifer boundaries (rift basins boundaries), caused the formation of a deep ground-water trough in each one of these basins. The general direction of ground-water flow in all these basins is from the basin boundaries toward the troughs which in all cases were occupying a distal end of a river system. Although no ground-water discharge is taking place at these troughs in the present time, it is logical to assume that in wet pluvial periods (which are well recorded) that these troughs would act as ground-water discharge areas. E1 Boushi and Abdel Salam (1982), in their discussion of the troughs of the Khartoum and the Gezira basins, mentioned that: " Those two troughs represent windows of replenishment to the Nubian aquifer. If water did not leak to augment the Nubian water, those troughs would have been filled long ago". According to the results of this work it is evident that this trough represent a replica of a palaeohydrologic system, and there is no connection between the Gezira aquifer and the Mesozoic sandstone aquifer. The ground water in the lower Mesozoic aquifer is under pressure, and it can seep upward, but the water from Tertiary sediments cannot move downward against the pressure. This is also reflected in the marked contrast between the water qualities of each aquifer. The existing hydrogeological pattern is a very old phenomenon, which has been established since

RIFT BASINS OF THE SUDAN the inception of the rift basins. It has been slightly modified after the filling of the sedimentary troughs, and subsequent recharge during the wet periods take place at a slow rate, separated by periods of no recharge, during the dry periods. The continuous subsidence of the basins, and the fact that the areas where the troughs are existing are the deepest parts of the basins, which coincide with the ground-water trough, proves that the ground water is a replica of the surface water pattern that was existing at the time of formation. Ground-water resources Several research workers studied the alluvial basins, as the collection of data from the existing hand-dug wells was less expensive, and there were problems of over-withdrawal due to the development around those basins (Iskander, 1967; Salama, 1971; Saeed, 1969; Hussein, 1975). All the other studies carried out by consulting firms, were concentrated in the problem areas, i.e. hard basement rocks. Since 1966, several detailed investigations have been carried out by research workers in some of the groundwater basins (either in part or in whole): Mabrook (1972), Saeed (1974), Mohamed (1975), E1 Tohami (1977). Most of these studies delineated basement boundaries, aquifer characteristics and used analytical methods to estimate safe yield of the aquifer systems. Three basins were studied in detail using groundwater models (Salama, 1985a), they are Baggara basin of Bahr El Arab rift, the Umm Ruwaba basin of the White Nile rift and the three basins of the Blue Nile rift. In all three cases two simulations were carried out. The first one with the present ground-water recharge and discharge rates (including abstraction for domestic and irrigation) and the second one with increased rates of abstraction. Separate sets of simulations were carried out to test certain aquifers for heavy abstraction rates for irrigation purposes and for the new urban sites which depends mainly on ground water (i.e. Daein, Babanusa, Muglad (Baggara basin), Bara, Umm Ruwaba, Tendelti (Bara and Umm Ruwaba basins)) and the heavy abstraction for irrigation in Khartoum Province. The important conclusions from this study are: (1) The ground-water resources of these basins are quite adequate to sustain the water requirements of the expanding rural and urban centres. In nearly all cases it was found that better well design, better well completion and deeper wells to tap high yielding aquifer layers are required. (2) That abstraction rates can be safely increased ten folds and in some cases 50-fold without causing extensive drawdown in the area. (3) On the other hand it was found that using

135 these aquifers for irrigation purposes, will cause heavy drawdowns. In Bara basin where recent recharge is decreasing due to the long trends decrease in rainfall, it is highly recommended that irrigation developments be phased out. In Khartoum province, all the recent studies which excludes the Blue Nile as a substantial source of recharge as previously claimed, show that the heavy pumping will adversely affect the water levels. (4) In the alluvial aquifers (W. Nyala, K. E1 Gash, Arbaat, Tokar) which supply domestic and irrigation water for these urban centres, the studies show that the ground-water resource is limited and the expansion of these centres has to be controlled. Petroleum resources of the Sudanese Rift System and the role of ground water in its migration and accumulation Petroleum discovery and resources In 1975, Chevron Overseas Petroleum Inc. started a major petroleum exploration operation in the west and southern part of Sudan. During their 12 year operation in Sudan they acquired vast amount of geological and geophysical data. These included extensive aeromagnetic and gravity survey, 58,000 km of seismic data and drilled 86 wells (Schull, 1988). These extensive detailed investigations shed more light on the history and development of the rift basins of Sudan. The productive and prospective structures resulting from extensional movement and the compressional forces created within resulted in a complex structures created by rotated fault blocks, drape folds, and reverse drag folds. These structures created producing oil traps. The reservoir rock range from quartz arenites and wackestones to arkosic arenites and wackestones. These include sandstones deposited in fluvial channel, lacustrine delta-plain-distributary channel, and delta front environment. Schull (1988) summed up the characteristics of the reservoirs from data compiled from 30 cored wells; reservoir quality decreases with increasing depth (due to compaction, quartz overgrowth), with decreasing grain size and with increasing amounts of feldspars and lithic grains. Petroleum was discovered in the three explored rift systems; Bahr El Arab rift, in the White Nile rift and in the Blue Nile rift. The first oil was recovered from a well in the Muglad basin, the first significant oil flow occurred in Abu Gabra basin and the first important discovery was made in the Unity basin (Anon, 1981a, b, 1982; Schull, 1988). Oil was discovered in Abu Gabra Formation Cretaceous sands in the Muglad basin, in the Bentiu Formation in Babanusa basin, in the sand deposits of Darfur Group in the Unity basin. In the Tertiary oil was discovered in the sands of the Areal Formation.

136 The lacustrine claystones deposited in suboxic environment provide good oil-prone source rock. The depositional environment of these claystones and shales are within large lakes distal from the primary elastic influx. The organic material deposited in the lake was preserved in the suboxic conditions. The primary sources are degraded algal and plant material. The lithological description of the sedimentary sequence of the Sudanese rift basins indicate that it has been deposited in shallow lacustrine environment, occupied by intermittent swamps, lakes, surrounded by continental edge delivering terrestrial organic material which is periodically exposed to subaerial degradation and water flooding. These conditions usually produce high wax crude oil, which is the case for Sudan oil. In the shallow parts of the recharge area, infiltrating meteoric water flush hydrostatic trapping positions, on the other hand the deeper parts of a recharge area are comparatively favourable for the entrapment of hydrocarbons. Lateral hydrocarbon migration towards discharge areas enhances the volumes of hydrocarbons available for entrapment in these areas. Biodegradation and water washing through continuous discharge resulted in an increase in the density of the residual hydrocarbons. Total organic carbon content of the source rock of Sudan averages 1.3%. The generated oils are paraffinic, low sulfur, high pour point. The recoverable reserve in Unity and Heglig areas has been estimated to be 250,300 million bbl (Schull, 1988). Several other wells have recovered significant amounts of oil during stem tests. Flow rates as high as 4000 BOPD on a 5 cm choke have been measured. All oils have low gas/oil ratios and high pour points (Schull, 1988).

Hydrocarbon migration and accumulation through ground-water flow in the rift basins Three major mechanisms control the primary hydrocarbon migration (Verweij, 1993): (!) primary migration of continuous separate phase hydrocarbons driven by hydrocarbon potential gradients; (2) ground-water driven primary migration of hydrocarbons in aqueous solution, and (3) diffusion-driven primary migration of hydrocarbon in aqueous solution, and diffusion-driven primary migration of hydrocarbons through organic matter network. On the other hand secondary phase hydrocarbon migration is driven by hydrocarbon potential gradients which in turn are controlled by: magnitude and direction of the force of gravity, densities of the hydrocarbon and the ground water, magnitude and direction of the net driving force for ground-water flow and magnitude and direction of the capillary pressure gradient. The actual rate of secondary hy-

R.B. SALAMA drocarbon migration is controlled by the hydrocarbon potential gradient, the density and viscosity of the hydrocarbons and the effective permeability of rocks to hydrocarbons. On a basin wide scale, the pattern of secondary migration is determined by (Verweij, 1993): The hydrogeological frame work of the basin; the hydrodynamic condition of the basin and the associated ground-water flow systems in the basin; the density differences between the hydrocarbons and water. Under hydrostatic conditions, the hydrocarbons will become trapped in the reservoir rock when buoyancy-induced lateral upward hydrocarbon migration in the carrier-reservoir rock is stopped by a capillary pressure boundary. Hydrostatic trapping positions include structural traps, stratigraphic traps and combination traps. Hydrodynamic conditions affect the sealing capacity of a rock or a fault and consequently influence the holding capacity of hydrostatic traps. Vertically downward ground-water flow may increase the resistant force to hydrocarbon movement of certain layers and make them impermeable to hydrocarbons, creating hydrodynamic trapping possibilities. This does not seem to be the case in Abu Gabra trough where hydrocarbon signatures have been noticed in the shallow top 1000 m. Which coincide with a relatively high salinity zone in the aquifer and from the ground-water flow direction a possible discharge area. The general pattern of ground-water flow systems which controlled hydrocarbon migration and accumulation in the rift basins can be summarised in: (a) Several studies have been published showing the relation between gravity induced ground-water flow and hydrocarbon accumulation (Toth, 1980; Toth and Corbet, 1986; Toth and Otto, 1989). The results of the study suggest that long distance lateral migration of hydrocarbons can be explained by a basin-wide gravity-induced ground-water flow focusing ground-water and hydrocarbons into laterally continuous hydrogeological units and provide an additional driving force to transport the hydrocarbons laterally across the basin. The relationship between the location of known hydrocarbon accumulations and the regional hydrodynamic condition has been identified by several authors (Toth, 1980; Toth and Otto, 1990). Favourable regions for accumulation and entrapment of hydrocarbons are created by the combined influence of buoyancy forces, net driving forces for ground-water flow and capillary forces. Effective recharge in the Sudanese Rift basins is taking place mainly along the basin boundaries, while discharge is most probably taking place in the trough areas of the closed basin, although there is no contemporary evidence of discharge. The lithologi-

RIFT BASINS OF THE SUDAN cal and chemical evidence indicate that ground-water discharge was taking place at these areas at different time during the pluvial periods. (b) Tectonic processes may also influence both the ground-water pressure condition and the hydrogeological framework in a basin. These forces might lead to the escape of ground-water and hydrocarbons from such tectonically geopressured zones vertically upward along fractures and active faults. Oil contaminated ground-water has been noticed in several wells in the Abu Gabra basin, with some oil showing during shallow drilling for ground-water in Abu Gabra trough. This area is also associated with a fault system which generated high temperature ground water along the fault area. (c) Generally, the frequency of hydrocarbon accumulations was observed to increase and be maximum in areas of ground-water discharge and associated stagnant zones (Toth, 1980). In the stagnant zone, ground-water flow is negligible and the separate phase hydrocarbons introduced into these zones, will be moved by vertically upward directed buoyancy forces alone. In discharge areas the ground-water flow is vertically upwards and consequently the net driving force for separate phase hydrocarbons is directed vertically upwards. This will eventually lead to the entrapment of hydrocarbons in available hydrostatic traps of sufficient sealing capacity. Or even in the absence of these traps hydrocarbons may be trapped by the opposite lateral hydrodynamic forces (Verweij, 1993). The gravity induced ground-water flow conditions enhances lateral migration of hydrocarbons towards ground-water discharge areas. In case of Sudan basins where each rift system is characterised by the presence of ground-water troughs which indicate a closed hydrogeological basin and possible ground-water discharge areas. These sites happens to be the best oil producing areas in the rift basins. Buried saline lakes of the Sudanese Rift System Salama (1985a, b, 1987, 1990, 1994) presented evidence for the presence of highly saline groundwater bodies that occupy the flowing end of each of the rift systems. These have been interpreted as buried saline lakes, sabkhas or playas (Fig. 19). The widespread presence of calcrete, kanker and other carbonate deposits, over, and in, most of the Tertiary deposits, showed that conditions were favourable for the deposition of carbonates. It is postulated that the shallow standing waters evaporated forming salt crusts. In the next flood, (fresh water) dissolved the most soluble salts; (NaC1 and NazSO4), and transported these towards the deepest part of the basin, leaving carbonates behind in the form of kanker nodules. This explains the wide

137 distribution of kanker nodules over the upper Tertiary horizons, and the concentration of the sodium, chloride and the sulphates in the saline zones. At the same time, in the deepest part of the graben which was always a lake, playa or sabkha, the lake water evaporated, and became increasingly saline. During dry arid periods the lakes were completely or partly evaporated thus creating layers of salts, which were later dissolved by ground water to form saline ground-water bodies. Lake Sudd Salama (1987) showed that a series of fresh water lakes exists at the edges of Bahr E1 Arab; Lake Keilak, Lake Abyad and Lake Kundi, together with the main Sudd lake which covers most of the central part of this river system. Based on hydrological data he concluded that if the Victoria Nile was not connected to the White Nile the Sudd would be a closed lake system. Independent evidence showing that the White Nile was not connected to the main Nile prior to 12,500 yr BP was presented by Shukri (1949) and Hassan (1975), based on mineralogical analysis of the Nile deposits and by Kendall (1969), Livingstone (1980) and Adamson and Williams (1980). All this evidence indicates that the Sudd depression of southern Sudan was a closed lake system (Fig. 20). Willcocks (1904) postulated that the ancient lake was 250 miles in length from north to south and that the Blue Nile flowed southwards to join this lake. Lawson (1927) elaborated the Lake Sudd hypothesis and was first to call it by this name. Ball (1939) developed the Lake Sudd hypothesis further, assigned a length of over 655 miles to the lake. Ball (1939) made calculations similar to those of Lawson (1927) and concluded that for a lake of these dimensions, an average evaporation rate of 3 mm day -! over the lake would be sufficient to dispose of all rain and river water entering the lake, since the average annual rate of evaporation from open water surfaces at Mongalla, Malakal and Khartoum are 3.0, 4.5 and 7.5 mm respectively (Hurst and Phillips, 1931). Berry and Whiteman (1968) and Whiteman (1971) went to great lengths to prove that Lawson and Ball were wrong. Salama (1987) agreed that the lake may never have achieved the size assumed by the pioneering workers, yet there is enough evidence to show that there was always a water body in the Sudd region (Berry, 1962; Salama, 1987). However, the lake was not formed by a dammed up Nile. It was formed in a closed basin caused by the block subsidence in Bahr E1 Arab rift (Salama, 1985a, b). The size fluctuated according to the palaeoclimatological events prevailing at the time. Salama (1987), using saturation indices of minerals and salinity parameters, showed that a saline

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RIFT BASINS OF THE SUDAN

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lake occupied the central area of the Sudd. He estimated that the lake fluctuated in size from a maximum of 32,000 km 2 to a minimum of 336 km 2 (the area enclosed between isosalinity lines of 1000 mg 1-l and 30,000 mg l-l; Fig. 19). The depth of the lake ranged from few meters at the borders of the lakes to about 200 meters in the central part (calculated from ground-water gradients, versus thickness of saline zone). The presence of thick carbonate deposits at Bahr E1 Arab, suggest that the size of the lake is larger than the figures calculated or perhaps indicate the presence of another separate saline lake in that part. The age of the lake can be estimated from the amount of salt deposited as related to annual flow, assuming that the salt content of the inflow was 120 mg 1-~. Although this is very high, even for today's rates assuming that the palaeoflow would

range between 2 and 20 km 3 annually, then the lake age would range from 46,000 to 110,000 years for a lake area of 336 km 2 and 800 km 2, respectively. W h i t e N i l e rift, t h e N u b a s a l i n e l a k e s

The Nuba buried saline lakes extend from near Ed Dueim in the north, to W. Adar in the south (Fig. 20). They were formed by the palaeo Abu Habil and wadi Adar drainage system. Today the mean annual flow of Abu Habil is 100 million m 3, while that of wadi Adar is estimated to be 1.5 to 2.0 • 109 m 3 (N.C.R., 1982). ERTS.1 satellite imagery (October 1972) shows a large alluvial fan on the left bank between Kosti and Keri Kera with numerous small meandering distributaries which has been formed by Khor Abu Habil (Gunn, 1982). Salama (1985b) showed that the trough which occupies the central part of the White Nile rift is

140 the lowest water level contour (320 m) in the central part of Sudan. This indicates that the Nuba lakes received water from the Gezira and Sudd lakes at one time or another during its depositional history. This also seems likely as the very slow deposition rates of Abu Habil compared to the White and Blue Niles system, always kept the Nuba lakes area at a lower depositional level. Salama (1985b) using salinity parameters and saturation indices postulated that the area enclosed by the isosalinity line of 1000 mg 1-~, defines the margins of the saline lake zone. The high salinity zones within this saline zone, represent four smaller separate buried saline lakes. Each of these lakes occupied one of the troughs within the rift structures of the Umm Ruwaba graben (Abu Habil saline zone, in Abu Habil trough, East and West Er Rawat saline zones, in East and West er Rawat troughs). Another saline water body along W. Adar, in the Adar trough, with salinities higher than 10,000 mg 1-~, indicates the presence of another saline lake in this trough. The presence of high sulphate concentrations, in the Nuba lakes show that it had similar depositional environment to that at the Sudd lake. Williams and Adamson (1980), in their study of the late Pleistocene evaporites near Esh Shawal indicated that the absence of clastic components, the well ordered crystalline structure and uniform elevation of the evaporites, and their radiocarbon ages all suggest prolonged late Pleistocene evaporation of an extensive body of still, saline water along what is now the White Nile flood plain. The rough accordance in age between these White Nile evaporites and those that crop out in the Blue Nile south of Khartoum may indicate prolonged aridity and an absence of outflowing drainage in the Gezira at this time. The radiocarbon ages are more than 40,000 years (Adamson et al., 1982). Blue Nile lakes Salama (1985b) showed that the area west of Ed Dueim is characterised by the presence of recent sand dunes; below those sand dunes thick layers of evaporites are present. The analyses show that the deposits are very similar to the carbonate deposits, near Bahr E1 Arab (Salama, 1987). E1 Boushi and Abdel Salam (1982) noted that the presence of Gypsum and Carbonates in the Gezira sediments imply that saline waters accumulated in an internal drainage basin in the past. Adamson et al. (1982) reported the presence of thick carbonate deposits in the Gezira, they also showed that the conditions were favourable for the deposition of carbonates during much of the Pleistocene and probably earlier. The presence of massive carbonate deposits or calcretes suggest prolonged and high input of dissolved carbonate under conditions

R.B. SALAMA suitable for its precipitation. These were sluggish stream flow and wide dispersal of the water at the downslope end of a fan such as the Gezira. Salama (1985b) postulated that the high salinity zone which is located in Matug trough and roughly coincide with the left branch of the palaeochannel of the Blue Nile (Figs. 11, 18 and 19) forms the Gezira buried saline lake. He also postulated that the lake size would fluctuate between a minimum of 250 k m 2 to a maximum of 15,000 km 2. From the presented evidence, it is clear that the Gezira depression was a closed basin for a long time during the Pleistocene. The cyclic pattern of wet and dry periods during the late Tertiary and Quaternary is well recorded (Gasse, 1977; Gasse and Street, 1978; Gasse et al., 1980) and also noted above. Adamson (1982) has shown that these cyclic events are applicable to all east and central parts of Africa and can even be extended to the north part of Africa. This cyclic pattern is reflected in the Gezira area as well as in all the other buried lakes in Sudan. Soba lake Williams and Adamson (1980), described the extensive presence of thick late Pleistocene carbonate deposits cropping out in the west bank of the Blue Nile some 5-8 km upstream of Khartoum. They also described the presence of saline, alkaline, calcareous, sandy and massive clays sediments in the area between Khartoum and El Masid. This saline zone also extends northward on both sides of the River Nile, the salinity increasing to very high levels near Sabaloka. Salama (1985b) showed that Soba lake covers the area extending from south of Sabaloka cataract in the north to the margins of the Khartoum graben in the south. Based on soil and ground-water salinity he postulated that all this area was covered by evaporite deposits. During wet periods the Blue Nile eroded its course in the Sabaloka cataract, the surface water passing through the northern outlet caused the leaching of salts from the top layers. This leaching is more noticeable around the Blue Nile areas. Relative age of the saline lakes and relation to East Africa lakes Is it possible to correlate the saline lakes events in Sudan with similar lakes in eastern Africa, Ethiopia and Afar. The available data are not adequate to make direct correlation, but in the following discussion, an attempt is made to correlate the possible events (Fig. 21). (a) All lac ages determined from ground water in the central parts of these basins, gave an age limit of above 40,000 yr BP (Salama, 1985b; Malmberg and Abdel Shafie, 1975).

RIFT BASINS OF THE SUDAN

141

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in Egypt and in the far north of the Sudan. He postulated that the more southerly basins drained either into the Red Sea, the Atlantic Ocean or internally. Butzer and Hansen (1968) concluded that the White Nile and Blue Nile basins certainly did not merge with the Saharan Nile before the early Pleistocene. Wendorf and Schild (1976) agree with Butzer and Hansen (1968) conclusions, and differ with De Heinzelin in his assumption of a late connection of the Ethiopian basin. Said (1981) supports De Heinzelin and gives further evidence that the Nile in Egypt is of local sources. Said (1981) stresses the fact that Egypt itself supplied most of the waters of the Nile during the early part of its history. He also concluded that arid conditions did not set in over the Sahara in general and Egypt in particular except during the Pleistocene. Previous to this epoch and during most of the Cainozoic, there is evidence that the climate in Egypt was wet. There was a good mat of vegetation, little surface denudation, and during several epochs, moderate to intense chemical weathering. The other theory is that the drainage from Ethiopia via rivers equivalent to the Blue Nile and the Atbara/ Takazze flowed to the Mediterranean via the Egyptian Nile since well back into Tertiary times (McDougall et al., 1975; Williams and Williams, 1980). Adamson (1982) concluded that the Nile has been supplying Ethiopian water and sediment to the Mediterranean since well back into the Tertiary, that is from several to many millions of years. He accepted the idea that several basins might have been present at the early stage, but the coincidence between the volume of debris eroded from Ethiopia and the volume of the Nile cone in the Mediterranean, argue for very early integration of sub Saharan drainage with the Egyptian Nile.

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(b) Malmberg and Abdel Shafie (1975) using 180 results of ground water from Bara basin showed that the ground water from the lower horizon is from rains of much wetter periods and the temperature during that time was - 4 to -8~ less. (c) Assuming equal rates of erosion and sedimentation throughout the sedimentary column during the Quaternary, the top 100-300 meters of sediments, which contain the saline zones, would require about 10,000 to 30,000 years to accumulate if calculated, using the rate of sedimentation of Yuretich and Cerling (1983), and 100,000 to 300,000 years if we used the average rate of 1 mm yr -1 . (d) The approximate age of the lakes was calculated from the salt content (Langbein, 1961), to range from 40,000 to 110,000 years. (e) From the rough estimates of ages made by sedimentation rates of (c) above and the age determined by the salinity data (d), the approximate age may be comparable with the ages of the lakes in Abhe (Gasse and Street, 1978). At the same time it can be correlated with the wet pluvial periods of Said (1981). From this limited evidence of age, it can be postulated that the formation of those saline water bodies, would most probably be in the period between 90-120 thousand years ago, i.e. within the late Quaternary.

The evolution of the River Nile There are two theories in relation to the age of an integrated Nile. The first one is that the integrated drainage of the Nile basin is of young age. De Heinzelin (1968) proposed that the Nile basin was formerly broken into series of separate basins, only the most northerly (the Proto Nile basin) feeding a fiver following the present course of the Nile

The Blue Nile in Ethiopia The great Abbai emerges from Lake Tana flowing southeast before it turns west and northwest towards Sudan to form the Blue Nile. Williams and Williams (1980) summarised the development of the Blue Nile; Ethiopia became updomed during EoceneOligocene and Miocene-Pliocene times, with major volcanic eruption between about 25 my and 19 my ago. Initial entrenchment of the Abbai-Blue Nile and Tekazze-Atbara fiver systems probably began during the Miocene or very Late Oligocene, and accelerated during Late Pleistocene times, with the eventual removal of some 100,000 km 3 of basalt, sandstone, limestone and granite from their basins. Gasse et al. (1980) have shown that during the last two to three million years extensive lakes covered parts of the Rift Valley and Afar. Studies at Hadar have shown the continuity of these lakes and of their

142

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general climatic setting through nearly a million years. Adamson and Williams (1980), in their synthesis of the development of the Ethiopian drainage system, concluded that the Tertiary uplift along the axes of the Red Sea and Ethiopian rift arches, together with the rifting itself, disrupted the eastward flow-

ing drainage. The western rim of the rifted arches formed a barrier which reversed the direction of river flow. Drainage over an area of some 180,000 km 2 was thus diverted westward into the Nile catchment. The increased elevation caused higher precipitation, greater stream flow, greater stream incision and greater sediment load.

RIFT BASINS OF THE SUDAN BAHR E f i A R ~ RIFT NO LOCATION

AI A2 k3 A3 A4 A5 A6 A7 A8 A9 A10 All AI2 k13 AI4 AI5 AI6 AI7 AI8 AI9 A20 A21 A22 A23 A24 A25 A26 A27 A28 A29 A]0 A31 A32 A33 A34 A35 A36 A3~ A38 A39 A40 A41 A42 A43 A44 A45 A46 A47 A48 A49 AS0

Abu G a b r a Abu M a t e r l q kbyel Abyad Akoke r i d g e Babanuea B a h r E1 A r a b B a h r E1 G h a z a l Baraka Bentlu (Bantlu) Bor Buram Dam Gamad D a r e s Salam Duk F a d l a t E1 D,a e l n El F a r d o s El P u l a F.I G l d a d l El G e l l a b l El Nahud (F.nNuhud) E1 S o n t a Gllelzan GoZ Dango J. Rllla (El Hllla) J. Meldob J. Marra J. TaJabo Jut (R. Jut) Abu Salama Kulkul Kurkur Hills Lake K e l l a k bake Rundl

143

BLUE NILE RIFT NO LOCATION BI B2 B3 B4 B5 B6 B7 B8 B9 BI0 Bll BI2 BI3 BI4 BI5 BI6 BIT BI8 BI9 S20 B21 B22 B23

Dlnder(R.Dlnder) Ed Duelm El A t s h a n Er Rahad(R.Rahad) Gezlra Hag Y o u a l f Rasahels~ Ingesenna Hills J. Biuyt J. Dall J . Doud J . Fau J . Mazmum J . Moya Kiteir Balla Khartoum Malakal Hatug Sabaloka Slnga Sennar Umm Udam wad Medanl

WHITE NILE RIFT NO LOCATION CI Abu Rab11 C2 Abu R u k b a C3 Adar C4 Akoraweng C5 Bara C6 El O b e l d C~ Er R a w a t CS J e b e l Kon C9 Jebel Zalata CI0 Kerl Kerr CII K h o r Abu R a b l l C12 KhOr A d a r C13 Khor Y a b u s c14 Kostl C1S Muba M o u n t a i n s CI& R a b a k CI~ R a h a d ( E r R a h a d ) C18 Renk C19 Tendel t l C20 lYmm Ruwaba C21 Wanklr

RIVER ATBARARIFT NO LOCATION

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J. = J e b e l K. - Khor W. - Wadl

Leben

Honqalla Ramies (Rammls) Sewar Sharer Sobat River sidecar SuE El Gamal Talbya Tarqalla Umm K e d a d d a Unity W. Bulbul W. E1 G h a l l a W. I b r a W. H y a l a W. Shellango

Plate 1 (continued). Alphabetical location map index.

The White Nile in Equatoria Doornkamp and Temple (1966) suggested in their investigation of the evolution of part of south Uganda, that the original course of rivers Katonga and Kagera were to the west. Regional earth movements during the Early to Middle Tertiary resulted in swell uplift in eastern and southeastern Uganda and the development of a tectonic depression along the site of the present Albert rift. Sediment accumulation and continued uplift during the Late Tertiary and Early Pleistocene reversed the westward flow into an easterly direction. Temple (in Williams and Williams, 1980) considered that the modern Lake Victoria basin is no older than late Middle Pleistocene on the grounds that it post-dates Miocene and Pliocene earth movements in the Kavirondo Gulf

and in the Albert rift, and contains sediments no older than the late Middle Pleistocene Nsongezi Series which have late Acheulian implements near the top of the sequence. Williams and Williams (1980) cautioned against regarding Lake Victoria as Pleistocene, Lake Albert as Mio-Pliocene and Ugandan drainage through into southern Sudan as Late Cainozoic. The presence of lake deposits up to 100 m above the present level of Lake Victoria discovered by Doornkamp and Temple (1966), definitely prove the damming of the U gandan drainage, at least to the 12,500 date given by Livingstone (1980), and supported by Adamson and Williams (1980), who postulated that it is possible that Lake Victoria itself had never supplied water to the Nile prior to 12,500 BE

144 Adamson and Williams (1980), agreed with Cooke (1958) that the Kafu, Katonga and Kagera rivers flowed westwards across Uganda before the rift influence, and that the present positions of Lakes Victoria and Albert were within the Zaire drainage system and ultimately discharging into the Atlantic Ocean. Adamson (1982) showed that the Nile drainage system originated with the rifting of Africa. He emphasised the fact that tectonic events of great age, probably dating back to the Precambrian, have exerted a powerful influence over the development of the Nile basin and its river system. Late Quaternary River Nile Adamson (1982), concluded that throughout the late Quaternary, the Nile behaved as an integrated river system with events downstream being determined by those in the headwaters. The characteristics of water discharge and sediment load which originated in the headwaters, in turn controlled the nature of aggradation, erosion and flooding on the low angle flood-plains far downstream. He clearly showed two important periods; the late Pleistocene arid period c. 20,000 to c. 13,000 BP and the terminal Pleistocene-mid-Holocene moist period c. 12,000 to c. 5000 BP. Butzer et al. (1972) in their study of the fluctuations of Lakes Rudolf, Nakuru, Naivasha, Magadi, Rukwa, Chad and Victoria, showed that all those lakes were high in the early Holocene between 10,000 and 8000 years ago. They showed that transgressions leading to this high stand began about 12,000 years ago, and evidence from three basins (Victoria, Nakuru, and Chad) indicates a pause or minor recession just at or before 10,000 years ago. Wherever information is available for the period preceding 12,000 years ago, it has consistently been shown that lakes are much smaller. Gasse et al. (1978) have shown that three lacustral phases can be recognised in Abhe; first the 60,000 years ago, second the 40,000 to 30,000 years ago and the third 29,000 to 17,000 years ago. They also showed that all the lakes transgressed rapidly after 11,000-10,000 yr BP reaching maximum elevations towards 9400 yr BP. Gasse et al. (1980) mentioned that the discovery of large tropical lakes at Hadar from 3.3 to 2.6 my has confirmed the picture emerging from East Africa (Olduvai, Omo, East Turkana, etc). The site has yielded an exceptionally fine collection of fossil remains, all of which indicate, as in East Africa, a Sudanian or Sahelian climate. Studies at Hadar have shown the continuity of these lakes and of their general climatic setting through nearly a million years. This period contrasts strikingly with the following one. Aridity set in between 2.5 and 2 m.y. in the Omo based on the pollen and rodents) and in the Afar, where it probably marks the onset

R.B. SALAMA of desert conditions. Williams and Adamson (1973), made the following summation; About 12,000 yr BP Lake Victoria overflowed, the level of the White Nile rose, and the dunes between Kosti and Ed Dueim became partly buried by alluvial clay. After 8000 yr BP the level of the White Nile fell. Dark clays accumulated in the swamps bordering the river until about 4000 yr B P. The fixed dunes of Kordofan and the even flow of the modern White Nile indicate the late Pleistocene dry period was more arid than the present semi arid climate of central Sudan, and it is possible that the White Nile may have ceased to flow into the main Nile, which would thus have been more seasonal in its regime than now The evolution of the River Nile and the buried saline rift lakes in Sudan From the discussion given in the previous sections, and from the discussions and evidence presented in this section, the writer is lead to accept the first theory which postulates that the integrated River Nile is of young age. The following conclusions support this case: (a) Each one of the rift systems, was a closed basin (Fig. 20). (b) That these basins were not interconnected, except after their subsidence ceased, and the rate of sediment deposition was enough to fill up the basins to such a level that would allow connection to take place. (c) That the Nuba basin was connected to both the Gezira basin and the Sudd basin, this can be shown from the cross-sections. The contour of 380 m above mean sea level, passing parallel to the White Nile, on the east and west banks, from south of Khartoum to Jebelin, makes all this area of the White Nile lower than the Blue Nile at any point south from Hasaheisa (Fig. 17). (d) It would be natural, with a gradient much higher than the existing White Nile gradients, that the Blue Nile would move westward towards the White Nile, especially if the exit from Sabaloka was choked or at a higher level. Which means that the palaeochannels of the Blue Nile which shows a southwesterly bend, were all moving towards the Abu Habil trough in the White Nile rift. Another proof of the movement of the Nile southward is the discovery of pure glassy volcanic ash by Abdalla and Adamson reported in Adamson et al. (1982), 1.6 m below the surface of Holocene White Nile clays south of Kawa on the White Nile. This volcanic ash is similar to the ash deposits discovered along the Blue Nile north of Wad Medani. These workers reached the conclusion that this ash is almost certainly entered the White Nile from the Sobat river or from the seasonal streams which drain the Ethiopian highs north of the Sobat. They also concluded that the deposits are possibly

RIFT BASINS OF THE SUDAN younger than the Blue Nile deposits, since they lie below a skin of Holocene clays. But from the new evidence available, it is now possible to interpret the presence of the volcanic ash in the White Nile as being deposited by the Blue Nile. There is nothing to prevent the Blue Nile waters from reaching this spot, as shown by the gradients of the Blue and White Niles. Also the interpretation of the age as younger than the Blue Nile ashes, based upon the thin thickness of the Holocene deposits, can be interpreted by the fact that the sedimentation rate of the White Nile rift is much slower; that is the reason for the shallow cover. Although Adamson and Williams (1982) gave an age of 25,000 years by dating carbonate deposits above them, it is most likely that the volcanic ashes were brought about by the wet period of 70,000 years ago. (e) This connection between the Gezira and Nuba basins, means that the overflow from the Blue Nile would move southward towards the Nuba basin and not northward towards the Sabaloka. (f) From the inscriptions on the rocks of Semna cataract (about 70 km south of Wadi Halfa) (Willcocks, 1904; Ball, 1939), it was assumed that the River Nile lowered its channel in the rocky barrier by about 8 m within the last 3800 years. That means about 0.0021 m yr -~. It is postulated here that the Semna cataract was rising at this rate, and that the Nile was cutting its way through its trapped course. It is also postulated that all the cataracts along the course of the River Nile are uplifts, adjacent to subsiding basins. It is not a coincidence that all the cataracts are the only Basement outcrops along the course of the River Nile. They are the result of isostatic adjustment of the subsiding basins. From the natural gradient northward, the flow started to move northward following the Tertiary rifting patterns; N W - S E and SW-NE. Due to the compensational effects, the uplifted areas started to rise, this rise in the area between the Sabaloka cataract in the south to the Semna cataract in the north started only 70,000-80,000 years ago. This can be calculated using the uplift rate from the Semna cataract. This age agrees with the date given by Said (1981) for the connection of the Blue Nile to the main River Nile in Egypt. It also agrees with the wet periods of Gasse (1977), which supports the case for an increase in sedimentation during the wet periods, filling of the basins and overflowing northward, so as to connect with the northern river system. (g) The Darfur dome in the north was the controlling ridge, that distributed the northerly and southerly water shed areas. As mentioned earlier in the section on the Sudanese Rift System, the Darfur dome highest point was in the west in the J. Marra area and that this elevation decreases in an easterly direction. It was also shown that the

145 intensity of rifting decreased in an eastward direction; which means that the River Atbara rift has the shallowest grabens. (h) Due to this difference in elevation, and to the difference in graben size to be filled by sediments, it is logical to expect that the rivers in the eastern part will be the first rivers to connect to the Egyptian drainage system. The River Atbara, was the first river to fill up its subsiding graben and moved northward towards the main River Nile. (i) The Blue Nile was always overflowing towards the Nuba basin. It was not connected to the main Nile, before 70,000-80,000 years ago. This connection occurred due to the damming effect of the outlet towards the south by sand dunes during arid periods. (j) The White Nile system, in Bahr E1 Arab and White Nile rift areas, remained a closed lake until, the connection of the Victoria Nile some 12,500 years ago. This can easily be proven using the hydrological budget, if the losses from the White Nile basin now are about 42 km 3 annually, and if the Victoria Nile is cut off, then there would be no outflow from the Sudd. The other proof is the presence of the evaporites on the Nuba lakes area, and all the areas, extending from south of Ed Dueim to near Renk which indicates that the troughs and grabens of the White Nile rift were occupied by a saline standing water body. (k) At the same time the present existence of the Sudd and the Machar marshes in Bahr El Arab and the White Nile rifts systems, indicate that these areas are lower than the normal gradient of the area to the north. These observations may indicate that the two areas are not yet fully filled with sediments. Another possibility is that these areas are still active, but are subsiding at a slow rate. This might agree with the continuous tremors occurring at J. E1 Rajaf near Juba, at the southern edge of Bahr El Arab rift (Whiteman, 1971; pers. observ., 1975) for Wadi E1 Melik and Wadi E1 Muggadam. The second one in Wadi E1 Qaab, west of Dongola, where the evaporites extends on the surface of depression for more than ten kilometres. This used to be the lake area of Wadi Howar. This shows, that this phenomenon of closed lake systems is existing northward and thus complete the pattern of the Sudanese lakes northward to the borders of Egypt. (1) The filling up of the saline lakes, or depressions, led to the connection of the Egyptian Nile with the Sudanese Nile, which captured the Ethiopian and Equatorial head waters during the tectonic activities that formed the rift structures of eastern and Central Africa and the Sudanese Rift Systems. (m) From all the above discussions it can be postulated, after making all the required reservations, that the age of the integrated River Nile is very recent. The River Atbara overflowed its closed basin

146 about 100,000 to 120,000 yr BP, The Blue Nile overflowed its closed basin 70,000 to 80,000 yr BP and the White Nile overflowed its closed basin 12,500 yr B P. (n) Williams and Adamson (1980), in their study of the late Pleistocene evaporites near Esh Shawal, made the following conclusion: The absence of clastic components, the well ordered crystalline structure and uniform elevation of the evaporites, and their radiocarbon ages all suggest prolonged late Pleistocene evaporation of an extensive body of still, saline water along what is now the White Nile flood plain. The rough accordance in age between these White Nile evaporites and those that crop out in the Blue Nile south of Khartoum may indicate prolonged aridity and an absence of outflowing drainage in the Gezira at this time. Adamson et al. (1982) interpreted the presence of the evaporites by the following: During the Pleistocene, microcrystalline dolomite and high Mg-calcite accumulated in what was evidently a desiccating White Nile lake, the extent of which remains unknown. The radiocarbon ages are more than 40,000 years (Adamson et al., 1982). (o) Alternating dry and wet periods are well recorded (Fig. 21) (Kendall, 1969; Gasse, 1977; Gasse et al., 1980; Wendorf et al., 1976; Adamson et al., 1980; Livingstone, 1980). The wet periods are characterised by high rates of deposition which partially fill up the basins, and form zones of fresh ground water within the saline water bodies. Due to the rapid sedimentation rates, the saline zones were quickly covered by sediments. These cycles eventually filled up the basins and led to their interconnection to form the existing Nile System. It is postulated (Salama, 1987) that Nile in Egypt was first connected to the Sudanese Nile System about 100,000-120,000 yr BP when the River Atbara overflowed its closed basin. This was followed by the Blue Nile which overflowed its basin 70,00080,000 yr BP and lastly the White Nile overflowed its closed basin 12,500 yr BP (Fig. 21).

CONCLUSIONS The deep lineaments and shear patterns of Sudan follow two main directions: NNW (Red Sea trend) and ENE (Gulf of Aden trend). Precambrian mobile belts trend NE and NW. Palaeozoic(?) sediments occupy N E - S W aligned grabens. Mesozoic continental sediments with NW palaeotrends were deposited in two major depressions also aligned NW. Cainozoic up-doming, volcanicity and tensional stress along NE and SW axes, associated with the movement of the African plate created rift structures in Sudan which were formed by successive block faulting along palaeotrends, followed by subsidence and lin-

R.B. SALAMA ear uplift. The Sudanese Cainozoic Rift System forms the largest rift system in Africa, extending from the eastern borders of Sudan with Ethiopia to the western borders with Central African Republic and Chad. It includes from west to east; Bahr E1 Arab rift, Wadi E1 Kuu rift, White Nile rift, Blue Nile rift and River Atbara rift. The thick Tertiary sediments filling the deep grabens were eroded from the elevated blocks; Jebel Marra, Nuba mountains Inqessena Hills, Darfur Dome and the Nile Congo divide. With continuous subsidence in the grabens and troughs, and continuous uplift in the flanking areas, hydrological and hydrogeological closed basins were formed in each of the rift systems; the Sudd in Bahr E1 Arab rift, the Nuba in the White Nile rift, the Gezira and Soba in the Blue Nile rift, the Atbara and Gash in the Atbara rift. At the flowing end of each of the rivers discharging in the rift system, alluvial fans were formed; Gash and Atbara fans of the River Atbara and River Gash of the River Atbara rift; Soba and Gezira fans of the Blue Nile river in the Blue Nile rift system; Abu Habil and Mashar fans of Abu Habil and Wadi Adar rivers in the White Nile rift; W. El Kuu, W. El Ghalla, W. Shallengo, W. Nyala, W. Bulbul, and E1 Sudd fans of W. E1 Kuu, W. E1 Ghalla, W. Shallengo, W. Nyala, W. Bulbul and Bahr el Arab river system in Bahr El Arab rift. The deltas and fans were always found at the distal end of the graben or trough, against an uplifted block of Basement, which seems to act as sill or dam, forming a hydrologically closed basin, causing the deposition of the sediments, the formation of evaporites and/or the saline lakes. The study of the ground-water salinity in the rift systems showed that a saline ground-water body or bodies exist within the Tertiary sediments, while the ground water in the lower Mesozoic sediments were found to be fresh. These saline water bodies are identified as: (a) Sudd saline zone in Bahr E1 Arab rift; (b) Nuba and Adar saline zones in the White Nile rift; (c) Gezira and Soba saline zones in the Blue Nile rift; and (d) Atbara and E1 Gash saline zones in the River Atbara rift. It was shown that the highly saline ground-water bodies coincide with the hydrogeologically closed basins, in each of the systems. The saline body or bodies have been formed by evaporation coupled with alkaline earth carbonate precipitation and resolution of capillary salts. It has also been found that all the closed basin systems, fans, deltas and saline ground-water bodies, lie in one line N - S / N E - S W , which is roughly parallel to the closed basin lakes in the East Africa Rift in Ethiopia, Kenya and Tanzania. This indicates similar conditions for the formation of both rift lakes. The East Africa Rift saline lakes still exist, because they are in the high uplifted areas from which sediments are eroded and transported. The Sudanese saline

RIFT BASINS OF THE SUDAN lakes on the other hand, are in the lowest part of another subsiding rift system, collected vast amounts of sediments. The rapid rate of sedimentation accelerated the rate of burial of the saline lakes under a thick cover of Tertiary sediments. On the other hand the rapid rates of deposition did not allow the lakes to b e c o m e highly saline. During dry arid periods the lakes were completely or partly evaporated to dryness, thus creating layers of high salinity, which were latter dissolved by ground water to form saline g r o u n d - w a t e r bodies. The alternating dry and wet periods caused the layering of the formations into saline and fresh zones, the wet periods are characterised by high rates of deposition which partially filled up the basins, and formed layers of fresh ground-water zones within and above the saline water bodies.

ACKNOWLEDGMENTS

I am indebted to m a n y people for the accomplishment of this work. First, I would like to acknowledge the generosity, cooperation and hospitality of my c o u n t r y m e n , the nomadic people of Sudan, who generously offered us all the necessary help required for our field operations during the period from 1966 to 1982. Second, I would like to acknowledge the assistance, cooperation and fruitful discussions of my colleagues, special thanks to M.S. E1 Tohami, K.M. Kheiralla, B. Karkanis, W. Iskander, S.I. Asaad, A.M. M o h a m e d , A.H Ishag, L. Wahdan, and to M.S. E1 Rabaa with w h o m I spent lengthy hours discussing the origin of these basins. Third, I ack n o w l e d g e the assistance of my wife Marsail and for her support and contribution to many of the ideas in this work. I would like also to thank C. Barber and M. C h u r c h w a r d from the CSIRO, Division of Water Resources for their constructive criticism of earlier drafts of the manuscript.

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R.B. S A L A M A Holmes, A., 1965. Principals of Physical Geology. Ronald Press, New York, N.Y., 1288 pp. Hunting Technical Services, 1970. A hydrogeological study of the Qoz resettlement area, Eastern Kordofan, Sudan. Report submitted to Rural Water Corporation, Khartoum (unpublished). Hunting Technical Services and Macdonald, M., 1976. Savanna development project, Phase 2, Annex 2, Part 1, Hydrogeology. Report submitted to Rural Water Corporation, Khartoum (unpublished). Hurst, H.E. and Phillips, P., 1931. The Nile Basin. General description of the basin. Phys. Dep. Paper, Govt. Press., Cairo. Hussein, M.T. 1975. Hydrogeological investigation of Khor Arbaat basin. Bull. Geol. Surv. Sudan, 28:161 pp. Jones, B.E, Eugster, H.P and Rettig, S.L., 1977. Hydrochemistry of lake Magadi basin. Geochim., Cosmochim. Acta 41: 53-72. Kabesh, M.L., 1962. The geology of Mohamed Qol Sheet. Mem. Geol. Surv. Sudan, 3:61 pp. Karkanis, B.G.Y., 1965. Hydrochemical facies of groundwater in western Province of Sudan. M.Sc. Thesis, Univ. of Arizona, Tucson (unpublished). Kendall, R.L., 1969. An ecological history of the Lake Victoria basin. Ecol. Monogr., 39: 121-176. Kheiralla, M.K., 1966. A study of the Nubian Sandstone Formation of the Nile Valley between 14N and 1742N with reference to groundwater geology. M.Sc. Thesis, Khartoum Univ. (unpublished). Klitzsch, E.H., 1983. Geological research in and around Nubia. Episodes, 1983, 3:15-19. Kr/Sner, A., Stern R.J., Dawoud, A.S., Compston, W. and Reischmann, T., 1987. The Pan-African continental margin in northeastern Africa: evidence from a geochronological study of granulites at Sabaloka, Sudan. Eart. Planet. Sci. Lett, 85: 91-104. Langbein, W.B., 1961. Salinity and hydrology of closed lakes. U.S. Geol. Surv. Prof. Pap. 412: 1-20. Lawson, A.C., 1927. The valley of the Nile. Univ. Calif. Chronicle, 29. Livingstone, D.A., 1980 Environmental changes in the Nile headwaters. In: M.A.J. Williams and H. Faure (Editors), The Sahara and the Nile. Balkema, Rotterdam, pp. 339-359. Malmberg, G.T. and Abdel Shafie, M., 1975. Application of environmental isotopes to selected hydrologic studies in Sudan. Int. Atomic Agency, Vienna (unpublished). McDougall, I., W.H.Morton and M.A.J. Williams (1975). Age and rates of denudation of Trap series basalts at Blue Nile gorge, Ethiopia. Nature, 254: 207-209. McKee, E.D., 1962. Origin of the Nubian and similar sandstones. Geol. Rundsch., 52: 551-587. Medani, A.H. and Vail, J.R., 1974. Post-Cretaceous faulting in Sudan and its relation to the East African Rift System. Nature, London, 248:133-135. Mohamed, A.M., 1975. Hydrogeology of Shagara basin with reference to the water supply of El Fasher town. M.Sc. Thesis, Univ. of Khartoum (unpublished). Mula, A.H.G., 1971. A geophysical survey of the Jebel Aulia region and a study of fault solutions of gravitational anomalies. M.Sc. Thesis, Univ. of Khartoum (unpublished). Mula, A.H.G., 1972. A geophysical survey of Jebel Aulia region. Sudan Notes Rec., 53: 164-168. N.C.R., 1982. Water Resources in Sudan. National Council for Research, Khartoum. Pomeyrol, R., 1968. Nubian Sandstone. Bull. Am. Ass. Pet. Geol., 52(4): 589-600. Qureshi, I.R., and Sadig, A.A., 1967. Earthquakes and associated faulting in central Sudan. Nature, 215: 263-265. R.E.G.W.A., 1979. Ed Dueim area, Hydrogeological study. Final

RIFT BASINS OF T H E SUDAN report. Report submitted to Rural Water Corporation, Khartoum (unpublished). Rhodis, H.G., Hassan, A. and Wahdan, L., 1963. Availability of groundwater in Kordofan Province, Sudan. Bull. Geol. Surv. Sudan, 12:16 pp. Ruxton, B.P., 1965. The major rock groups of the northern Red Sea Hills, Sudan. Geol. Mag. 3: 314-330. R.W.C., 1977. Hydrogeological map of Sudan. Piezometric surface map. Edited by R.B. Salama. Rural Water Corporation, Sudan. R.W.C., 1977. Hydrogeological map of Sudan. Water quality map (Total dissolved solids). Edited by R.B. Salama. Rural Water Corporation, Sudan. Salama, R.B., 1971. Geology and Hydrogeology of Wadi Nyala. M.Sc. Rep., London University. Salama, R.B., 1972. Salinity problem in Kiteir Balla area. R.W.C. Open File Rep. (unpublished). Salama, R.B.(1976), Groundwater resources of Sudan. R.W.C. Open File Rep. (unpublished). Salama, R.B., 1977. Groundwater resources of Sudan. In: Biswas (Editor), Water Development and Management. Proc. United Nations Water Conference, Mar Del Plata, Argentina. Part 4. Pergamon Press, p. 1796. Salama, R.B., 1985a. Buried troughs, grabens and rifts in Sudan. J. Afr. Earth Sci., 3(3): 381-390. Salama, R.B., 1985b. The evolution of the River Nile, in relation to buried saline rift lakes and water resources of Sudan. Ph.D. Thesis, Univ. of New South Wales (unpublished). Salama, R.B., 1987, The evolution of the River Nile. The buried saline rift lakes in Sudan w I. Bahr El Arab Rift, the Sudd buried saline lake. J. Afr. Earth Sci., 6(6): 899-913. Salama, R.B., 1990. The role of Megastructures on the development of salinity in the River Nile Basins. Proc. Int. Conf. on Groundwater in Large Sedimentary Basins, Perth. Aust. Water Resourc. Council Ser., 20: 288-297. Salama, R.B. 1994. The Sudanese buried saline lakes. In: M. Rosen (Editor), Paleoclimate and Basin Evolution of Playa Systems. Geol. Soc. Am. Spec. Pap., 289 (in press). Salama, R.B. and Salama, M.N., 1974. Geology and hydrogeology of the Southern Sudan. R.W.C. Open File Rep. (unpublished). Said, R., 1981. The Geological Evolution of the River Nile. Springer-Verlag, Berlin, Heidelberg, New York. Said, E.M., 1969. Hydrogeology of the Gash River. Geol. Survey Sudan, Bull., 19. Saeed, E.M., 1974. Geological and hydrogeological studies of Khartoum Province, Sudan. Ph.D. Thesis, Cairo Univ. (unpublished). Schull, T.J. 1988. Rift basins of interior Sudan: Petroleum exploration and discovery. Am. Assoc. Petrol. Geol. Bull., 72: 1128-1142.

S.G.E.P., 1979, Groundwater resources in Khartoum province. Sudanese-German Exploration Project Tech. Rep. Part 2. Submitted to the Mineral Resources Dep., Khartoum (unpublished). Shafie, A.I., 1975. Lithology of Umm Ruwaba Formation and its palaeogeography in connection with water problems. Ph.D. Thesis. Inst. Geol. Res., Moscow (unpublished). Shukri, N.M., 1949. The mineralogy of some Nile sediments. Quart. J. Geol. Soc., 105:511-531. S.I.K.R., 1983. Satellite image of Khartoum region. Regional Remote Sensing Facility, Nairobi. Sly, P.G., 1978. Sedimentary processes in Lakes. In: A. Lerman (Editor), Lakes, Chemistry, Geology, Physics. Springer-Verlag, New York, N.Y., pp. 65-84. T.N.O., 1979. Groundwater investigations, El Jebelein. Report submitted to Rural Water Corporation, Khartoum (unpublished).

149 T.N.O., 1982. Water resources assessment. Kassala Gash basin. Final report. Report submitted to Rural Water Corporation, Khartoum (unpublished). Toth, J. 1980. Cross-formational gravity-flow of groundwater: A mechanism of the transport and accumulation of petroleum (The generalised hydraulic theory of petroleum migration). In: W.H. Roberts III and R.J. Cordell (Editors), Problems of Petroleum Migration. Am. Assoc. Pet. Geol., Stud. Geol., 10: 121-167. Toth, J. and Corbet, T., 1986. Post-Paleocene evolution of regional groundwater flow-systems and their relation to petroleum accumulations, Taber area, Southern Alberta, Canada. Bull. Can. Pet. Geol., 34(3): 339-363. Toth, J. and Otto, C. J. 1990. Hydrogeology and oil deposits at Pechelbronn-Soultz, Upper Rhine graben: Ramifications for exploration in Intermontane Basins. Int. Symp. on Intermontane Basins: Geology and Resources, Chiang Mai, Thailand, 30 January-2 February, 1989. Vail, J.R., 1972. Jebel Marra, a dormant volcano in Darfur Province Western Sudan. Bull. Volcanol., 19: 1-14. Vail, J.R., 1974. Distribution of Nubian Sandstone Formation in Sudan and vicinity. Bull. Am. Assoc. Petr. Geol., 58(6): 10251036. Vail, J.R., 1976. Outline of the geochronology and tectonic units of the Basement complex of North East Africa. Proc. R. Soc., London, A, 350: 127-141. Vail, J.R., 1978. Outline of the geology and mineral deposits of the Democratic Republic of Sudan and adjacent areas. Overseas Geol. Miner. Resour., 49:66 pp. Vail, J.R., 1982. Distribution and tectonic setting of post-kinematic igneous complexes in the Red Sea Hills of Sudan and the Arabian-Nubian Shield. Pre-Cambrian Res., 16 (Abstracts A41). Verweij, J.M. 1993. Hydrocarbon Migration Systems Analysis. Elsevier, Amsterdam, 276 pp. Weissbord, T., 1970. Nubian Sandstone, discussion. Bull. Am. Assoc. Petr. Geol., 54(3): 526-529. Whiteman, A.J., 1971. The Geology of the Sudan Republic. Clarendon Press, London. Wendorf, E, and Schild, R., 1976. Prehistory of the Nile Valley. Academic Press, New York, N.Y., 404 pp. Willcocks, N., 1904. The Nile in 1904. E. and F.N. Spon, London, 225 pp. Williams, M.A.J., 1966. Age of alluvial clays in the western Gezira. Nature, 211 : 270-271. Williams, A.J. and Adamson, D.A., 1973. The physiography of the central Sudan. Geogr. J., 139: 498-508. Williams, M.A.J. and Adamson, D.A., 1974. Late Quaternary desiccation along the White Nile. Nature, 248: 584-588. Williams, M.A.J. and Adamson, D.A., 1980. Late Quaternary depositional history of the Blue and White rivers in central Sudan. In: M.A.J. Williams and H. Faure (Editors), The Sahara and the Nile. Balkema, Rotterdam, pp. 281-304. Williams, M.A.J. and Williams, E, 1980. Evolution of Nile Basin. In: M.A.J. Williams and H. Faure (Editors), The Sahara and the Nile. Balkema, Rotterdam, pp. 207-224. Williams, M.A.J., Adamson, D.A. and Abdula. H.H., 1982. Land forms and soils of the Gezira: A Quaternary legacy of the Blue and White Nile rivers. In: M.A.J. Williams and D.A. Adamson (Editors), A Land Between Two Niles. Balkema, Rotterdam, pp. 111-142. Yuretich, R.E and Cerling, T.E., 1983. Hydrogeochemistry of Lake Turkana, Kenya: Mass balance and mineral reactions in an alkaline lake. Geochim. Cosmochim. Acta, 47: 10991109.

Chapter 7

The Niger Delta Basin

T.J.A. REIJERS, S.W. PETTERS and C.S. NWAJIDE

INTRODUCTION The Niger Delta Basin occupies the Gulf of Guinea continental margin in equatorial West Africa, between lats. 3~ and 6 ~ N and longs, 5 ~ and 8~ E. It ranks among the world's most prolific petroleum-producing Tertiary deltas, comparable to the Alaska North Slope, the Mississippi, the Orionoco, and the Mahakam (Indonesia) deltas (Fig. 1). These Tertiary deltaic hydrocarbon provinces account for about 5% of the world's oil and gas reserves and about 2.5% of the basin areas

1 MACKENZIE DELTA

(Alaska north slope) 2

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on earth. The sedimentary fills in Tertiary deltas are entirely clastic, and are supplied by large continental drainage systems which have constructed arcuate and "bird's foot" deltaic wedges that prograde from continental onto oceanic crusts along both divergent and convergent margins. Major Tertiary deltas in most parts of the world have constructed circular sedimentary basins in which the ratio of sedimentary volume to surface area is very high. Thus, the Niger Delta, with a total area of about 75,000 km 2 and a clastic fill up to 12,000 m thick, is the largest in Africa (Figs. 1 and 2).

9

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MAJOR DRAINAGE

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PRODUCING DELTAS

BASINS

A

NON-PRODUCINGDELTAS

Fig. 1. World's major petroleum-producing deltas. African Basins. Sedimentary Basins of the World, 3 edited by R.C. Selley (Series Editor: K.J. Hsti), pp. 151-172.

9 1997 Elsevier Science B.V., Amsterdam. All rights reserved.

152

T.J.A. REIJERS, S.W. PETTERS and C.S. NWAJIDE

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Fig. 2. Regional setting of Niger Delta Basin (after Emery et al., 1975; Dingle, 1982).

As a sedimentary basin, the Niger Delta encompasses a region that is much larger than the geographical extent of the modern delta constructed by the Niger-Benue drainage systems. It embraces other deltas "which are not members of the Niger system" (Allen, 1970), notably the Cross River Delta (Evamy et al., 1978), and extends into the continental margins of neighbouring Cameroon and Equatorial Guinea, which therefore "have portions of the Niger Delta that contain hydrocarbons" (Clifford, 1986). Also, the Niger Delta sedimentary wedge contains major submarine parts that constitute a complex continental margin which projects pronouncedly into the Gulf of Guinea so that it overlaps the outlines of Africa and South America when both

continents are fitted together along the 500-fathom bathymetric contours of Bullard et al. (1965) (Fig. 3, inset) The tectonic setting and geologic evolution of the Niger Delta Basin transcend and pre-date the post-Eocene regressive clastic wedge that is conventionally ascribed to the delta (Frankl and Cordry, 1967; Short and Stauble, 1967; Weber and Daukoru, 1975), hence the entire megastructure within which the Niger Delta lies is reviewed in this paper since that is the basin. This megastructure is the southern Benue Trough. Apart from the fact that Cretaceous strata are exposed along the northern margins of the Niger Delta, and have been intercepted on its eastern and western flanks, the probable occurrence of 5-6

THE NIGER DELTA BASIN

153 oceanic fracture zones (Emery et al., 1975; Delteil et al., 1974), initiated the Benue Trough, and also later controlled the location of the main axis of subsidence of the Niger Delta. The Chain Fracture Zone coincides with the Benin Hinge Line of the western Niger Delta (Fig. 4), whereas the eastern delta frame, generally referred to as the Calabar Flank (Reyment, 1965), is more complicated, with NW-SE-tending structures such as the Ikang Trough, the Ituk High, and the Calabar Hinge Line. Sinistral transcurrent shearing along the fracture zones caused deformation in the Benue Trough and modified the Gulf of Guinea continental margin from the simple pull-apart basement structures with half-grabens that underlie the West African continental margins north and south of the Gulf of Guinea. In the Benue Trough, sedimentary infillings lie in Cretaceous subbasins such as the Gongola, the Yola, the Abakaliki, the Anambra, and the Afikpo subbasins. The Niger Delta can therefore be considered as the youngest sub-

km of Jurassic-Lower Cretaceous strata beneath the delta (Dingle, 1982) (Fig. 2), and the location of the Niger Delta on the oceanward extension of the Benue Trough have regional tectonic and stratigraphic implications for the delta.

TECTONIC SETTING The Niger Delta Basin occupies the coastal and oceanward part of a much larger and older tectonic feature, the Benue Trough. The Benue Trough is a NE-SW folded rift basin (Figs. 2 and 4) that runs diagonally across Nigeria. It formed simultaneously with the opening of the Gulf of Guinea and the Equatorial Atlantic in Aptian-Albian times, when the equatorial part of Africa and South America began to separate (Benkhelil, 1987). Taphrogenic subsidence along fundamental transform faults which had cut through the lithosphere and are the landward continuations of the Chain and Charcot

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T.J.A. REIJERS, S.W. PETTERS and C.S. NWAJIDE

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basin in the Benue Trough, a region whose stratigraphic and paleogeographic evolution have been controlled by southward shifting deltaic depocentres (Fig. 5) (Benkhelil, 1989; Petters 1985); westward post-deformational displacement of depocentres, and northward directed marine transgressions (Murat, 1972; Petters, 1981).

ANTECEDENT DELTASIN THE BENUETROUGH The basal lithic fill in the Benue Trough, best exposed in the northern parts (Fig. 5A), are the Lower Cretaceous alluvial fan, braided river, lacustrine and deltaic clastics of the Bima Sandstone (Allix and Popoff, 1983) which also extend into the central Benue Trough (Nwajide, 1990). In the southern part and in the Niger Delta, basal continental clastics are not well known, except along the basin margins where Lower Cretaceous fluviatile deposits are exposed. Deposits of the earliest marine transgression, the Asu River Group of Middle Albian age, are however, more widespread (Fig. 5A), and following

a mild mid-Cenomanian regressive pulse with local deltaic accumulations (Awe and Keana Formations), there was a more regionally extensive transgression in the Late Cenomanian-Early Turonian (Fig. 5B). The deltaic Makurdi and Agala Sandstones marked Late Turonian regression, followed by another major transgression in the Coniacian. Although restricted Santonian marine beds have been encountered in the southern Benue Trough (Petters et al., 1983). The Santonian-Early Campanian was marked by regional folding in the Benue Trough, after which the Abakaliki trough was uplifted and the axis of deposition displaced to the Anambra Basin and the Afikpo Syncline (Fig. 6), where deltaic sedimentation became permanently established in the southern Benue Trough, leading to the present-day Niger Delta. In order to fully understand the evolution of the Niger Delta Basin, it is appropriate first to examine closely its immediate precursor, the Anambra Basin (Fig. 7), particularly as both basins are stratigraphically connected (Fig. 8). Deposits of earlier marine cycles are exposed on the Calabar Flank of the Niger Delta and in the Anambra Basin, whereas

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those of Tertiary cycles form the bulk of the succession which are encountered in oil wells in the Niger Delta. Deltaic sedimentation in the A n a m b r a Basin

In the Anambra Basin the deltaic complexes, over 2500 m thick, are divisible lithostratigraphically into three lithofacies (Agagu et al., 1985), namely: the Agbani (Coniacian), the Owelli (Campanian: Fig. 5D), and the Mamu-Ajali Sandstones (Figs. 5D and 7). Each of the deltaic packets comprises fluvial, deltaic plain, and pro-delta lithofacies, of which the Enugu Shale and the Nkporo Shale represent the brackish marsh and fossiliferous pro-deltaic facies respectively of the late Campanian-Early Maastrichtian Nkporo depositional cycle. Another depositional phase, the Sokoto cycle (Petters, 1979), followed in the Late Paleocene and culminated in the deposition of the Imo Shale in the Anambra Basin and the Kalambaina Formation in the Sokoto Embayment (Fig. 5E). Two major transgressions therefore occurred during the Late Campanian and the Paleocene in the Anam-

bra Basin (Fig. 10), and by Late Eocene times, the funnel-shaped, tide-influenced shelf sea in the southern Benue Trough (Fig. 5D) gradually changed into a wave-influenced, high-energy shelf sea off a curvilinear coastline where lobate deltas formed. The deposits of the Late Eocene regression which marked the beginning of the Niger Delta progradation constitute the outcropping Ameki "group" (Fig. 10) which includes continental facies (Nanka Formation) and paralic as well as pro-deltaic facies. The subsurface equivalents of the Imo Shale and the Ameki Group in the Niger Delta are the Akata and the Agbada Formations respectively (Fig. 10). The critical factors that governed deltaic sedimentation and the paleDgeographic evolution of the Anambra Basin are its shape, the proximity of sediment source areas, regular incursions of the sea during the CampanianMaastrichtian, and the paleocirculation pattern. The paleomorphology of the coastline was also significant. As can be deduced from Figs. 5C-E there was a gradual straightening out of the coastline over time. Like the modern Niger Delta, the shape of the coastline, together with the shelf bathymetry and wave dy-

THE NIGER DELTA BASIN

157

Fig. 7. Geological map of Southern Benue Trough. namics, driven by the prevailing wind direction, determined the nature of the longshore drift. Assuming there was, throughout the Late Cretaceous and Early Tertiary, a precursor of the Guinea current which today prevails along the shelf of the Niger Delta, and that the Anambra Basin was located within the tropical belt (Berggren and Hollister, 1974) with mon-

soonal winds approaching from the southwest, then the only fundamental difference between the Anambra Basin and the present-day Niger Delta was the shape of the coastline. The funnel-shaped Anambra shelf sea probably generated two drift cells as opposed to the four independent cells of the Recent Guinea current. In the Anambra Basin there was a

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THE NIGER DELTA BASIN southwest to northeast-directed drift cell, and a southeast to northwest directed cell. Unavoidably, therefore, the converging flows, both pointing towards the apex of the embayment, must have led to constriction, and caused sedimentation dominated by storm waves and tides, especially during periods when the paleocirculation was reinforced by transgressive tendencies. During pronounced regression, however, another flow factor, that of riverine input from the protoNiger River, might have been stronger. Thus, the stratigraphic succession in the Anambra Basin can be interpreted not only in terms of transgressiveregressive cycles as evident above, but within the context of an evolving coastline configuration, and variable sediment input.

The Campanian Formations of the Nkporo Group (Fig. 10) reflect a funnel-shaped shallow marine shelf setting that graded into channelled low-energy marshes. The concave inward shape of the coastline gave scope for gradual filling with the small amounts of sediments brought in by short rivers draining the sediment-dispersal centres, that were carried along by northeastward converging longshore drifts. Behind poorly developed foreshores and shorefaces, extensive coastal swamps could be expected. In many respects this configuration was comparable with that currently prevailing on Nigeria's southwest coast. Within the swamps, fluviatile point bars were formed that are preserved as the Owelli Formation. The shallow open marine shelf sea was alternatively storm- and tide-dominate.

The Maastrichtian Sediments of the coal-bearing Mamu and the tidally influenced Ajali Formations (Fig. 10) accumulated during this epoch when the Nkporo cycle tended towards an overall regression with associated progradation. The Mamu coastal plain was drier than the earlier swamps. Muddy shorefaces separated it from mud-dominated shelfs. The Ajali Sandstone marks the height of the regression when the coastline was still concave, albeit less "embayment-like". Two converging littoral drift cells governed sedimentation and are reflected in the characteristic unidirectional tidal sand waves in the Ajali sandstone.

The Paleocene The Nsukka Formation marks the onset of the Sokoto transgression and documents a return to paludal conditions. Sedimentation was mainly by fluvial input. However, it was punctuated by marine incursions reflecting continued convergence of the two littoral drift cells controlling shoreface sedimen-

159 tation. The Imo shales reflect shallow-marine shelf conditions in which foreshore and shoreface sands formed occasionally. The shales contain a significant amount of organic matter which are potential source rock for hydrocarbons in the eastern part of the Niger Delta and in the Anambra basin.

The Eocene The progradational Nanka Formation marks the return to regressive conditions. It offers an excellent opportunity to study tidal deposits (Fig. 9A). Frequently exposed, strongly asymmetric sandwaves suggest the predominance of one total current direction over the other and the operation of reverse currents is suggested by the bundling of laminae separated from each other by mud drapes, reflecting neap tides.

Sequence stratigraphy of the Anambra Basin Although the Anambra Basin has not yet been extensively investigated with the aim of recognizing sequence stratigraphic features, the repeated allocyclic incursions of the sea however, resulted in characteristic basin-wide genetic sequences (Bush, 1971; Galloway, 1989), or parasequence sets (Van Wagoner et al., 1988). Two allocyclic events have been recognized in the Anambra Basin that encompass large time intervals (Fig. 10), in contrast to the better studied Niger Delta with a more complete log record, where at least eleven events, usually involving considerably smaller time increments, are reflected by a cyclic pattern of transgressiveregressive lithologic units (Reijers et al; in preparation). Typically, such lithologic units contain fines with characteristic marine faunal assemblages, vertically followed by a variety of coarser clastic parasequences, usually of paralic affinity (Figs. 10 and 11) and often of an autocyclic nature. Some of the fines in the Anambra Basin and most of the shales in the Niger Delta have been correlated with major global flooding events (Petters, 1983). The allocyclic events recognized in the Anambra Basin and the Niger Delta have, however, lumped together numerous world-wide transgressive-regressive events reflected by global maximum flooding surfaces (Haq et al., 1988). In the Anambra Basin the sedimentary fill, summarized above, is bounded at the base by the Santonian (76 m.y.) and at the top by the Eocene (38.6 m.y.) unconformities (Fig. 10). It reflects a second order cycle composed of two transgressiveregressive parasequence pairs, reflecting relative sealevel fluctuations. The best exposure in the Anambra Basin studied so far and aimed at identifying sequence stratigraphic features, is located along the Enugu-Port Harcourt express-way, at Leru, where

160

T.J.A. REIJERS, S.W. P E T T E R S and C.S. N W A J I D E

Fig. 9A. Flood tidal sandwaves in the Eocene Nanka Formation exposed 18 km northeast of Onitsha in the Anambra Basin. Note reactivation surfaces (Sr), tidal bundles (Bt) and wave ripple laminated master bedding (Bm) formed by coalescent mud drapes on sigmoidal foresets. Note also the thinning upward trend in the bedding unit.

Fig. 9B. Part of the Campanian succession (Nkporo Formation) in the Anambra Basin exposed on the scarp slope of the Enugu Cuesta at Leru, near Okigwe. Note clean beach sands (Bs) overlain by an incised valley fill (ivf) bounded base and topped by a scoured surface (Ss) and ravine surface (Sr) respectively. Sb = black shale; Bt = tidal bundles.

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162

T.J.A. REIJERS, S.W. PETTERS and C.S. NWAJIDE

Fig. 11. Recent depositional environments in the Niger Delta.

the road breaches the Enugu Escarpment (Fig. 7). Here, a type 1 sequence boundary, associated with significant erosion, divides the section into lower and upper parasequence sets. An incised valley fill underlies the sequence boundary. Other erosional sandstone bases in the section are 4th or 5th order breaks. Two maximum flooding surfaces (MFS) have

been inferred; the lower may be associated with the Early Maastrichtian (73.5 m.y.) condensed sequence indicated by Haq et al. (1988). The upper MFS is positioned close to the base of the Mamu Formation. A ravinement surface (transgressive surface of erosion) terminates the incised valley fill and underlies a transgressive systems tract.

THE NIGER DELTA BASIN

163

Table 1 Lithofacies characteristics and sedimentary environments of the Recent Niger Delta and lithofacies characteristics Environment

Environmental characteristics

Lithofacies

Lower floodplain

Strong currents in channels, meander migration, periodic flooding of topstratum levees and backswamps, abundant plant growth

Channels and point-bars, mottles, backswamps; cut-off channels similar to backswamps

Mangrove swamp

Strong reversing tidal currents, tidal flats

Channels and point-bars: mainly layered cross-stratified f. to v.c. sand and organic-rich silty clay; abundant drifted plant debris; inter-creek flats and inter-swamp deltas

Beach

Strong wave attack on active beaches, shore currents diverging from delta tip; soil formation and plant growth on beach ridges

Delta tip: mainly evenly laminated clean f. to m. sand

River mouth bar

Very strong wave action and reversing tidal currents; longshore current; energy conditions decrease inland and seaward from bar crest with increase in depth

Crests: mainly clean, v.f. to m. sand with even lamination, cross-stratification or cut-and-fill. Bar flanks: layered v.f. sand, clean v.c. silt and clayey silt; drifted plant remains sometimes in thick layers

Delta-front platform

Strongly to moderate wave and tidal current action; longshore currents; Guinea currents; rip currents. Energy conditions decrease from shoreface to outer edge

Delta tip: on inner platform coarse v.f. sand and v.c. silt with even laminations; on outer platform layered v.f. sand, v.c. silt, clayey silt and silty clay with plant debris

Pro-delta shape

Moderate to weak wave and tidal current action; Guinea current

V.f. sand, v.c. silt, clayey silt and silty clay; coarser layers with even lamination, cross-stratification, plant debris and mica flakes; common to abundant mottles. Delta flanks: layered v.c. silt, clayey silt and silty clay in shallower parts; uniform fine clayey silty clays in deeper areas; abundant mottles

Open shelf

Weak wave and tidal current action; deep ocean currents of unknown strength flowing northward over shelf edge

Delta tip: layered v.c. silt, clayey silt and silty clay; mainly uniform fine clayey silt and silty clay; abundant mottles and pelagic foraminifera. Delta flanks: mainly uniform fine silty clays; abundant mottles and pelagic foraminifera

Nondepositional

Weak wave and tidal current action; strong to moderate action inshore; no or very slow deposition of suspended fines; abundant benthos, organic debris concentrated

Mainly mottled v.f. to v.c. quartz sands largely out of equilibrium with prevailing current conditions; shell debris; glauconite, foraminifera and clay-silt increase from shallow to deep water; partly Late Pleistocene in age. Deposits interpreted as of strandplain origin

Delta top Benin facies

Delta front Agbada facies

Pro-delta Akata facies

From Weber and Daukoru (1975) and Whiteman (1982).

CENOZOIC NIGER DELTA

The m o d e r n delta

Recent depositional environments and sedimentary facies in the Niger Delta described by Allen (1965, 1970) have furnished a lithogenetic model that relates facies variations of modem high energy, wave-dominated, constructional, arcuate-lobate tropical deltas to depositional process (Fig. 11, Table 1). Apart from enhancing sedimentological, stratigraphic, and paleoenvironmental interpretations of the thick Cenozoic succession in the Niger Delta (Weber, 1971; Weber and Daukoru, 1975), the crite-

ria established for the Recent Niger Delta have been successfully applied to ancient deltas in other parts of the world, such as the Early Devonian Bokkeveld Group in the Cape Province of South Africa, where a similar prograding wave-dominated deltaic system was inferred (Hiller and Theron, 1988). It is therefore appropriate to examine the physiographic and depositional processes in the modem Niger Delta before considering the lithostratigraphy and stratigraphic development of the thick Cenozoic deltaic fill. As shown in Fig. 11, the Niger Delta displays a concentric arrangement of terrestrial and transitional depositional environments. These environments can be broadly categorized into three distinct facies belts

164

T.J.A. REVERS, S.W. PETTERS and C.S. NWAJIDE

THE NIGER DELTA BASIN

165

(Table 1). These are the continental delta top facies; the paralic delta front facies; and pro-delta facies. Fluvial processes control sedimentation in the lower floodplain of the delta top environment, whereas from the mangrove swamps coastward, tidal influences prevail. Semidiurnal tides approach the

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Niger Delta from the southwest with a range that increases from 1 m at Lagos in the west to 2.8 m, over mouth bars at the Calabar River (Allen, 1970). Tidal currents in the mangrove swamps are strong, and also sweep the fiver front platform. But the dominant process which constructs the beachridge and barrier complexes along the Niger Delta coast are a combination of large, perennial swell waves which approach from the southwest, and the vigorous longshore drift which the waves generate. Offshore, the warm Guinea Current prevails, operating as four independent cells, under the influence of the convex, seaward coastline of the Niger Delta and the predominantly NE-directed trade winds. Consequently, nowadays a west-east directed drift cell of the Guinea Current flows from Cape Palmas far to the west toward the northwestern flank of the delta; a north-westward drift cell flows from Akassa Point in the Niger Delta to the NW flank of the delta; whereas a west--eastward drift cell prevails from Akassa Point to the Cross River Estuary. Along the eastern shelf, there is a northward drift cell from the Congo River to the Rio Del Rey River. Below the thermocline, currents of uncertain strength flow to the north or northwest. From the shoreline to shelf break at the 90 m bathymetric contour, 50100 km offshore, the submarine topography of the Niger Delta is almost fiat. In front of the delta, the continental shelf is wider and shallower than further south. Allen (1965) showed the presence of three terraces on the Niger Delta shelf, which he attributed to the drowning of barrier-beach or barrier island complexes that were formed during stillstands of the Holocene transgression. These submarine terraces are crossed by several shallow-channels, believed to be former distributaries of the Niger Delta, that were incised during a lowstand of sea-level. Apart from these relict channels, major submarine canyons notch the shelf and continental slope off the Niger Delta (Fig. 3). These include the Qua Iboe and Calabar Canyons in the east (Houbolt, 1973), the Niger Canyons off the central bulge of the delta (U.S. Naval Oceanographic Office, 1965), and the Avon and Mahin Canyons on the western re-entrant of the delta (Burke, 1972). Turbidity currents in submarine canyons off the Niger Delta have deposited deepsea fans on the continental rise of the delta. The continental slope off the delta comprises a gentle upper slope, and a hilly, steep lower slope known as the Nigeria Escarpment. A belt of mud diapirs occurs along the seaward side of the Nigeria Escarpment (Delteil et al., 1974). As shown below, the aforementioned depositional processes, namely, fluvial, coastal, and marine, including turbidity currents, coupled with the rises and falls of sea-level have determined the stratigraphic fill of the Cenozoic Niger Delta.

166

Lithostratigraphy The lithostratigraphy of the Cenozoic Niger Delta is a direct product of the various depositional environments outlined above. Ever since on-going deltaic progradation ensued in the Early Tertiary, these environments and their characteristic lithofacies have prevailed. Well sections generally display vertical subdivisions (Short and Stauble, 1967) (Fig. 12), in which an upper delta top lithofacies, the Benin Formation, consists of massive continental sands and gravels; and is underlain gradationally by the delta front paralic lithofacies, the Agbada Formation, comprising mostly sands with minor shales in the upper part, and an alternation of sands and shale in equal proportion in the lower part. Pro-delta marine shales belonging to the Akata Formation occur lower in the section, with sandstone units of deep-sea fan origin. As shown in Figs. 12 and 13, Niger Delta lithostratigraphic units are strongly diachronous having began to accumulate since deltaic progradation commenced in the Early Tertiary; in which case the Akata Formation ranges from Paleocene to Recent; the Agbada Formation from Eocene to Recent; and the Benin Formation from Oligocene to Recent. Along the northern perimeters of the Niger Delta where the proximal parts of these lithostratigraphic units are exposed (Figs. 7, 10 and 11), different formation names have been assigned, namely: Imo Shale (Akata), Ameki (Agbada), and Ogwashi-Asaba (upper Agbada facies). As evident in Fig. 13, large clay fills of ancient submarine canyons, containing turbidites are common in the eastern and western parts of the Niger Delta succession (Burke, 1972; Orife and Avbovbo, 1982; Petters, 1984) The submarine canyons that were subsequently filled by these clays were entrenched mainly during Late Tertiary lowstands of sea-level (Fig. 13).

Evolution of depobelts and sequences Stacher (1995) showed that the Niger Delta sequence consists of a series of discrete depocentres or depobelts (Fig. 14) which were the main belts of deposition of the Agbada Formation that succeeded each other progressively as the delta shifted its loci downdip through time. Major structure-building growth faults determined the location of each depobelt. As evident in Fig. 14, the entire sedimentary wedge in the Niger Delta were laid down sequentially in five major depobelts, each 30-60 km wide, with the oldest depobelt lying furthest inland and the youngest located offshore. Sedimentation in the depobelts was a function of the rate of deposition and the rate of subsidence with syn-sedimentary growth faults upsetting the delicate balance (Evamy

T.J.A. REIJERS, S.W. PETTERS and C.S. NWAJIDE et al., 1978). The combined effect of continued sediment supply, syn-sedimentary faulting, and relative sea-level fluctuations gave rise to the continuous vertical cyclic stacking of proximal fluvio-marine interlaminated silt, sand, and clay; usually followed by various types of lower to upper shoreface sands and coastal plain deposits; each vertically terminated by the next transgressive event (Weber, 1971). Active growth faults control the development of each cycle (Fig. 11B). The sedimentation mechanism in the depobelts was termed the "escalator regression" model by Knox and Omatsola (1989). On the downthrown side of the major structure-building growth faults, the paralic facies of the Agbada Formation accumulated in the depobelts, when the ratio of sedimentsupply versus subsidence was approximately equal to or less than 1. This corresponded to the active phase of subsidence in a particular depobelt. Thus, syn-sedimentary faults trapped dispersing sediments into a local depocentre at the downthrown side of the fault, where sediment accumulation ensured continued equilibrium and on-going deposition along the shoreface (Fig. 11B). When subsidence rates decreased, the rate of addition of accommodation space reduced and alluvial continental sands of the Benin Formation rapidly advanced across the depobelt to maintain the base-level of sedimentation which had been transferred in the down-dip direction into the adjacent depobelt or depocentre that was created by a younger syn-sedimentary fault system. Again, shoreface lithofacies mark the base in the down-dip direction of the adjacent depobelt, thus rendering the base of the new depobelt equivalent to the continental facies of the older up dip depobelt, in which case the "time" lines in the Niger Delta cross different lithofacies, from marine into paralic, and into continental (Fig. 12), hence the delta lithostratigraphic units are strongly diachronous and difficult to subdivide and correlate using marine biostratigraphic criteria. Sequence stratigraphy is therefore potentially applicable to the Niger Delta, in that the fundamental building blocks of the Niger Delta succession are well defined cyclic offlapping parasequence sets (Figs. 11B and 12). Each parasequence set consists of a marine clay which represents the marine flooding surface, that change upward into proximal fluviomarine interlaminated silt, sand and clay; usually followed by various types of lower to upper shoreface sands; and coastal plain, continental deposits. The above parasequence sets are repeated many times and stacked vertically into the thick highstand systems tracts that are found in the various depobelts (Figs. 12 and 14). During lowstands, regressive parasequence sets followed with progradation sometimes occurring beyond the shelf edge, leading to the accumulation of sediments in the deeper

THE NIGER DELTA BASIN

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part of the delta, in which lowstand systems tract fan systems commonly developed. A comprehensive delta-wide sequence stratigraphic subdivision of the Niger Delta basin is being undertaken by Reijers et al. (in preparation). Occasionally, delta-wide marine transgressions (Petters, 1983) that coincided with eustatic sea-level changes (Haq et al., 1988), interrupted deltaic progradation and caused the Niger Delta paleoshoreline to retreat landward over a distance of 70-100 km. Planktonic foraminifera from some of the major regional marker shales that ac-

cumulated during the delta-wide marine transgressions, suggest several major short-lived transgressive pulses within the Late Paleocene, Early Eocene, Late Eocene, Late Oligocene-Early Miocene, Late Early Miocene, Middle Miocene, Late Miocene, and Pliocene. Foraminiferal and palynological ages have enabled detailed paleogeographic reconstructions of the Cenozoic Niger Delta (Fig. 15). During the Paleocene-earliest Eocene, marine shales accumulated over much of the Niger Delta, with paralic sediments being restricted to what was the incipient

168

T.J.A. REIJERS, S.W. PETTERS and C.S. NWAJIDE

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Cross River Delta (Evamy et al., 1978). Deposition of coarse material of the Agbada Formation took place in what is now known as the North-

ern Delta depobelt in the Early-Middle Eocene (Fig. 14), while fines were simultaneously deposited further offshore as the marine Akata Shale For-

THE NIGER DELTA BASIN

169

Fig. 15. Paleodrainage trends in Niger Delta (after Ejedawe, 1981). mation. Syn-sedimentary growth faulting developed close to the coastline, which trapped sands and fines and enabled the delta to prograde. During the Oligocene and Earliest Miocene, the riverine systems that fed the emerging Niger Delta debouched at a coastline along which longshore currents began to play an active role in sediment distribution and a thick sequence of sandy sediments accumulated in the Greater Ughelli depobelt. Evamy et al (1978) relate this localized deposition to pronounced subsidence and a relatively slow advance of the delta front. A paleohigh, the prolongation of the Abakaliki uplift, was active during the Oligocene-Early Miocene (Evamy et al., 1978); and separated the emerging Niger Delta into segments each characterized by its own palaeodrainage pattern (Fig. 15). Sediments were deposited during the Middle-Late Miocene in the Central Swamp and the northern zones of the present-day Coastal Swamp depobelts. During the Late Miocene to Pliocene, the delta prograded steadily and the present-day youngest depocentres (Coastal Swamp and Offshore depobelts) were formed. From Late Miocene onward a large depocentre developed as the Cross River Delta, including its eastern part, the Rio Del Rey Basin in Cameroon (Fig. 16).

STRUCTURE

The influence of basement tectonics on the structural evolution of the Niger Delta was largely limited to movements along the Equatorial Atlantic Oceanic fracture zones which extend beneath the delta and determined the initial locus into which the protoNiger built its delta. As the delta advanced onto oceanic crust, repeated subsidence of the oceanic basement (Fig. 8) created more space for the thick sedimentary pile of the prograding Cenozoic Niger Delta (Evamy et al., 1978; Hospers, 1971). On the eastern flank of the Niger Delta Basin, the wedgeshaped geometry of Late Tertiary clastic facies suggests the uplift and impact of the Cameroon Volcanic Line as the dominant sediment source area into the Cross River Delta (Rio Del Rey Basin) (Seme Obomo et al., 1994). The paleogeographic evolution of the vicinity of the Cameroon Volcanic Line (Fig. 16) reveals the dominance of sedimentation in the Cross River Delta from Pliocene onwards, unlike the neighbouring Doula Basin which was the main Oligocene-Miocene deltaic depocentre (Regnoult, 1986). Growth faults triggered by penecontemporaneous deformation of deltaic sediments, are the common structures in the Niger Delta (Merki, 1972; Evamy et al., 1978). They are generated by rapid

170

T.J.A. REIJERS, S.W. PETTERS and C.S. NWAJIDE

Fig. 16. Paleogeography of eastern Niger Delta Basin showing influence of the Cameroon Volcanic Line. sedimentation and gravitational instability during the accumulation of the paralic Agbada deposits and the continental Benin sands over the mobile, undercompacted Akata pro-delta shales. Lateral flowage and extrusion of the Akata pro-delta shales during growth faulting and related extension, also account for the diapiric structures on the continental slope of the Niger Delta in front of the advancing depocentre of paralic sediments (Figs. 3 and 16). Growth faults in the Niger Delta are restricted to the paralic Agbada Formation. They comprise (Fig. 14) the major structure-building faults, some of which bound the depobelts; steep, parallel crestal faults which cut the rollover structures; and antithetic faults. Associated with the structure building faults are the rollover anticlinal structures (Fig. 14) Growth faults

and rollover structures are the dominant hydrocarbon traps in the Niger Delta.

PETROLEUM GEOLOGY The Niger Delta Basin holds enormous petroleum reserves, estimated at about 30 billion barrels of oil and 260 trillion cubic feet of natural gas, ranking the delta seventh in world production, with a current average production of about 1.8 million bbl of oil/day. A few giant oil and condensate fields, with reserves exceeding 500 million bbl occur in the Niger Delta (Fig. 14), for example Edop, Jones' Creek, Oso, Imo River, Nembe, Forcados Yorki, Cawthorne Channel, Meren, Delta South, and Okan fields. About 937 mil-

THE NIGER DELTA BASIN lion barrels of oil and condensate have been discovered in the Rio Del Rey sector in Cameroon, and 45 million barrels around Equatorial Guinea. Oil and gas reserves in the Niger Delta Basin are concentrated in sandstone reservoirs throughout the paralic Agbada Formation, usually trapped in rollover anticlines associated with growth faults (Fig. 14). Additional to the conventional growth fault-related traps are the stratigraphic traps related to paleochannel fills, regional sand pinch-outs, and truncation (Orife and Arbovbo, 1982; Stacher, 1995). Gross reservoir properties are a function of sand/shale ratios and the sealing potential of faults, while the transgressive marine shales form important regional top seals, with faults often providing lateral seals (Fig. 14). Because of sand/ shale alternations and the repetitive nature of the traps, most oil fields in the Niger Delta have multiple, stacked reservoirs, with oil column heights ranging from 15 to 50 m. Mature Eocene and Miocene shales of the Akata and Agbada Formations constitute the major source rocks (Ekweozor and Okoye, 1980; Ejedawe, 1981; Ejedawe and Okoh, 1981; Nwachukwu and Chukwura, 1984; Bustin, 1988). Niger Delta crudes originated mostly from land plant material, hence they are high in resins and waxes, with significant contributions of structureless organic matter from marine sources. The low sulphur contents of Nigerian crudes (below 0.4%) also confirm land plant source material. Overall, two types of crudes are found in the Niger Delta, a light, paraffinic, waxy crude with pour points of about 20 to 90~ and a naphthenic, nonwaxy medium crude with specific gravity of less than 26~ and pour points below - 13~

CONCLUSION

The Niger Delta Basin is divisible into a Cretaceous succession that is exposed in the Anambra Basin to the north, and a much thicker Tertiary fill which underlies the Recent Niger Delta. Deltaic sedimentation in both basins, located in the Southern Benue Trough, was controlled by global sea-level fluctuations which produced two transgressive-regressive pairs of parasequence sets in the Anambra Basin, whereas in the Tertiary Niger Delta southward migrating depobelts contain more parasequence sets in the strongly diachronous Agbada Formation. Another control was the shape of the coastline which changed from convex, through curvilinear, to concave, thus modifying the longshore drift cells from convergent to divergent. This resulted in a gradual change from strongly tidal Cretaceous sedimentation to river dominated deltaic sedimentation in the older parts of the Niger Delta, and to wave-dominated deltaic sedimentation in the Cenozoic. Basin fill was continuous but ir-

171 regular, being influenced by growth faulting that controlled subsidence and sedimentation rates, and whether proximal or distal sediment types of growing deltas predominated.

ACKNOWLEDGEMENTS

We acknowledge the permission of Shell Petroleum Development Company (SPDC) and their Joint Venture partners Nigerian National Petroleum Company (NNPC), Elf, and Agip as well as the Department of Petroleum Resources of the Ministry of Petroleum Resources, Nigeria, to publish this paper. We also thank professor Dr. Ken Hsti for inviting us to contribute to this volume.

REFERENCES

Agagu, O.K., Fayose, E.A. and Petters, S.W., 1985. Stratigraphy and sedimentation in the Senonian Anambra Basin of eastern Nigeria. Nigerian J. Min. Geol., 22: 25-36. Allen, J.R.L., 1965. Late Quaternary Niger Delta, and adjacent areas; sedimentary environments and lithofacies. AAPG. Bull., 49: 547-600. Allen, J.R.L., 1970. Sediments of the modern Niger Delta, a summary and review. In: J.P. Morgan and R.H. Shaver (Editors), Deltaic Sedimentation. Modern and Ancient. SEPM Spec. Publ., 15: 138-151. Allix, P. and Popoff, M., 1983. The Lower Cretaceous of the northeastern part of the Benue Trough (Nigeria): An example showing the close relationship between tectonics and sedimentation. Bull. Cent. Rech. Explor. Prod. Elf-Aquitaine, 7: 349359. Benkhelil, J., 1987. The origin and evolution of the Cretaceous Benue Trough (Nigeria). J. Afr. Earth Sci. 8:251-282. Benkhelil, J., 1989. Structural frame and deformations in the Benue Trough.Bull. Cent. Rech. Explor. Prod. Elf-Aquitaine, 7: 349-359. Bergggren, W.A. and Hollister., C.D. 1974. Paleogeography, paleobiogeography and the history of circulation in the Atlantic ocean. In: W.W. Hay (Editor), Studies in Paleo-Oceanography. SEPM Spec. Publ., 20:126-185. Bullard, E.C., Everett, J.E. and Smith, A.G., 1965. The fit of the continents around the Atlantic. In: P.M.S. Blackett, E.C. Bullard, and S.K. Runcorn (Editors), Symp. on Continental Drift. R. Soc. Lond. Philos. Trans. Ser. A, 258: 41-51. Burke, K.C., 1972. Longshore drift, submarine canyon and submarine fans in development of Niger delta. AAPG. Bull., 56: 1975-1983. Bush, D.A. 1971. Genetic units in delta prospecting. AAPG Bull., 55:1137-1154. Bustin, R.M., 1988. Sedimentology and characteristics of dispersed organic matter in Tertiary Niger Delta: Origin of source rocks in a deltaic environment. AAPG Bull., 72: 277-298. Clifford, A.C., 1986. African oil m past, present, and future. In: M.T. Halbouty (Editor), Future Petroleum Provinces of the World. AAPG. Mem., 40: 339-373. Delteil, J.R., Valery, P., Montadert, L., Fondeur, C., Patriat, P. and Mascle, J., 1974. The continental margin in the northern part of the Gulf of Guinea. In: C.A. Burk and C.L. Drake (Editors), Geology of Continental Margins. Springer-Verlag, New York, N.Y., pp. 297-311.

172 Dingle, R.V., 1982. Continental margin subsidence: A comparison between the east and west coasts of Africa. In: R.A. Scruton (Editor), Dynamics of Passive Margins. Geodynamics Ser. 6, Am. Geophys. Union/G.S.A., Washington, D.C., 59 pp. Ejedawe, J.E., 1981. Patterns of incidence of oil reserves in Niger Delta basin. AAPG. Bull., 65: 1574-1585. Ejedawe, J.E. and Okoh, S.U., 1981. Prediction of optimal depths of petroleum occurrence in the Niger Delta Basin. Oil Gas J., 79:190-204. Ekweozor, C.W. and Okoye, N.V., 1980. Petroleum source bed evaluation of Tertiary Niger Delta. AAPG. Bull., 64: 12511259. Emery, K.O., Uchupi, E., Phillips, J., Bowin, C. and Mascle, J., 1975. Continental margin off western Africa: Angola to Sierra Leone. AAPG Bull., 59: 2209-2265. Evamy, B.D., Haremboure, J., Kamerling, P., Knaap, W.A., Molloy, EA. and Rowlands, P.H., 1978. Hydrocarbon habitat of Tertiary Niger Delta. AAPG Bull., 62: 1-39. Frankl, E.J. and Cordry, E.A., 1967. The Niger delta oil province recent developments onshore and offshore. 7th World Petroleum Congr., Mexico City Proc. IB, pp. 195-209. Galloway, W.E., 1989. Genetic stratigraphic sequences in basin analysis 1: Architecture and genesis of flooding-surface bounded depositional units.AAPG Bull. 73: 125-142. Haq, B.U., Hardenbol, J. and Vail, P.R., 1988. Mesozoic and Cenozoic chronostratigraphy and cycles of sea level changes, In: C.K. Wilgus, B.S. Hastings, H. Posamentier, J. van Wagoner, C.A. Ross and C.G. St. Kendall (Editors), Sea Level Changes: An Integrated Approach. SEPM Spec. Publ., 42:71108. Hiller, N. and Theron, J.N., 1988. Benthic communities in the South African Devonian. In: N.J. McMillan, A.E Embry and D.J. Glass (Editors), Devonian of the World, Can. Soc. Pet. Geol. 3: 229-242. Hospers, J., 1971. The geology of the Niger delta area. In: F.M. Delany (Editor), Continental Margin, Geol. E. Atlantique, SCOR Conference, Cambridge 1970. HMSQ, pp. 123-142. Houbolt, J.J.H.C., 1973. The deep-sea canyons in the Gulf of Guinea near Fernando Po. Verh. K. Ned. Geol. Mijnbouwk. Genoot., pp. 1- ! 8. Knox, G.J. and Omatsola, M.E., 1989. Development of the Cenozoic Niger Delta in terms of the escalator regression model. In: Coastal Lowlands. Geology and Geotechnology, Proc. K. Ned.. Geol. Mijnbouwk. Genoot., pp. 181-202. Merki, P., 1972. Structural geology of the Cenozoic Niger Delta. In: T.F.J. Dessauvagie and A.J. Whiteman (Editors): African Geology. Ibadan University Press, Ibadan, pp. 635-646. Murat, R.C. 1972. Stratigraphy and paleogeography of the Cretaceous and lower Tertiary in southern Nigeria. In: T.F.J. Dessauvagie and A.J. Whiteman (Editors), African Geology. Ibadan University Press, Ibadan, pp. 251-266. Nwachukwu, J.I. and Chukwura, P.I., 1986. Organic matter of Agbada Formation, Niger Delta, Nigeria. AAPG. Bull., 70: 48-55. Nwajide, C.S., 1990. Cretaceous sedimentation and paleogeography of the Central Benue Trough. In: C.O. Ofoegbu (Editor),

T.J.A. REIJERS, S.W. P E T T E R S and C.S. N W A J I D E The Benue Trough Structure and Evolution. Friedr. Vieweg and Sohn, Braunschweig/Wiesbaden, pp 19-38. Orife, J.M. and Avbovbo, A.A., 1982. Stratigraphic and unconformity traps in the Niger Delta. In: M.T. Halbouty (Editor), The Deliberate Search for the Subtle Trap. AAPG Mem., 32: 251-265. Petters, S.W., 1979. Maastrichtian-Paleocene cyclothems and paleoclimate in SE Lullemmneden Basin, West Africa. Newsl. Stratigr., 8:180-190. Petters, S.W., 1981. Paleoenvironments of the Gulf of Guinea. Oceanol. Acta Proc. 26th Int. Geol. Congr. on Geology of Continental Margins Symposium, Paris, July 7-17, 1980, pp. 81-85. Petters, S.W., 1983. Gulf of Guinea planktonic foraminiferal biochronology and geological history of the South Atlantic. J. Foraminiferal Res., 13: 32-59. Petters, S.W., 1984. An ancient submarine canyon in the Oligocene-Miocene of the western Niger Delta. Sedimentology, 31: 805-810. Petters, S.W., 1985. Foraminiferal biofacies in the Nigerian rift and continental margin deltas. In: M.N. Oti and G. Postma (Editors), Geology of Deltas. A.A. Balkema, Rotterdam, pp. 219-235. Petters, S.W., EI-Nakhal, H.A. and Cifelli, R., 1983. Costellagerina, a new Late Cretaceous globigerine foraminiferal genus. J. Foraminiferal Res., 13: 247-251. Regnoult, J.M., 1986. Synthese G6ologigue du Cameroon. Ministbre des Mines et de l'Energie, Yaounde, 300 pp. Reyment, R.A., 1965. Aspects of the Geology of Nigeria. Ibadan Unversity Press, Ibadan, 145 pp. Seme Obomo, R., Rosendahl., B.R., Loule, J.P., 1994. Seismic evidence of the Cameroon volcanic impact on the stratigraphy of the northern Douala Basin. 12~ Coll. Afr. de Micropaleontologie, 2~ Coll. de Stratigraphie et de Paleogeographie de l'Atlantique Sud, Angers, France, 16-20 Juillet ! 994, 13 i pp. Short, K.C. and Stauble, A.J., 1967. Outline of geology of Niger delta. AAPG Bull., 51:761-779. Stacher, P., 1995. Present understanding of the Niger Delta hydrocarbon habitat. In: M.N. Oti and G. Postma (Editors), Geology of Deltas. A.A. Balkema, Rotterdam, pp. 257-267. U.S. Naval Oceanographic Office, 1965. Oceanographic Atlas of the North Atlantic Ocean. Pub. No. 700, Washington D.C., 71 PP. Van Wagoner, J.C., Posamentier, H.W., Mitchum, R.M., Vail, P.R., Sarg, J.E, Loutit, T.S. and Hardenbol, J., 1988. An overview of the fundamentals of sequence stratigraphy and key definitions. In: C.K. Wilgus, B.S. Hastings, H. Posamentier, C.A. Ross and C.G. St.C. Kendall (Editors), Sea-Level Changes: An Integrated Approach. SEPM, Spec. Publ. 42: 39-46. Weber, K.J., 1971. Sedimentological aspects of oil fields in Niger delta. Geol. Mijnb., 50: 559-576. Weber, K.J. and Daukoru, E.M., 1975. Petroleum geology of the Niger Delta. Proc. 9th World Petr. Congr., 2:209-221. Whiteman, A.J., 1980. Nigeria: Its Petroleum Geology, Resources and Potential, Vol. !. Graham and Trotman, London, 166 pp.

Chapter 8

The West African Coastal Basins

M.A. ALA AND R.C. SELLEY

INTRODUCTION

From the Niger delta south to the Cape of Good Hope a series of basins occurs along the west African coast (Lehner and De Ruiter, 1977). On land they are separated from one another by Precambrian basement, but when traced offshore they merge into a single wedge of sediment that thins westward into the South Atlantic Ocean. Each basin is related to a major river system that drains the adjacent hinterland of continental Africa. From north to south the rivers and their sedimentary basins are as follows: The Rio del Rey basin of Equatorial Guinea and Cameroon occupies the eastern edge of the Niger river system. The Sanaga River drains into the Douala basin of Cameroon. The Ogooue River drains into the Gabon basin of Gabon. The River Congo drains into the Congo basin of Congo, Cabinda and Zaire. The River Cuanza drains into the Cuanza basin of Angola, and the Orange River drains into the Orange River basin of Namibia and South Africa (Fig. 1). These basins have been extensively explored for petroleum, with marked success in the Gabon, Congo and Angolan basins. Exploration in the Rio del Rey and Cameroon basins is proving exciting at the present time. Exploration in the Orange River basin has been slow up until recent years for political reasons. There is thus an extensive amount of geological information on these basins. Most of it, however, is maintained confidential in oil companies, and relatively little has been published. The following account draws extensively on papers published by Logar (1983) and Clifford (1986). The structure and stratigraphy of the coastal basins of west Africa are broadly similar in tectonic style and sequence of facies, and are indeed similar with the coastal basins of South America (Fig. 2). This similarity is because all the basins were formed in response to rifting, sea floor spreading, and the resultant separation of the South American

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and African continental plates (Bullard, et al., 1965). As the Atlantic Ocean gradually opened from north to south, however, the timing of structural events and facies sequences is diachronous in a southerly direction. Salt deposits, largely of Aptian age occur along the west African marginal basins between the Cameroon volcanic line in the north, and the Walvis

African Basins. Sedimentary Basins of the World, 3 edited by R.C. Selley (Series Editor: K.J. HsiJ), pp. 173-186. 9 1997 Elsevier Science B.V., Amsterdam. All rights reserved.

174

M.A. ALA AND R.C. SELLEY

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ridge, in the south (Baumgartner and Van Andel, 1971). Variations in the distribution of the evaporites and other facies reflect changes in patterns of oceanic circulation as the Atlantic widened, and also reflect changes in global climate (Evans 1978). The various coastal basins of west Africa will now be described in turn, from north to south.

THE RIO DEL REY AND DOUALA BASINS

The Rio del Rey and Douala basins are so small, similar and adjacent, that they can conveniently be considered together (Fig. 3). The Rio del Rey basin, or more strictly embayment, forms the eastern limb of the Benue trough and of the great Niger delta basin that infills it described in Chapter 7. Onshore the oldest sediments of the Rio del Rey basin that crop out are the Mundek Formation. This consists of coarse pebbly cross-bedded fluvial sandstones that are broadly comparable to the Bima Sandstone of Nigeria to the west. They are often

unfossiliferous, and thus hard to date, but in some instances can be attributed to Late Aptian to Albian ages. These deposits are overlain by shallow marine limestones, sandstones and shales of the Mungo and Logbaba formations. These range in age from Cenomanian to Maastrichtian. There is then a major unconformity, that is succeeded by nonmarine sands and silts of Miocene and Pliocene age. In the subsurface offshore, however, drilling activity by oil companies has revealed a stratigraphic sequence broadly comparable to that of the Niger delta (Short and Stauble, 1967). There is a diachronous sequence of pro-delta shales (the Akata Formation), overlain by delta front sands and shales (the Agbada Formation), in turn overlain by fluvial sands (the Benin Formation). These deposits range in age from Eocene to Recent. Not only the stratigraphy, but also the structural style of the Niger Delta extends into the Rio del Rey basin, with both growth fault and overpressured clay diapir zones being recognizable (Fig. 4). The Rio del Rey basin is separated from the Douala basin by a major tectonic feature, the Cameroon volcanic line. This is a line of volcanic centres that extends from the mainland of Africa out into the Atlantic Ocean, to include the volcanic islands of Principe, Sao Tome and Annabon. The Cameroon volcanic line marks a major change in tectonic style. To the west is a province of east-west aligned transverse faults. To the east of the line is a province of north-south aligned normal faults that are downthrown towards the Atlantic Ocean to the west. The volcanic line was established at least as far back as the Late Cretaceous (Santonian), and continues to be intermittently active down to the present day. It is correlated with the Pernambuco ridge of South America (Fig. 2). The Douala basin is larger than the Rio del Rey basin, and contains a more continuous stratigraphic section. Continental basal Cretaceous sands are overlain by shallow marine limestones, sandstones and shales of Late Cretaceous, Palaeogene and Neogene ages. At the western margin of the basin these formations are overlain by basaltic lavas from the Cameroon volcanic centre. When traced offshore the sedimentary formations thicken markedly, with evidence for over 7 km of subsidence since the middle of the Cretaceous Period. There are more sands and fewer shales than in the Rio del Rey basin, and growth faulting and diapirism due to overpressured shales are largely absent (Fig. 5).

THE GABON BASIN

The Gabon basin extends from southern Cameroon, through Equatorial Guinea and into

THE WEST AFRICAN COASTAL BASINS

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176

M.A. ALA AND R.C. SELLEY

Fig. 5. South-west to north-east cross-section of the Douala basin. For location see Fig. 3. Note the absence of growth faults and clay diapirs (based on Logar, 1983). Gabon, where it reaches its maximum thickness and extent, before narrowing and thinning south into Congo. The Gabon basin is much larger, and more important economically, than the Rio del Rey and Cameroon basins put together. It is the mirror image of the Campos basin of Brazil, with a comparable importance as a petroleum province. The Gabon basin contains up to 18 km of sedimentary fill, and is divided into two subbasins by the northwestsoutheast trending Lambarene horst. This separates the Eastern subbasin, to the northeast, from the Atlantic subbasin to the southwest. Both subbasins are cross-cut by 2 major northeast to southwest trending transfer zones, the Ascension and N'komi faults (Fig. 6). The tectonic and sedimentary history of the Gabon basin can be conveniently grouped in terms pre-salt, salt, and post-salt phases. These will now be described in ascending order. Pre-salt sequence

The sedimentary sequence of the Gabon basin unconformably overlies Precambrian igneous and metamorphic rocks. The lowest sediments are largely barren arkosic conglomerates, sands and minor shales, whose age is open to speculation. The fresh water ostracod Estheria, and fish scales have been found, but these are of little stratigraphic value. The first dated fossils are of Barremian age in what is termed the Cocobeach Group Thus these basal sediments are referred to as the pre-Cocobeach Group. This has been subdivided into various local names. The pre-Cocobeach sediments crop out onshore around the edge of the Eastern subbasin, but are overstepped by younger formations in the Atlantic subbasin. The pre-Cocobeach sediments are of nonmarine, largely fluvial origin, and may include formations laid down on the African Shield long

Fig. 6. Structural map of the Gabon basin (after Brink, 1974; Belmonte et al., 1965, and other sources).

before the Atlantic rifting phase commenced. The upper part of the section, however, may well have formed in response to rift initiation. The overlying Cocobeach Group contains fossils that provide dates from Barremian to Aptian age. Its base with the barren underlying pre-Cocobeach

THE WEST AFRICAN COASTAL BASINS Group is gradational. The upper contact, however, is well marked by a transgressive sand that indicates the arrival of the saline influx from which the Aptian salt deposits were formed. The Cocobeach Group consists of interbedded and laterally gradational fluvial sands and lacustrine shales. Turbidites are interbedded with the latter. The lacustrine shales are rich in organic matter and serve as a petroleum source bed. The interbedded turbidites and fluvial sands serve as reservoirs in a variety of fault-related structural and stratigraphic traps. The Cocobeach Group was deposited in a series of nonmarine rift basins that formed due to the onset of rifting between the African and South American continental plates. The combined thickness of the pre-salt deposits, the pre-Cocobeach and Cocobeach groups together, is in the order of 9000 m (Brink, 1974).

Salt sequence As the floors of the rift basins gradually subsided there came a time when the sea entered and disrupted the earlier fluvial and lacustrine depositional systems. This ingress of the sea is marked by a basal sandstone, the Gamba Formation. This is up to some 50 m in thickness in the east, but it gradually thins out in a westerly direction. It is dated as Aptian. The Gamba Formation is an important petroleum reservoir, by virtue of the Cocobeach source rocks beneath, and the succeeding salt seal above. The Aptian evaporites have a thickness of some 300 m, though this is very variable, and locally greatly exceeded due to halokinesis. The salt is distributed widely across both the Eastern and Atlantic subbasins, but is absent over the Lambarene horst. Halokinesis occurs in the Atlantic subbasin, but is absent in the Eastern subbasin, where the evaporites are thinner and shallower (Belmonte et al., 1965). The Aptian evaporites consist largely of halite, with sylvite and bischoffite in the areas of Lucina and M'bya. These localities would seem to mark

177 two depocentres of evaporite formation. Sylvite, being more mobile than halite, is commoner over the crests of diapirs. There are two major cycles of desiccation and evaporation. The evaporites pass up into anhydrite and the dolomites of the Madiela Formation (Logar, 1983).

Post-salt sequence Once the sea had invaded the rifts the separation of Africa and South America occurred rapidly. This resulted in the replacement of the evaporite episode, by an open marine regime that continues to the present day. Marine sediments of Albian to Recent Age occur in the western subsurface parts of the Gabon basin, showing a passage into nonmarine sediments eastwards towards the African coast. The sediments consist of a mixture of sands, shales and carbonates. The sands include paralic formations, such as the late Cretaceous Weze delta system. The coeval Batanga Formation, interpreted as of tidal channel origin, is an important reservoir (Clifford, 1986). Shelf carbonates, such as the Madiela (Albian) and Azile (Turonian) formations, pass basinward into shales. Organic-rich shales are a characteristic part of the post-salt stratigraphic section. Black shale deposition commenced in the Gabon basin in the Albian, reaching a maximum in the Cenomanian, and dying out in the Turonian stage (Tissot, et al. 1984). The distribution of post-salt facies in the Gabon basin is due to an interplay of three processes. The narrow seaway between Africa and South America favoured restricted oceanic circulation, hence anoxic conditions, and hence organic rich mud deposition. But this control declined as the Proto-Atlantic Ocean widened. Uniform global temperatures during the Cretaceous Period allowed whole ocean basins to become stagnant and anoxic, and also caused a global rise in sea level. High stands of the sea, such as occurred in the late Cretaceous, allowed anoxic conditions to spread across continental shelves, lead-

Fig. 7. Cross-section through the Gabon basin (compiled from Brink, 1974; Belmonte et al., 1965, and other sources).

178

M.A. ALA AND R.C. SELLEY

Fig. 8. Geophysical survey of WASP Line 8 crossing the North Gabon basin. From Meyers et al. (1996), courtesy of Elsevier. For location see Fig. 6. Upper: Magnetic and gravity anomaly profiles. Middle: uninterpreted seismic section. Lower: Interpreted seismic section showing Aptian evaporites. POC = Proto-Oceanic crust. PSR = Top of Pre-Salt reflector. RM = Reflection Moho. ZU = Zone of underplating. SRS? = Syn-rift sediments? AU = Anguille unconformity (Coniacian).

-,,,I

180 ing to the deposition of blankets of organic-rich clay. Conversely low stands of the sea allowed sand to be transported across the emergent shelves, over the shelf edge, to end up deposited as turbidites, grain flows and debrites in deep water fans at the foot of the continental slope. The Gabon basin is a major petroleum province with production from both beneath and above the salt. The pre-salt habitat has been discussed earlier. Though the post-salt sequence contains Cretaceous black shales, gas chromatography reveals that the post-salt hydrocarbons are of nonmarine origin. Furthermore they show identical chromatograms to those found in pre-salt reservoirs that were clearly sourced from the lacustrine shales of the Cocobeach Group (Clifford, 1986). Migration from the Cocobeach source rocks has been favoured by the laterally extensive Gamba sands. Petroleum has escaped across the salt horizon, either where the seal is breached by local salt solution, or where it is breached by listric faults. Both these mechanisms have allowed pre-salt petroleum to migrate up into post-salt shallow and deep water sand reservoirs. The main traps are provided by the Aptian salt domes (Fig. 7). Naturally the genesis of these structures has prompted much research, since the interplay of salt movement and petroleum generation is crucial to an understanding of petroleum entrapment (Ross and Hempstead, 1993, and Turner, 1995). Halokinesis has taken the form of both diapirs and salt walls. The salt walls occur on the uplifted, landward, side of listric faults that extend out towards the Atlantic Ocean. The faults curve down to coalesce at the base of the salt sequence to form a subhorizontal thrust (Fig. 8). This structural style is not restricted to the Aptian evaporites, but also occurs in plastic overpressured Tertiary clays higher up the sequence. These listric faults are clearly gravity-induced slump structures that developed in response to the separation of South America from Africa. This pattern of westerly dipping listric faults associated with the plastic deformation of evaporites and clays extends from the Gabon basin southwards to affect all the coastal basins of west Africa.

THE CONGO BASIN

The Congo basin occurs where the River Congo debouches into the Atlantic Ocean. Geographically it extends from southern Gabon, across Congo, Cabinda, Zaire and into northern Angola. Sediments crop out along the coast continuously connecting the Congo basin with the Gabon basin to the north, and the Angola/Cuanza basin to the south. As mentioned earlier, the sequence of sediments in the Congo basin is broadly comparable to that of the Gabon basin to

M.A. ALA AND R.C. SELLEY

AGE

LITHOLOGY

STRATI GRAPHY

" "" "'" :": ::'.:'." ...... ." -'.

Miocene La{e Eocene ~ 1 Oligocene hiatus ~

-

~

.'.

~

~

----

---- -

~

~

/ z" 9 .4../... ........9 9 ~

~ Madingo

9 ,

Loango

dolomite

ADrian

~

Sendji limestone

'.". :. " . ' . " . . . ' . :' . 9 . ,

Chela ~

to

.....

divers

-. :.~.........-.:. 9

,'"

-.

-.-....... -.:. :... 9

9

~

o,

.

~

",'

~

salt

Dre-

formation

, 9 9 , ,

9 ,,

9 "

,o. "-

9". ".. :.:.'. "..'.! Precambrian

salt

sandstone

9.'.':':

" 9 : .: . . . . .

9

;

Tchal a s a n d s t o n e

o

Loeme

Neocomian

marl ,-i

/~!'

2EEEEk__~

Albian

Paloukou

" -"

~

- -- |

Cenomanian

" ."

,

E o c e n e sen o n i a n Turonian

"o" o .

Vandji

sandstone basement

Fig. 9. Stratigraphic section of the Congo Basin, based on Logar et al. (1983).

the north. There are however some very significant differences. It is again convenient to describe the sequence and history in terms of the tri-partite division of pre-salt, salt, and post-salt sediments (Fig. 9). Pre-salt sequence

Precambrian basement is unconformably overlain by conglomerates, coarse arkosic sandstones and shales, variously termed the Vandji Formation in the southern part of the basin, and the Luculla Formation in the northern part. These are largely unfossiliferous and of continental origin. The absence of fossils means that their precise age is unknown. The Vandji and Luculla formations are overlain by a sequence of interbedded and laterally interfingering limestones sandstones and shales. Palaeontology shows that this sequence is non marine and of Neocomian to Barremian age. Lithostratigraphic nomenclature for these sediments is confusing, and varies across the basin. The limestones are developed on palaeohighs, and are referred to as the Toca Formation. They pass laterally into sandstones and shales that were deposited in palaeolows. These are collectively referred to as the Bucomazi Formation, but a distinction is sometimes made between the organic rich lacustrine shales, termed the Point Indienne Formation, and lenses of turbidite sands that are referred to as the Mengo Formation. This Neocomian section is attributed to deposition during the initial phase of rifting during which algal reefs formed around the edges of a series of anoxic lacustrine basins.

THE WEST AFRICAN COASTAL BASINS

181 east

west

~ea level

5000

0000

5000 Jepth tmetres) 50 km I

i

Fig. 10. Cross-section through the Congo basin, based on Clifford (1986) sequence is some 60 m in thickness, and of AIbian age. The succeeding Albian to Cenomanian sequence shows considerable lateral facies changes. Along the coast, in the east, paralic sandstones and shales are attributed to the Vermelha Formation. These sediments pass westwards into shallow marine carbonates of the Pinda Group. The open marine shelf was subjected to differential subsidence due to halokinesis of the underlying Loeme salt. This took the form of a series of listric faults sheafing out towards the Atlantic Ocean in the lower part of the evaporite sequence. The positive salt features lifted the floor of the sea to within the photic zone. This favoured carbonate sedimentation on these uplifts, coeval with clay deposition in the adjacent lows. once this pattern of sedimentation was established, differential loading enhanced the lateral differentiation of facies (Fig. 11). The Pinda and Vermelha groups are overlain by the Iabe Group. This ranges in age from Turonian up to Maastrichtian. In the eastern part of the basin the Iabe Group consists of interbedded shales and limestones which are locally phosphatised. Towards the base of the group, however, the carbonates are interbedded with sandstones that serve as petroleum reservoirs in the Malongo, Kungulo, Takula, and

Salt sequence A major unconformity separates the pre-salt sequences just described, from the salt sequence. This unconformity cross-cuts both the Neocomian and older barren formations to locally overly preCambrian basement (Fig. 10). The unconformity is immediately overlain by a laterally extensive transgressive basal unit, some 60 m in thickness, known as the Chela Formation. This is comparable to the Gamba Formation already described from the Gabon Basin to the north. It is largely sandstone, but there are some local occurrences of shales, calcareous siltstones and dolomites. The Chela Formation is of Aptian Age. The overlying evaporite sequence is referred to as the Loeme Formation. It consists of alternations of anhydrite, halite and sylvite, together with intercalations of black shale. Subsequent to deposition the evaporites have been deformed into both diapirs and fault-defined salt walls.

Post-salt sequence The Loeme evaporites are overlain by a thin sequence of anhydrites and dolomites, variously referred to as the Mavuma or Inhuca formations. This

sea level

, '

,

_

~

r

-

~

~

_

_

_

A

~

-

~

-

-

~

_

_

~

_

base of p h o t i c

_

9

9 9

.

.

9 9

~ o

9

ts" "

.

,

zone

_~_

.

-"" ,

Fig. 11. Cartoons to illustrate how halokinesis caused lateral facies variation in the Pinda Formation (Albian-Cenomanian) of the Congo basin. Carbonate sedimentation took place on the crest of salt growth structures where the sea floor was above the photic zone, this was synchronous with mud deposition below the photic zone. These lateral facies variations enhanced the original salt movement by differential loading. Developed from Logar et al. (1983).

182 other oil fields. Traced westwards towards the Atlantic Ocean, the carbonates and sandstones die out, and the whole of the Iabe sequence is composed of shales and marls, often referred to as the Madingo Group. The Cretaceous sediments are unconformably overlain by Palaeogene and Early Eocene limestones, siltstones and marls of the Landana Formation. There is then a second major depositional break that is common to all the west African coastal basins. Oligocene sediments are absent. The Early Eocene sediments are unconformably overlain by a thin sequence of Miocene to Recent sediments. A basal conglomerate is overlain by interbedded paralic sands and shales. The Congo basin, like the Gabon basin to the north, is a major petroleum province. There are many similarities in the style of petroleum generation and entrapment, but also some significant differences. Again, the pre-salt Neocomian lacustrine shales provide the source for both pre- and post-salt accumulations. Again migration pathways are essential to allow the petroleum to migrate across the regional salt seal into the post-salt section. Reservoirs occur in the presalt sands, such as the Luculla and Bucomazi formations in the Malongo West and North fields, but also in the algal reefs that bordered the lacustrine depocentres and in the transgressive Chela basal sand of the salt sequence. These accumulations occur in faulted and truncated combination traps. Post-salt reservoirs are largely to be found in the sands and limestones of the labe Group where they are structurally closed above salt diapirs.

T H E A N G O I , A (OR CUANZA) BASIN

The Congo basin thins southwards into a narrow coastal strip of sediment only a few kilometres wide, before it widens out again and thickens to become the Angola basin where the Cuanza River enters the Atlantic Ocean. Hence the Angola basin is also referred to as the Cuanza (or Kwanza) basin (Fig. 12). The sedimentary and structural history of the Angola basin are so similar to those of the basins to the north already described, that it can be dealt with relatively briefly, though it too is a major petroleum province. The sediments of the Angolan basin can be divided into pre-salt, salt, and post-salt sequences (Fig. 13). Pre-salt sequence

The pre-salt sequence consists of interbedded clastics and volcanics of the Lower Cuvo Formation. The clastics include continental red sandstones and shales with occasional thin coals. The volcanics consist of basaltic lavas and tufts. The Cuvo Formation

M.A. ALA AND R.C. SELLEY

:-t-

4-

'P"

i. t -

,."~-t-- §

+ "+ @

~-

x~'.'...N.4- , -

Luanda

_

-4-

-4.

~

+

,

Ir

'

/

i

t

r

t,

/

i

I

,

t.

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r

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~L

Pre-Aptian

~-~ P r e c a m b r i a n

~

4q.

4-

4-

4

-t-4-

~

+

*

t"

-~

field limit of aptian salt basin fault

basement

Fig. 12. Map of the Angola (Cuanza) basin. Based on Brognon and Verrier, 1966.

appears to infill a half graben, being thickest along the eastern margin of the basin' and thinning westwards towards the present Atlantic coastline (Bouju, 1977). Because of a paucity of fossils it is hard to date the Lower Cuvo formation, but its geometry suggests that it formed during the rifting phase immediately prior to Cretaceous crustal separation (Duval et al., 1992). Salt s e q u e n c e

The Lower Cuvo Formation is overlain with an angular unconformity by the Upper Cuvo Formation. This varies in thickness from 100-200 m and consists of shallow marine well sorted sands and limestones. The Upper Corvo Formation marks the arrival of marine conditions and is thus correlatable with the Chela and Gamba formations described from the Congo and Gabon basins respectively. The Upper Cuvo Formation grades up via dolomites and bituminous shales into the Aptian evaporites of the Tuenza Group. Some 1500 m in thickness, these pass eastwards into red beds along the basin margin (the Dondo Group). Traced

THE WEST AFRICAN COASTAL BASINS LITHOLOGY

&

183 with occasional marginal bituminous layers suggestive of sabkha conditions. Evaporitic deposition was terminated with the widespread deposition of the Catumbela Limestone Formation9 Aptian evaporites can be traced southwards along the African coast as far as the Walvis ridge, which, together with the Cameroon Volcanic Line to the north, appears to have confined the extent of the salt basin9

STRATIGRAPHY

AGE west

east

. . . . . . . . . . . " " " ' " " " ' " ' ." ' ". . . . . .. . . . . .. . 9 , . . . . . . . ....:.:

Miocene

to Recent .._~.~_~~.

~ Hiatus

Oligocene

Eocene

Palaeocene

~

G

.

.

r

a

t

.

i

d

a

o

.

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~

"" " " "

" - " : " : :: " " : i i i i

9 9 .,........., 9 . 9

. . 9.

- R i o G r a n d e -

""

"

~N'golome

.

=-Itombe

Tu r o n i a n

.

.

.

'9 1

~

_

I

~

Post-salt sequence

"''"'"

:"." 9

9 " .'.

.... 9

.

.

Gabo

-

Cenomanian

,-...,. ............-.

. . . . . . . .

9 . 9 .

Senonian

A I bi

~

. Ledo--'

. "

'=

r

I

I

i

Following the main episode of evaporite formation, coastal evaporites continued to form throughout the late Cretaceous and early Eocene times, passing westwards into deep water shales and turbidite sands9 In common with the rest of the west African coast, Late Eocene and Oligocene sediments are absent. Lower Eocene sediments are unconformably overlain by Miocene and younger clastics. Paralic sands in the eastern part of the basin pass westwards into shales and marls9 Halokinesis started in the Angola basin before the end of the Cretaceous Period9 Salt domes in the east pass westwards into a terrain of diapirs (Fig. 14). An east-west alignment of salt features implies basement-related structural control9 Movement was continuous throughout much of the Cenozoic Era, resulting in a strong control on the distribution of Eocene and Miocene sedimentary facies9 This was exacerbated by the fact that many of the salt structures are of piercement type, and actually reached the surface, causing salt solution hollows to form. Withdrawal of salt adjacent to active piercement structures resulted the formation of rim synclines and concomitant "turtle-back" structures in the adjacent lows (Lundin, 1992). Petroleum seepages and asphalt deposits occur in Angola, so exploration for petroleum has a long history. Pre-salt reservoirs are of poor quality and hard

-

-

'

an

Qu i sson d e I

-

1

L'Lr LC

"

L L_ k_

~ C a t u m b e l a T u e n z a .

_

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L. " 9 ~ ~ ~ ~ : : : : : : L L. L.-."Quiang'a-l-b'nga~_ kS. L. LDgn.do reef~---l-' Masslve . ." . [ . [ . " ~. /I / / /salt,. ~ - ' . ' . ' ./' . ' . / /

Aptian

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Neocomian

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V

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:

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\ V o / c a n" i c s N / -

.

.

.

.

Precambrian

Fig. 13. Stratigraphic column of the Angola (Cuanza) basin. Based on Brognon and Verrier, 1966. westwards they pass into a barrier reel the Quianga Group, along the Cabo Ledo ridge9 This distribution of facies suggests that the evaporites may have formed in a classic barred basin9 The Tuenza Group consists of halite, anhydrite, dolomite cycles

~

9

9

-

""

"

. . . . .

9

~

~

9

9

.

9

_

9

_

9

Miocene

to Recent

" ~

9

9

9

9

9

~

. . . .

9

9

9

9

9

.

"

i

~ ,

~ ~

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9

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,

sealever

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../m

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~ ~ ~

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9

~ . - - ~ 9

o

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.

~

!

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4-

4-

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.,

2000

9

-3000

uto,

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--i'/4.'4-

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+

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+

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depth

Ld n O t,

Fig. 14. Cross-section through the Angola basin (based on Clifford, 1986).

5 |

10 k m 1

(metres)

184 to locate. Most of the production comes from within and above the salt sequence. Production comes from carbonate reservoirs within the salt sequence, such as the Binga Limestone in the Tobias, Glenfica and Galinda fields, and in the overlying carbonates, such as the Catumbela Limestone in the Mulenvos Field. Oil is also produced from Eocene and Miocene sands. All of the traps are related to salt tectonics. They are either simple domal traps above a diapir, or more commonly, complex faulted, and/or truncated reservoirs on the flanks of salt domes. Production also comes from "turtle-back" structures in the adjacent lows (Walgenwitz et al., 1990).

M.A. ALA AND R.C. SELLEY

",',k I, , ~ s I

-, '),~waJvis

Bay

I

Z',I, , .

, , : . .. " - -

,

i

~__~....o

.,'

',i". I

i

I

Namibia

2

-

|

a

j

9

B.,,~

!

I I

I

I q

Orange River

South

THE ORANGE RIVER BASIN

b

%

The last of the west African basins to describe is the Orange River basin. This is infilled by sediments brought down by the Orange River drainage system to reach the Atlantic coast at Alexander Bay. The Orange River marks the boundary between Namibia to the North and South Africa to the south (Fig. 15). Early accounts of the Orange River basin described a total sediment column of some 4 km in offshore Namibia (Lehner and De Ruiter, 1977). More recent accounts from the South African end of the basin, supported by drilling results, show that the total sediment column may be in excess of 10 km (Muntingh, 1993, Muntingh and Brown, 1993). Regional seismic lines show that the sedimentary rocks can be divided into two megasequences. The lower one is related to rifting, the upper one related to drifting. Faults in the lower sequence die out at the drift-onset unconformity. Faults in the upper megasequence are listric and westerly dipping, toeing out in overpressured shales close to the drift onset unconformity (Fig. 16). As mentioned earlier, evaporites are absent south of the Walvis ridge. Thus the tripartite stratigraphic description used for the northern basins is replaced by 2 sections on the syn-rift and post-rift drift sequences.

Syn-rift sequence Unlike the coastal basins described to the north, sediments of the Orange River basin do not crop out onshore. The coastal outcrops consist of Precambrian igneous and metamorphic basement overlain by a veneer of Quaternary deposits. Seismic studies show that the syn-rift sediments occur in a series of marginal rift basins adjacent to the present-day continental margin and a central rift sequence that lies beneath the present-day continental shelf margin (Fig. 16). In common with the rift-related sediments

,

\

\

.

%

%

..J

%

%

Tow n

\ '

O

Africa

~

\

""

"3?,,

'~%

.

~ m.,-

...

500 km f

key ---1---

isopachs of sedimentary section(kin) faults . . . . . . . approximate boundary of Namibian:South Africa median I~ne 9 Kgf=Kudu gas field A ~ ~ line of cross-section B

Fig. 15. Map of the Orange River basin, based on Muntingh (1993) and Muntingh and Brown (1993).

of the basins to the north, it is not possible to know the age of the oldest sediments within these rifts. There may be sediments of diverse Phanerozoic ages as yet undrilled. On the basis of seismic correlation with a 117.5 my magnetic anomaly in the adjacent Eastern Cape basin, Muntingh (1993) dates the drift onset as Late Hauterivian. In wells in the Kudu gas field in Namibia barren sediments are overlain at the drift onset unconformity by shallow marine sediments dated by McLachlan and Wickens (1990) as Barremian (113-116.5 my). These data suggest that the syn-rift sediments are Lower Cretaceous, and older. It is possible that pre-rift Phanerozoic sediments occur on the floors of the rifts. On the basis of wells drilled in Namibia and South Africa the sedimentology of the sediments truncated by the drift-onset unconformity is tolerably well known. They consist largely of red beds, that include barren conglomerates, sands and shales. These are believed to be largely of fluvial origin. Some of the sands however are very well-rounded, anhydrite cemented, and exhibit thick sets of high

THE WEST AFRICAN COASTAL BASINS

185

A

B

Barremian-Eocene ~ -~'~-:r

a

~

+

~

NW =

~ ~ +

//~

+" 4-::n't: : : t a:

'

depth (metres)

;us: 4 - : : : : : ~

500 km

~;:22

~ SE

y

Z

SW -~

300km

~ NE

Fig. 16. Diagrammatic cross-sections of the Orange River basin. Upper, based on Lehner and De Ruiter, (1977). Lower, based on Muntingh (1995) and Muntingh and Brown (1993). Age calibration of seismic reflectors is based on drilling results from the DSDP, the Kudu gas field of Namibia (McLachlan and Wickens, 1990), and wells drilled offshore South Africa.

angle cross-bedding. These have been interpreted as of eolian origin. Some wells have also encountered nonmarine black shales with significant quantities of oil-prone kerogen. These shales presumably formed in lakes during more humid episodes. Wells have also encountered igneous rocks. These consist largely of basaltic amygdaloidal lavas, with bombs, oxidized soil horizons and eolian sand interbeds (McLachlan and Wickens, 1990). Petrographically some of the sands are volcanoclastic, derived from the lavas with which they are interbedded, and others are arkosic, indication derivation from the Precambrian basement.

ACE L C~

9

.....

" . . . . .

9 ..

if3 (1) I::~u c

9

=- ".:.:.':. i

~ ffl

-

L

121. ffl D

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>,0 -C ~

Cn C

.

~ ~

~_ t_ "1:3

_

0 U ,u

." 9. I

. . . . .

u I

U~ C C

I black shale

ffl

Post-rift (drift) sequence According to the data presented above, the drift onset unconformity in the Orange River basin is dated at around 115 my (Hauterivian-Barremian). Where the unconformity has been drilled and cored, as in the Kudu gas field, the sediments immediately above the unconformity are shelly bioturbated sands, occasionally glauconitic, and with algal oncolites (McLachlan and Wickens, 1990). These parameters suggest a very shallow marine origin for the basal sands. This unit is about 100 m thick in the Kudu field, and invites correlation with the Upper Corvo, Chela and Gamba formations described from the Angola, Congo and Gabon basins respectively. Here the comparison ends, however, as the evaporites that overly the basal sands in the basins to the north are absent. Instead the shallow marine basal sand is directly overlain by up to 6 km of clastic sediments (Fig. 17). These consist of shallow marine

]

LITHOLOGY I w E..

. . . . . . . 9 .

I

(anoxic event) I

black sha e anoxic event) V

v

I v

Pr

caml

Fig. 17. Diagrammatic stratigraphic column of the Orange River basin, based on Muntingh (1995) and Muntingh and Brown (1993). Evidence for the Cretaceous anoxic events comes from the results of the DSDP; for sources see Lehner and De Ruiter (1977) and Tissot et al. (1984).

sands along the coast, that include both progradational deltaic and aggradational coastal sequences. The sands die out westwards into a thick shale sequence that is overpressured. These include organic

186

rich intervals dated as mid-Aptian and CenomanianTuronian. Submarine fan sequences have been identified seismically in upper Lower Cretaceous and lower Upper Cretaceous sediments. The intimate details of the seismic stratigraphy of the whole sequence has been resolved down to the very last high tide (Muntingh and Brown, 1993).

REFERENCES Baumgartner T.R. and Van Andel T.J.H., 1971. Diapirs of the continental margin of Angola, Africa. Geol. Soc. Am. Bull., 72: 783-802. Belmonte Y., Hirtz P. and Wonger R., 1965. The salt basins of the Congo. In: Salt Basins Around African. The Institute of Petroleum, London, pp. 55-74. Bouju, J.P., 1977. Etude Geologique d'une savane arborescente: le bassin du Cuanza (Angola). Notes et Mem. de C.EP. Paris, pp. 31-42. Brink A.H., 1974. Petroleum geology of the Gabon Basin. Am. Assoc. Pet. Geol. Bull., 58: 216-235. Brognon G. and Verrier G.1966. Oil and geology in Cuanza Basin of Angola. Am. Assoc. Pet. Geol. Bull., 50: 108-158. Bullard E.C., Everett J.E. and Smith A.G., 1965. The Fit of the Continents around the Atlantic. In: P.M.S. Blackett, E.C. Bullard and S.K. Runcorn (Editors), A Symposium on Continental Drift. R. Soc. London, Phil. Trans., Ser. A, 258: 4145. Clifford, A. C., 1986. African oil - - past, present and future. In: M.T. Halbouty (Editor), Future Petroleum Provinces of the World. Am. Assoc. Pet. Geol. Mem. 40: 339-372. Duval, B., Cramez, C. and Jackson, M.P.A., 1992. Rift Tectonics in the Kwanza Basin, Angola. Mar. Pet. Geol. 9: 389-411. Evans R., 1978. Origin and significance of evaporites in basins around Atlantic Margin. Am. Assoc. Pet. Geol. Bull., 62: 223234. Lehner P. and De Ruiter P.A.C., 1977. Structural history of Atlantic Margin of Africa. Am. Assoc. Pet. Geol. Bull., 61: 961-98 !.

M.A. A L A A N D R.C. S E L L E Y Logar, J.E, 1983. Afrique de l'Ouest. Well Evaluation Conference. Schlumberger, Paris, 205 pp. Lundin, E.R., 1992. Thin-skinned extensional tectonics on a salt detachment, northern Kwanza Basin, Angola. Mar. Pet. Geol. 9:405-411. McLachlan, I. and Wickens, H. de V., 1990. The stratigraphy and sedimentology of the reservoir interval of the Kudu 9A-2 and 9A-3 boreholes, offshore Namibia. In: Abstract 23, Earth Sci. Congr. Geol. Soc. S. Afr., pp. 359-362. Meyers, J.B., Rosendahl, B.R., Groschel-Becker, H., Austin, J.A. and Rona, P.A., 1996. Deep penetrating MCS imaging of the rift-to drift transition, offshore Douala and North Gabon basins, West Africa. Mar. Pet. Geol. 13: 791-836. Muntingh, A., 1993. Geology, prospects in Orange basin offshore western South Africa. Oil Gas J., Jan. 26, pp. 106-109. Muntingh, A. and Brown, L.E, 1993. Sequence stratigraphy of petroleum plays, post-Cretaceous rocks (Lower Aptian to Upper Maastrichtian), Orange Basin, Western Offshore, South Africa. In: P. Weimer and H.W. Posamentier (Editors), Siliciclastic Sequence Stratigraphy. Am. Assoc. Pet. Geol. Mem. 58: 43-70. Ross, D. and Hempstead, N., 1993. Geology, hydrocarbon potential of Rio Muni area, Equatorial Guinea. Oil Gas J., Aug. 30, pp. 96-100 Short K.C. and Stauble A.J.1967. Outline of geology of Niger Delta. Am. Assoc. Pet. Geol Bull., 51: 761-779. Tissot, B., Demaison, G., Masson, P., Delteil, J.R. and Combaz, A., 1984. Paleoenvironment and petroleum potential of middle Cretaceous black shales in Atlantic basins. In: Petroleum Geochemistry and Basin Evaluation. G. Demaison and R.J. Murris (Editors), Am. Assoc. Pet. Geol. Mem., 35: 217-228. Turner, J.P., 1995. Gravity-driven structures and Rift Basin Evolution: Rio Muni Basin, offshore Equatorial Africa. Bull. Am. Assoc. Pet. Geol., 79:1 ! 38- I 158. Walgenwitz, F., Pagel, M., Meyer, A., Maluski, H. and Monie, P., 1990. Thermochronologicai approach to reservoir diagenesis in the offshore Angola basin: A fluid inclusion 4~ and K-Ar investigation. Am. Assoc. Petrol. Geol. Bull. 74: 547563.

Chapter 9

The East African Rift Basins

L.E. FROSTICK

INTRODUCTION The East African Rift is perhaps one of the best known continental rifts which exists on the present Earths surface. It forms a striking geomorphological feature which runs from Afar in the north of Ethiopia to B lantyre to the south of Lake Malawi, a total length of approximately 35,000 km (Fig. 1). At its deepest, in the Ethiopian segment of the rift, it is over 3 km deep. Along most of its length it achieves a depth of over 1 km with its own localised microclimatic and hydrological environment which has, in places, led to the development of a unique flora and fauna (e.g. in the Magadi basin, Burgis and Gaudet, 1981). It cuts through the African craton and acts as both a north-south corridor for, and an east-west barrier to, the migration of animals and birds. It is not surprising that this area has played an important role in evolution, and that it is a prime candidate for the title "The cradle of Mankind" (Leakey, 1973). The hominid finds throughout the length of the rift have been both more numerous (Coppens et al., 1976) and more complete than in any other part of the World and they have allowed palaeontologists to construct an evolutionary sequence for Mans development which now has few missing links. Geological interest in the Rift dates back as far as Thomson and Von Hohnels' expeditions of 1883 and 1887-8 and their publications of 1887 and 1894 respectively. Since this time the intensity of research has waxed and waned, probably reaching a peak during the 1970s and early 1980s when many of the most important hominid sites were found (Leakey et al., 1976). There are numerous publications about the area dealing with the full spectrum of geological topics from plate tectonic setting (McKenzie et al., 1970) to details of the microflora and-fauna (Haberyan and Hecky, 1987). It is undoubtedly one of the best examples of an

active rift that exists today and has been used as an analogue for the interpretation of older, subsurface basins worldwide (e.g. Frostick and Reid, 1987).

GEOMORPHOLOGY

The East African Rift consists of a single valley in its northern most section. Between Afar and the Kenya-Ethiopia border the valley floor cuts through a domed area 1000 km in diameter which includes most of Ethiopia and southern Yemen. The floor of the rift is punctuated by a series of seven lakes which, although now largely ephemeral, have varied greatly in size during the Pleistocene and early Holocene (Grove, 1986). The doming has effectively diverted river drainage away from the rift (Fig. 2; Frostick and Reid, 1990) and the western half of the Ethiopian dome is drained by rivers which feed into the Nile. The eastern half of the dome is drier and all drainage is towards the Indian Ocean, although little water survives evaporation losses to reach the sea. South of the Ethiopian border is one of the larger lakes in the rift, Lake Turkana. This lies to the south of the domed area and its surface is 1300 m below that of the highest lake in Ethiopia. It is fed by a number of exotic perennial rivers including the Omo, Turkwell and Kerio and in the past has received overflows from Chew Bahir to the north and the Suguta trough to the south (Grove et al., 1975; Truckle, 1976, Frostick and Reid, 1989). For most of the period between 8500 and 3000 BP Lake Turkana stood more than 80 m above its present level and drained to a tributary of the White Nile (Butzer, 1980; Fig. 3). To the south of Turkana the rift bifurcates around Lake Victoria giving an eastern and a western branch with contrasting geomorphological characters. In the eastern branch the rift again crosses a domed area. The floor of the rift rises to 1890 m at lake

African Basins. Sedimentary Basins of the World, 3 edited by R.C. Selley (Series Editor: K.J. Hsti), pp. 187-209.

9 1997 Elsevier Science B.V., Amsterdam. All rights reserved.

188

L.E. FROSTICK

N

)~..~;

thiopian Rift

BLOCK

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regory R i f t

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ff ~4'

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g

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Naivasha and then falls again to 600 m at Lake Natron. Most of the lakes in this branch of the rift are small as a function of both the relatively low rainfall in the area and the small catchments of the streams which debouch into them (Fig. 2). The lake basins are generally closed, with volcanic and structural barriers preventing overspill from one basin into another. This, combined with high rates of evaporation and considerable addition of salts from hydrothermal and volcanic sources, has led to a number of the lakes becoming saline notably lakes Natron and Magadi (Eugster, 1986, Vincens and Casanova, 1987).

The lakes of the western branch are, in contrast, generally deeper and fresher the deepest being Lake Tanganyika (1400 m). This is partly a function of the low rates of sediment supply to the western rift where the main rivers have drainage basins on crystalline basement rocks which are resistant to erosion (Fig. 4). The rift branches reunite to the north of Lake Malawi which sits in a rift valley similar to that in the western branch. Other rift-like structures can be traced southwest from Rungwe by way of the Luangwa valley and Lake Kariba into the Kalahari where they are finally obscured by blown sand.

THE EAST AFRICAN RIFT BASINS

189

Fig. 2. River drainage and structure in the area around the East African Rift (after Frostick and Reid, 1989). CLIMATE AND VEGETATION The length of the rift, spanning the equator and extending from latitude 12 N to 15 S dictates a range of climate types. In addition the uplift which has accompanied rifting has led to the development of local microclimates, particularly on the rift flanks. Rainfall is lowest in the Afar region and increases southwards into northern Kenya. The whole of this area is arid to semi-arid and vegetation is limited to desert grass and scrubland (Fig. 5). The climate in the eastern branch of the rift is drier than in the western branch. Rainfall varies but is generally around 250-500 mm allowing the growth of Acacia bush and seasonal grasslands. In the western branch average annual rainfall is almost an order of magnitude greater, for example a value of 2000 mm

is recorded in the Ruwenzori mountains near Lake Mobutu. Further south Lake Malawi occupies the wettest section of the rift with rainfall approaching 3000 mm annually. Tropical forests dominate these areas. The uplifted flanks of the rift are wetter and cooler than the rift valley and as a result have a mountain forest type of vegetation, even in otherwise arid zones, for example in Ethiopia the Mount Abuye Meda stands 4000 m high above a rift valley floor at less than 500 m.

PRE-RIFT GEOLOGY

The geology of the area through which the rift is cut influences not only the way in which the structure can develop but also the character of the sedi-

190

L.E. FROSTICK

Fig. 3. Present lakes and rivers with the expanded systemsof 9000 BP superimposed (after Frostick and Reid, 1989) mentary fill. Softer, more easily weathered sedimentary rocks are quickly reworked into the evolving basins giving a dominantly clastic fill, whereas hard crystalline basement provides little detritus allowing the basin to become deep with a sedimentary fill which relies upon evaporitic and biogenic sources. The East African Rift crosses a variety of geology along its length, although its setting within one of the worlds major cratons dictates a strong influence of crystalline basement. In general the location and character of the rift is controlled by the mechanical anisotropy of the crust. The sedimentary basins are therefore more or less confined to the orogenic belts surrounding the stable cratons. The cratons have

been subject to repeated phases of metamorphism and as a result they are relatively homogeneous (Fairhead and Green, 1989). In contrast mobile belts have a strong fabric which gives them lines of weakness that are easier targets for rifting. In Ethiopia Precambrian rocks are associated with Palaeozoic metasediments and extensive marine Mesozoic rocks. Further south much of this cover disappears and both branches of the rift cut directly through Precambrian rocks, but generally located in the younger Precambrian metasediments particularly in the Eastern branch. To the west of the western branch stable cratonic rocks are resistant to weathering and erosion. This tends to restrict

THE EAST AFRICAN RIFT BASINS

Fig. 4. Distribution of Precambrian basement rocks and older fold belts.

the supply of clastic sediment to the associated basins. To the south, around Lake Malawi, the geology changes to incorporate Permo-Triassic Karoo sediments. These rocks sit in an earlier rift phase which trends northeast-southwest (Fig. 6). They are mainly continental deposits comprising mixtures of shales, sandstones and conglomerates not dissimilar to those found in the later rift basins of the area (Kreuser et al., 1990). The crust of East Africa has been put under stress a number of times in its long history and, as a result, exhibits a variety of structural trends which can act as lines of weakness during subsequent rifting. Lambiase (1989) suggests that there have been 7 phases of extension since the early Permian. Notable structural trends are those of the Central African shear zone also with a northeast-southwest trend similar to that of the Karoo, and the southern Sudan rift which is oriented northwest-southeast.

PLATE TECTONIC SETTING

The development of plate tectonic theory during the early 1960s helped geologists to place the East African Rift in its wider context, linking it with crust and mantle processes in the Red Sea-Gulf of Aden. In plate tectonic terms it is one arm of a

191 trilete system of rifts which meet in a triple junction in the Afar region of Ethiopia (Oxburgh, 1978; Fig. 2). In the Red Sea and Gulf of Aden there is evidence that sea floor spreading commenced in the Middle to Upper Miocene (Bonatti, 1985). The presence of high concentrations of heavy metals in this area is typical of the early phases of oceanic evolution. If present trends continue into the future it is to be expected that the Red Sea-Gulf of Aden system will become the next major ocean and that the East African Rift, although still seismically and volcanically active, is actually a failed rift or aulacogen which will slowly die in the next few million years. Rifts can develop in a variety of plate tectonic settings and rely only on the existence of crustal extension which may be generated in a variety of ways. Examinations of the structural evolution of the East African Rift and other present day and older rifts have led earth scientists to think that there are two distinct types of rift which pass through very different phases in their evolution. These have been termed mantle-generated or active rifts and lithosphere-generated or passive rifts. Mantle-generated rifts are thought to develop at sites of mantle upwelling and are characterised by doming, arching and uplift on a regional scale. The fundamental control on the development of this type of rift is thought to be the presence of upwelling hot mantle material beneath the crust (Bailey, 1972). The diagnostic features of active rifts are: (1) doming which occurs early in the development of the rift and which can divert river drainage away from the developing basins, and (2) very high heat flow accompanied by widespread and intensive volcanic activity which begins early in the history of the rift. In lithosphere-generated rifts the first topographic expression of tectonic activity is subsidence (McKenzie, 1978). Doming, if it occurs, is a later event related to the thermal anomalies generated by crustal thinning and faulting. Subsidence is initiated as a result of tension in the lithosphere which stretches and thins it (Fig. 7). The essential features of this type of rift are: (1) the early development of a topographic depression, a sag-type basin, (2) uplift later in the development of the basin, as a result of the asthenosphere rising towards the surface (McKenzie, 1978), and (3) little volcanism and generally restricted to later stages of basin development. If these ideas are correct, the East African rift would seem to be a good example of an active rift. Evidence of early doming is widespread (Fairhead, 1986; Fig. 2) particularly in the eastern branch. Volcanic activity varies in character and intensity along the rift but dating often proves that it commenced at an early stage of basin development (Baker, 1986). However, the lack of doming and small amount of

192

L.E. FROSTICK

Fig. 5. Vegetation map for the East African Rift and surrounding area. volcanism in the western branch of the rift is consistent with a passive rather than an active rift. It appears that the distinction between active and passive rifting is, in fact, unhelpful in this area and that the basins all result directly from lithospheric extension.

STRUCTURE

Mantle and lower lithosphere Geophysical studies of the area around the Rift show that it is underlain by a zone of hot, low density mantle which extends far beyond the rift shoulders. It is approximately 1000 km across and is interpreted

as an area where crustal thinning is accommodated by the rising asthenosphere (McKenzie, 1978; Fairhead, 1986). The extent of the thinning is delimited by a zone of strongly negative gravity anomalies, ranging from - 1 0 0 to - 2 4 0 mGals. The highest negative values occur in Ethiopia and Kenya where there are two domes approximately centred on Robit and Nakuru respectively. It is these areas which are the foci for intensive volcanism and where the development of the rift valley is most marked. Anomalously low mantle velocities and high upper mantle electrical conductivities (Fairhead and Reeves, 1977) also suggest a broad uplift in the lithosphere-asthenosphere boundary under these areas (Fairhead, 1976).

THE EAST AFRICAN RIFT BASINS

P

193 research groups, notably the PROBE and KRISP surveys (Rosendahl and Livingstone, 1983; Khan et al., 1986; Rosendahl, 1987). These data have revolutionised thinking about rift structure and led to the development of models for rift development which are closer to, and help to explain, field observations (Gibbs, 1984; Rosendahl, 1987; Sander and Rosendahl, 1989; Chorowicz, 1990). Rifting in the East African system is thought to have begun in the early Miocene and continues to the present day. Earthquake data suggest that the system is still propagating to the southwest (Fairhead and Henderson, 1977) and that new basins may still be developing. Indeed active surface faulting is recorded in many areas of the rift (Johnson et al., 1987). Extension is generally in a N W - S E direction but the major rift basins vary in orientation from almost N-S to N W - S E and N E - S W (Fig. 1). Orientation and the location and character of the main faults are controlled by the local structural grain and by pre-existing lines of weakness. The fundamental rift unit is now accepted as being the half graben. Strong asymmetry is the hall mark of almost all of the rift segments (Rosendahl, 1987; Rogers and Rosendahl, 1989) with a single or sometimes two main border faults on only one side of the basin (Fig. 8). These faults are normal faults with throws often in excess of 2 km (Chorowicz, 1989). Along the opposite margin antithetic and synthetic faults with much smaller throws cut the structure into a series of tilted fault blocks.

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~

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Our ideas of rift structure have changed considerably over the past decade as a result of the acquisition of seismic data for rifts in a variety of different plate tectonic settings, particularly the East African Rift. Here both deep and shallow seismic surveys have been carried out by a number of

(a)

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rift

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Fig. 7. Sketch diagram to show the processes of active and passive rifting (after Turcotte and Emerman, 1983)

194 This margin has been termed variously a "ramping margin", a "flexural margin", a "faulted flexure" or a "flexural monocline" (Frostick and Reid, 1987; Chorowicz, 1989). The width of the basins varies between 30 and 200 km the widest of the basins being in northern Ethiopia where the rift expands to meet the Red Sea. At intervals along the rift axis there are topographical barriers which vary in height. The highest of the barriers separate the rift into a series of hydrologically isolated drainage basins. River linkage between these particular basins is only possible during periods of higher rainfall than at present (for example during the last glacial period at around 10,000 yr BP). The topographical barriers reflect a subsurface compartmentalisation of the structure, the precise nature of which is still the subject of debate. The debate arises, at least in part, from the variety of features noted in different barriers: it is evident that some barriers have faulting as a major formative process whereas others do not; faulting varies greatly in its intensity and character; the geometry of the zone is variable; some barriers are dominated by volcanic vents while others have no volcanogenic features. What is accepted is that these transverse or oblique structural elements are an essential feature of rifting (Gawthorpe and Hurst, 1993). They have been variously called transfer zones, accommodation zones, relay zones, relay ramps and segment boundaries (Gibbs, 1984; Rosendahl, 1987; Larsen, 1988). Some zones occupy the sites of large intracontinental faults and are better termed "transform faults" (Chorowicz, 1990). A good example of this is the Aswa fault zone which forms the northern limit of the western branch of the African rift and then continues southeastward to the Indian Ocean. Research carried out during the 1980s suggested that there are two main types of transfer zone: (1) those comprising defined faults and called a hard-linkage system (Bally, 1982; Gibbs, 1984), and (2) ramp or relay zones without defined faults and linking en echelon border faults. These are considered to be soft-linkage systems (Rosendahl et al., 1986; Larsen, 1988). Subsequent work in East Africa has suggested that these two types of transfer form the two end members of a range of geometries (Morley et al., 1990). The distance between adjacent transfers varies between tens and hundreds of kilometres. Displacement along individual border fault segments is known to vary systematically (Peacock and Sanderson, 1991; Walsh and Watterson, 1991; Roberts et al., 1993). Displacement is zero at each tip of the fault and increases toward the centre. This means that the topography of the footwall diminishes towards the transfer zone whereas that of the hanging wall increases. The rift is therefore com-

UE. FROSTICK partmentalised into a linked chain of subbasins, the degree of connection between them depending on the height of the separating barrier. There has been some speculation on the precise nature of the border faults. Seismic sections shot across particular sections of the rift (e.g. Lake Tanganyika, Rosendahl, 1987) suggest that surface faults sole-out or detach within the crust rather than penetrating individually to the mantle. In addition border faults linked to detachment zones are consistent with the surface features mapped in many areas, for example Bosworth (1989) produces a structural interpretation of the Baringo area with an eastward dipping detachment at a depth of 15 km with extension estimated at 11 km. He maintains that this arrangement can explain all of the surface features reported by Chapman et al. (1978). The dominance of synthetic over antithetic faults is explained by kinematic considerations such as the intersection of fault planes at depth and their rotation to positions of low shear stress. The result is an asymmetric structure with antithetic faults "locking" due to their poorer access to the detachment surface (Fig. 9, after Bosworth, 1989). An alternative to the model of listric faults and defined detachment zones is brittle deformation in the upper crust, along essentially planar faults, which is substituted by uniformly distributed creep at depth (McKenzie and Jackson, 1986). The main evidence supporting this interpretation is the distribution of earthquake foci along major border faults. These tend to describe a planar pattern down to 15-20 km, below which deformation is essentially aseismic (Chen and Molnar, 1983). Either if these two interpretations can be used to explain the observed surface features in the East African Rift. It is impossible to resolve whether the faults are planar or listric and if a defined detachment exists at depth.

VOLCANICITY Volcanism is closely associated with the domed areas of the rift. From Afar southwards through Ethiopia and into the Kenyan or Gregory rift volcanic features are widespread and volcanogenic deposits are important features of the rift fill (King, 1978; Williams and Chapman, 1986). In the western branch of the rift volcanism is very limited, mainly to a small area around Lake Kivu to the north of Lake Tanganyika and the Rukwa rift (Fig. 10; Black, 1984). There are three main types of volcanism: (1) mixed-silica-saturated and undersaturated alkali basalt-phonolite-trachyte-rhyolite volcanism mainly associated with domed areas; (2) highly undersaturated K-rich and Na-rich lavas associated

THE EAST AFRICAN RIFT BASINS

195

Fig. 9. Interpretation of a section of the Eastern Rift as a series of listric faults linked to an eastward dipping detachment (after Bosworth,

1989).

Fig. 10. Volcanism in the area around the rift (after Black, 1984). with carbonatites and frequently found in parts of the rift which cut through basement rocks, and (3) silica-saturated transitional basalt-pantelleritecomendite-rhyolites linked to the formation of the Red Sea proto-passive margin (Black, 1984). In the Ethiopian rift volcanism associated with active deformation is known to have commenced between 28 and 22 my when widespread flood basalts erupted in the Afar depression. However the main period of activity began about 15 my with the eruption of ignimbrites, rhyolites and basalts which continued intermittently until at least 3 my (Mohr, 1983).

The extrusive volcanic rocks in the Eastern branch of the rift are both extensive and varied (Baker, 1986; Macdonald, 1987). It has been described as the most volcanically prolific continental rift in Phanerozoic Earth history (Karson and Curtis, 1989). It is estimated that there are approximately 500,000 km 3 of rift related volcanics in this area, over a third of which occur in Kenya (Barberi et al., 1982). Compositions range from basic through intermediate rocks and include basalts, trachytes, phonolites and mugearites. Volcanism is thought to have begun during the Oligocene and has been widespread during all subsequent stages of rift development. There has been a progressive change throughout the evolution of the Rift away from general volcanicity towards more localized eruptions within the rift valley floor. This parallels narrowing of the active fault zone through time (Karson and Curtis, 1989). Some large and active volcanoes sit outside of the rift structure, notably mounts Marsabit, Elgon, Kenya and Kilimanjaro (Fig. 11). Bosworth (1987) explains the location and occurrence of these vents as being controlled by major cross-rift structures or transfer faults. In support of his ideas it is interesting to note that the trends of several large transfer faults, when extrapolated to the east and west of the rift axis, do pass through major, ex-rift volcanoes. Natrocarbonatites are a particular feature of the East African Rift geology and they influence both the hydrology and the sedimentation in the lake basins in which they occur. One good example of this is the Lake Natron basin in northern Tanzania (Baker, 1986). Here the active volcano, O1 Doinyo Lengai, has erupted natrocarbonatite lavas and ashes consisting largely of soluble sodium-potassium car-

196

L.E. FROSTICK The subsequent filling of the basin is therefore asymmetric, with the thickest sequence adjacent to the fault and thinning onto the flexured margin. Available accommodation space increases after each bout of fault activity and then decreases as sediments attempt to fill the void. Ideally each phase of filling will be represented by a wedge-shaped deposit, the size of which depends upon the amount of accommodation generated. However, since sediment supply rarely keeps pace with subsidence and the basin is never completely filled between bouts of fault activity, the size of the depositional wedges owes more to the availability of material to fill the space than to tectonics. Sedimentary facies and depositional environments

Fig. 11. Relationship between major transfer zones and extra-rift volcanic centres (after Bosworth, 1987).

bonate minerals which are easily leached by surface waters to give sodium-rich solutions. Strontium and carbon isotope work carried out on the evaporitic deposits of Lake Natron, and other saline lakes in the Rift such as Lake Magadi, suggests that natrocarbonatites are the main source of sodium in the lake waters.

THE BASIN FILLS Evidence of the nature of the basin fills comes from a combination of seismic data, field observations on exposed stratigraphical sections and inference from present day processes and environments (Bowen and Vondra, 1973; Tallon, 1978; Frostick and Reid, 1989; Specht and Rosendahl, 1989). Patterns of sedimentation vary greatly from basin to basin and through time in response to local and regional differences in tectonic activity, volcanicity and climate (Frostick and Steel, 1993). However, the structural similarities between basins dictates some comparability at a large scale, even where the detail is dissimilar. Geometry of the fills The accommodation space available within the basin at any given time will control the geometry of the accumulating fill. The structural asymmetry which characterises the Rift dictates that subsidence is at a maximum adjacent to the faulted margin.

Clastic sediments tend to dominate rift environments, largely as a result of the topography generated by a combination of thermal uplift and block faulting both of which promote erosion and therefore increase the potential sediment supply. Two main depositional environments occur in the Rift, fluvial, in a variety of forms, and lacustrine. Aeolian activity is of only local importance and wind-blown deposits are rarely seen in the fills of those basins which have been studied in detail. Fluvial sediments In a continental rift with restricted marine influence rivers and their deposits play a vital r o l e in the architecture of the basin fill. Rivers control the locations at which both sediment and water are delivered to the basin as well as controlling the rate and timing of supply. Understanding the interplay between the evolving structure and the patterns of river drainage is of great importance in interpreting the basin fill. When present day rivers in the area surrounding the rift are analysed it becomes obvious that most of the large rivers, and many of the small ones, are diverted away from it (Fig. 2). This is clearest in the Ethiopian and Kenyan sections where thermal doming is the main cause. Doming in mantle-generated rifts may predate basin development. Such basins would be sediment-starved from their inception. In some areas of the Rift fiver diversion appears to result from the effects of footwall uplift and tilting of faulted blocks, both of which will obviously post-date fault development. Here the early stages of basin development are characterised by an influx of fluvial sediments into the evolving "sag". Footwall uplift effectively diverts all significant drainage away from the faulted margins of the basins and this restricts the development of alluvial fans. Present-day fans rarely extend more than a kilometre

THE EAST AFRICAN RIFT BASINS from the basin margin and ancient sequences show little evidence of large-scale progradation. This may be due partly to very active faulting and very rapid increases in accommodation (Blair, 1987) but the small size of the river catchments is an important contributing factor. The occurrence and character of the alluvial fans is influenced by climate. They are best developed in semi-arid areas, for example in the Suguta basin in the Kenyan Rift (Fig. 1). In wetter areas, such as that around Lake Tanganyika, the consistently high lake level favours the formation of fan-deltas (Tiercelin, 1990). Both modem and preserved fan deposits comprise coarse conglomerates and breccias deposited from debris flows, sheet floods and channelised braided streams in a seasonally, or spasmodically active, environment. Rivers draining into the Rift vary in hydrology from perennial to ephemeral with drainage basins which are generally small, rarely exceeding 600 km 2. There are a few exceptions to this, notably the Awash and Omo Rivers in Ethiopia/Kenya, and the Rusizi and Malagorasi Rivers which drain into Lake Tanganyika all of which are perennial. The smaller rivers are mostly ephemeral in character and characterise the flexured margin of the basin. They are bedload-dominated and either sinuous or braided in planform. The beds of these rivers can be either of sand or gravel and are surprisingly flat, irrespective of the calibre of the sediment, with few bedforms apart from low bars only a few centimetres high. Gravel-bed rivers are organised into coarser bars and finer fiats similar to those described for the Nahel Yatir in Israel (Laronne et al., 1994) The internal structures in the sand-bed channel fills echo the flatness of the surface with plane parallel laminations seen in almost all sections (Frostick and Reid, 1977). The few perennial rivers are very different in character, with a higher suspended load they often take on a meandering planform and produce fining upward sequences with distinctive coarser channel sediments set in finer overbank deposits characteristic of this type of river (Allen, 1978). In places where the border faulting is complex and consists of more than one parallel faults tilting on the fault blocks can trap river systems and force them to drain parallel to, but outside of, the main rift. A good example of this is the Kerio River in central Kenya which is caught between the Kamasia and Elgeyo faults and drains northwards bypassing Lakes Bogoria, Baringo and Logipi to eventually discharge into Lake Turkana through a transfer zone (Fig. 12). It is interesting to note that diverted rivers often gain access to the rift at, or close to, transfer zones. This is because they are the locations where throw on the border fault, and therefore topography, decreases to zero. As a result transfer zones are sites where thick

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sequences of fluviatile and deltaic sediments may accumulate (Frostick and Reid, 1990). Lake sediments The existence of topographic barriers at transfer zones segments the Rift into a series of hydrologically isolated basins. As a result the development of axial river systems is inhibited and runoff accumulates within the basins to form lakes. The size and depth of the lakes depends upon the hydrological balance of the basin. Of importance are the quantity and timing of rainfall, rates of evapotranspiration and the depth of the basin. The largest and deepest lake, Tanganyika, is in the western branch of the rift. It covers an area of over 40,000 km 2 and reaches a depth of 1470 m. Lakes in the Eastern branch are typically shallower and smaller, for example Lake Bogoria is less than 10 metres deep, although Lake Turkana is 250 km long by 40 km wide and is over 120 metres deep in its deepest part. The character of the lake sediments which accumulate is controlled by the location of river inputs, the quantity and calibre of clastic sediment supply, salinity and evaporation rates as well as the amount and type of organic productivity. In shallow waters rivers build up deltas which vary consider-

198

L.E. FROSTICK

Fig. 13. View westward across the River delta of the Laga Tulu Bor, Lake Turkana.

ably in character (Frostick and Reid, 1987; Fig. 13). They are generally dominated by fiver processes, but wind-wave reworking is of local significance. Sands may be reworked into beach ridges which run parallel to the shore and chart the rapid changes in Lake level in this area of strongly seasonal precipitation (Frostick and Reid, 1985). Offshore and on the steeper coastlines sudden influxes of clastic sediment can result in the development of sediment gravity flows and sublacustrine fans are recorded in the deeper basins of Lakes Tanganyika and Malawi (Le Fournier et al., 1985; Tiercelin, 1990). Organic-rich sediments occur in lakes with a variety of morphological and hydrological characteristics, ranging from deep, e.g. Tanganyika, Malawi and Kivu, to shallow, e.g. Baringo and Magadi. The prerequisites for the development and survival of organic material are high organic productivity coupled with a stratified water column. These conditions are most likely to occur in the deeper lakes but can exist in some of the shallower ones where productivity is enhanced by hydrothermal warming and enriching of the water. The main deposits are organic rich muds with high concentrations of algal remains which can give TOC levels in excess of 10%. In Tanganyika these muds are over 100 m thick in places. Diatomites also occur, particularly associated with hydrothermal activity in the Eastern Rift basins. Volcanicity can also favour the development of evaporitic sequences, especially under seasonally wet and semi-arid conditions. The deposits in the most spectacular of the evaporitic lakes, Magadi and

Natron, comprise interbedded centimetric laminae of trona (NaHCO3.Na2CO3.2H20) and wind-blown fine sands and silts. These deposits have been quarried from the lake beds since 1917. During each rain season the trona crust which has developed is broken up into "salt-bergs" and then dissolved by inflowing river waters (Fig. 14). During the next dry season the water evaporates to leave annual varves of trona approximately 2-3 mm thick (Fig. 15; Eugster, 1980). Lake Asal in the Afar depression of Ethiopia has a different chemistry as a result of its proximity to the evolving Red Sea-Gulf of Aden proto-ocean. The salinity of the waters is approximately 35% with CI, Na, K, SOn and Mg as the major ions (Loupekine, 1973). As a result the evaporites in this basin are dominantly halite and gypsum deposited in wide sabkhas. Aeolian reworking of sediments within the fill of the basins is spatially limited and relatively unimportant at the present time. Coastal sediments may be reworked into small dunes but extensive dune deposits are not characteristic of basins undergoing active faulting (Frostick and Steel, 1993). Rather they tend to develop towards the end of the active phase of rift development when subsidence slows.

Sequences and cyclicity in the fills The East African Rift basins have been accumulating sedimentary sequences since their inception in the Oligocene. Different basins have different fill histories and sequences, depending upon the timing and character of tectonic activity, fluctuations

THE EAST AFRICAN RIFT BASINS

199

Fig. 14. Lake Magadi during the dry season (A), and the rainy season (B). in hydrology and climatic change. However, some common patterns can be recognised as characteristic of particular structural and geomorphological settings in the basin. These are the faulted margin, the lake centre and the flexured margin.

The faulted margin In many basins the lake abuts the fault scarp. The topography on this margin is generally steep and only small rivers with restricted drainage basins

flow across it (Fig. 16). The classic picture of a scarp with large alluvial fans issuing from it does not apply in East Africa. Fans are small and only rarely project more than a few kilometres into the basin. The sequences developed on this margin reflect this general geomorphological picture and comprise interbedded restricted coarse fan and more extensive fine lake sediments. It is now accepted that fine lake sediments found close to a fault scarp are diagnostic of active faulting, since fans cannot prograde if

200

L.E. FROSTICK

Fig. 15. Evaporitic deposits of Trona on the surface of Lake Magadi.

accommodation is generated faster than it can be filled with sediment by the supplying rivers (Blair, 1987). Rises and falls in lake level in response to climatic fluctuations have only a limited affect on sedimentation at this margin. A small scale rise of a few metres will hardly move the shoreline and will be undetectable in the succession, a larger change can shift fine lake sediments closer to the fault scarp thus mimicking the effects of faulting. Lake centre The deposits which accumulate in the lakes of the rift vary somewhat with the hydrology and chemistry of the basin (see previous section). However laminated deposits are, in all cases, characteristic with the laminae or varves resulting from seasonal changes in rainfall, evaporation and organic productivity. In Lake Tanganyika, deposits comprise either green detrital muds with interbeds of diatom gels or black organic rich alternating with grey, more clayrich, layers. In Lake Turkana the laminations are thin beds which fine upwards, each bed representing the fall out from a single rain season's programme of flash floods.

Fig. 16. River drainage in the area around Lake Tanganyika (after Frostick and Reid, 1987).

The flexured margin This margin is characterised by the most complex sequences. Both tectonic and climatic fluctuations will change the distribution and character of the sediments deposited by altering both fiver gradients and lake depths. In Lake Turkana the sedimentary

THE EAST AFRICAN RIFT BASINS

201

Fig. 18. Lake level curves for the past 14,000 years for four rift lakes (after Frostick and Reid, 1989). this time in response to the changing water balance (Frostick and Reid, 1989; Fig. 18).

Examples of basin types

Fig. 17. Beach ridges around Kokoi on the eastern shore of Lake Turkana (after Frostick and Reid, 1985).

consequences of recent fluctuations in lake level are clearly visible on air photographs in the form of a flight of abandoned beach deposits now stranded over 100 m above the present lake shore (Fig. 17; Frostick and Reid, 1985). The hall mark of flexured margin sequences is successive phases of river/delta progradation, valley incision and then filling. A rising lake shifts the position of shoreline and deep lake environments landwards, particularly in the sections of the basin with lower gradients, resulting in the superposition of lake and shoreline sediments on fluvial deposits and progressive coastal onlap. As the water level falls fluvial systems will cut down into underlying lake sediments and prograde into the basin depositing a "lowstand fan" (sensu Vail et al., 1977) with a coarsening upward signature. Aggradation is rarely a consequence of base-level fall, contrary to the ideas of Posamentier and Vail (1988). All of the sequences on this margin c a n be analysed using sequence stratigraphic concepts suggested for fluvial deposits by Posamentier and Vail, however the criticisms levelled by Miall (1991) should be borne in mind. Cyclicity occurs at a variety of scales as a result of fluctuations in weather and climate as well as pulses of tectonic activity. Much of the fill of the basins is Plio-Pleistocene in age, a time of rapid and extreme climatic change. Many of the lakes show evidence of large scale rises and falls during

There is a great deal of variety in basin character in the East African Rift. They range from broad, shallow structurally complex and volcanic rich to narrow, deep and structurally simple with little volcanism. The segments of the rift containing lakes Tanganyika, Turkana and Baringo-NaivashaMagadi have been selected for detailed description as representatives of the continuum of basin types.

The Lake Tanganyika basin The lake occupies a narrow faulted trough which varies between 50 and 80 km wide. It is the largest of the rift lakes and is divided structurally into 6 subbasins each approximately 100 km long. The river catchment area is relatively small when compared to the volume of the lake. Only approximately 160,000 km 2 drains into Tanganyika and over 30% of this is the Malagorasi river and its associated tributaries. All of the sub-basins are seen to be asymmetric on seismic sections (Rosendahl, 1987; Fig. 19) with main border faults which alternate on either side of the rift axis. The sedimentary fill is wedge shaped, thinning and onlapping onto the flexured margin. The Malagorasi excepted, only small rivers cross the flexured margin of the basin and drainage diversion away from the fault scarps is very marked (Fig. 16). In addition the rift at this point cuts largely through crystalline basement which is slow to weather and difficult to erode. As a result clastic sediment supply is restricted over most of the lake and this encourages the accumulation of both stromatolites and shell banks along its shores. Rates of sedimentation in deeper waters are slow (less than 0.5 mm yr -1) and persistent thermal stratification at more than 70 m below surface combines with anoxic waters at depth to allow fine-grained, organic-rich muds to accumulate over much of the basin (Hecky and Degens, 1973). Productivity is high as a result of

202

L.E. FROSTICK is due to the larger quantities of sediment supplied by the Malagorasi and Ruzizi rivers both of which have built significant delta deposits. The narrowness and depth of the basin have important consequences for the development of sequences on the flexured margin. The relatively steep gradients on this margin allow only limited migration of the lake shore as water level rises and falls. This makes it more difficult to recognise previous high lake stages and also simplifies the sequences which develop (Fig. 20). Similar patterns of sedimentary fill are found in the other deep basins of the western rift, notably Lakes Malawi (695 m) and Kivu (500 m). Differences arise largely from the distribution of river inputs and hydrothermal activity.

Fig. 19. Structure of the subbasins of Lake Tanganyika as interpreted from seismic sections (after Rosendahl, 1987). both the warm tropical setting and sublacustrine hydrothermal activity which adds nutrients to the waters. The planktonic flora is characterised by a few species, notably diatoms (Talbot, 1988), and diatom gels are found interbedded with the black muds. The northern-most sub-basin is the only one where clastic sediment dominates both the flexured margin and the deeper lake environment. This

The Lake Turkana basin Turkana is the largest of the lakes in the Eastern Rift but is still less than 120 m deep at its deepest point. It is a saline lake which, due to its shallowness, is generally unstratified. It occupies parts of 3 major structural subbasins the most northern of which is complicated by much small scale faulting (Fig. 21). The structural basins spread far beyond the limits of the lake on the flexured margins and average over 90 km wide by 75 km long. The overall impression is of a broader and slightly more complex structure than that of Lake Tanganyika. Basin width and length have both been influenced by the rise of a series of volcanic centres, especially to the south of the lake. Volcanism also plays an important role in the sedimentological development of the basin. Widespread basalts and tufts are present in the Tertiary and Quaternary sequences and these act as chronostratigraphic markers (Bowen and Vondra,

Fig. 20. Schematic diagram showing a hypothetical cross-section through the Tanganyikarift with details of the sedimentary fill.

THE EAST AFRICAN RIFT BASINS

203

I

Fig. 21. Structure and cross-sections across the subbasins of the Lake Turkana rift (after Frostick and Reid, 1990).

ronment favourable to the preservation of terrestrial fossils and these sequences are rich in the remains of vertebrate animals, including hominids (Leakey and Leakey, 1978). The faulted margin of the basin is crossed by only a few small rivers which carry limited amounts of coarse detritus. Most of the rivers drain away from the basin depocentre as a result of footwall uplift and tilting. Some are caught between two parallel faults in the footwall which create small saddle-type basins (Fig. 24). Fans are small and during periods of active fault movement fine lake sediments abut the scarp. Lake sediments are generally dominated by clastic material and have organic contents of less than 1%. Turkana is typical of the moderately deep lakes in volcanically active settings. Few of the other lakes in the Eastern rift are similar in character since most achieve maximum depths of less than 10 m (e.g. Bogoria is 9 m, Baringo only just over 4 m). Similarly all of those in the Ethiopian rift are either shallow or ephemeral (e.g. Chew Bahir). However the smaller lakes of the Western rift, Mobutu and Rutanzigel, are of moderate depth and well-mixed like Turkana. These are likely to have similar patterns of sedimentation.

1973, Cerling and Brown, 1982; Fig. 22). There are distinct volcanic vents some of which act as topographical barriers to sediment transfer, for example North, South and Central Islands are all volcanic, as are mounts Kulal and Teleki. Interestingly a number of the vents are associated with transfer zones. Volcanism has important implications for sedimentation in the basin. Volcanic rocks weather rapidly yielding large quantities of debris which ranges in clast size from boulders to clays (Frostick and Reid, 1980; Cohen et al., 1986; Frostick and Reid, 1986). This contrasts with areas of the rift where the only source of sediment is slow-weathering crystalline basement and rates of sediment supply are low and dominated by quartz sand (e.g. Tanganyika). The Turkana rift, and basins like it, are therefore dominated by the deposition of clastic sedimentary sequences. Carbonate and organic rich lake sediments are rare with the exception of localised stromatolitic and diatomitic bands, e.g. those reported by Johnson (1974). The flexured margins of the Lake Turkana subbasins are less steep than the equivalents in Lake Tanganyika. As a result changes in lake level have produced rapid and extensive transgressions and regressions which complicate the stratigraphy (Fig. 23). Interbedding of fluvial, deltaic and littoral deposits at metre to decimetre scale characterise the Plio-Pleistocene sequences of the flexured margin. Rapid transgression and burial produces an envi-

This 400 km long section of the Eastern rift is one where small lakes cover only a negligible proportion of the basin area and do not necessarily abut the main border faults. There are 3 subbasins in this area: the Baringo-Bogoria, Naivasha-Nakuru and Magadi-Natron basins. Each of the subbasins is approximately 100 km wide by 125 km long with a major border fault on either the western (Baringo-Bogoria and Magadi-Natron) or eastern (Naivasha-Nakuru) margin. The lakes are shallow and generally unstratified and are often separated from the main border fault by a series of minor antithetic and synthetic faults. The structure of the rift is broad and open with a very complex pattern of faulting which controls the river drainage patterns of the area (Fig. 12). The basins are compartmentalized across their widths by a number of tilted fault blocks which limit the transfer of sediment in directions orthogonal to the general rift axis. They cut the rift into a series of small and shallow sections each with its own depositional history. Thick and extensive volcanic deposits characterise this section of the rift and the weathering of volcanic rocks provides a major source of the clastic sediments being deposited in the basin as well as affecting lake chemistry and the character and extent of evaporitic sediments. Stratigraphic evidence suggests that outpourings of volcanic debris have often filled the small amounts of accommodation available, for example in the Baringo area there

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L.E. F R O S T I C K

Fig. 22. Satellite image of the Eastern Rift showing extensive basalts and basaltic sediments (black areas, especially to the east of lake Turkana which is situated in the central north of the photograph).

Fig. 23. Schematic diagram showing a hypothetical cross-section through the Turkana rift with details of the sedimentary fill (after Frostick and Reid, 1990).

THE EAST AFRICAN RIFT BASINS

205

Fig. 24. Print lay-down for the area around Fergusons Gulf, Lake Turkana. The river draining southwards (from top to bottom) is caught behind the fault scarp of the Katulenyang Hills and only cuts eastward towards the lake at a transfer zone.

Fig. 25. A. General structure of the Baringo-Bogoria-Magadi basin. B. Schematic diagram showing a hypothetical cross-section through the Baringo rift with details of the sedimentary fill (after Frostick and Reid, 1990).

are sequences of lava flows over a kilometre thick (Fig. 25). Sedimentary deposits are often of subsidiary importance. They comprise small stringers of shallow-water, sometimes ephemeral and evaporitic, lake deposits in sandy or finer alluvium (e.g. the Miocene Ngorora Formation, Pickford, 1978). The fine nature of the fiver sediments is due to low

local gradients and the small size of the drainage networks. Diatomites may be deposited in the lake, encouraged by the introduction of silica from volcanic sources. Similar patterns of sedimentation exist in the rift segments containing the shallow lakes of Ethiopia (including Chamo, Abbeya, Abbe, Awasa, Shala,

206 Ziway and Chew Bahir), as well as in other areas of the Eastern rift (e.g. around Lakes Eyasi and Manyara) the main variations depending on the degree of ephemerality and the salinity of the lake, the amount and character of volcanism and the number of antithetic and synthetic faults.

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The sedimentary basins of the East African Rift vary greatly in structural style and fill-character. The main contrast is between the deep and structurally simple basins of the Western branch of the rift and the shallower and more complex nature of those in the Eastern and Ethiopian sections. The amount of volcanicity appears to correlate with the complexity of faulting. Areas where the faults are many but small are the foci of volcanic activity, for example the Baringo area of Kenya, whereas the Lake Tanganyika basin, which has a relatively simple structure, has experienced only limited volcanic activity. Considerable research has been carried out into the underlying controls on the variations in basin structure and volcanicity in different parts of the Rift. One possible control is the location of mantle plumes and the associated doming (Cerling and Powers, 1977), the domed areas having high heat flow and acting as centres for volcanic activity (Fairhead and Stuart, 1983). Fairhead (1986) has speculated that the degree of crustal thinning may exert an important control, the thinnest crust being in the domed and volcanic areas. Kuznir and Park (1987) have suggested that rates of strain play an important role in controlling the style of faulting in rifts. They consider that high rates help to maintain crustal ductility and tend to localise brittle failure into the development of few but large faults. The resulting structure looks very similar to the Tanganyika type of basin. In contrast, areas with lower strain rates have a more brittle crust which fails in more places giving many but small faults which develop over a wider area. This is a Baringo-Bogoria type of basin. They also speculate that crustal thickness may control strain rates and that this is therefore the fundamental control on structure. It appears that the conditions in the crust prior to faulting are the critical factor in controlling the variety of basin type. There is a spectrum of basins which range from simple to complex depending on the character of the pre-rift crust and differences which developed during incipient rifting due to localised doming and high heat flow (Fig. 26).

x

Fig. 26. Schematic diagram showing the relationships between basin-fill type, structure, volcanism, rates of strain and crustal thickness.

I m p o r t a n c e o f the E a s t A f r i c a n Rift b a s i n s

Continental rifts offer the conditions favourable to the development of a number of strata-bound economic deposits which are rare in the cratonic areas beyond the rift basin margins. Some of the basins contain sediments which can both produce and trap hydrocarbons in economic quantities, given the right burial history. The best source rocks occur in Tanganyika-type lake deposits, which also contain littoral deposits which are suitable reservoirs. Evaporitic deposits are extracted as sources of industrial minerals and thick sand and gravel deposits occur in many of the basins. These will be economic if the local demand for building aggregates is sufficient. In addition littoral processes can lead to the concentration of heavy minerals, particularly metals, in beach deposits (Frostick and Reid, 1985). The area has yielded important geological and geophysical data which has contributed significantly to our understanding of rift dynamics. These data give us important insights into crustal processes as a continent begins to split apart and add greatly to our understanding of plate tectonics. Many of the models used to interpret ancient rift sequences are based upon ideas originally developed in East Africa. The East African Rift basins are of academic as well as economic significance. They will therefore remain the targets for research far into the foreseeable future.

ACKNOWLEDGEMENTS My work in East Africa was carried out whilst in receipt of a NERC grant held jointly with Professor

T H E E A S T A F R I C A N RIFT BASINS

207

I a n R e i d . H e h a s b e e n a h e l p a n d i n s p i r a t i o n to m e o v e r t h e p a s t 20 y e a r s a n d I a m v e r y g r a t e f u l to h i m . I w o u l d a l s o like to t h a n k m y f o r m e r colleague Professor Ken McClay for many enjoyable d i s c u s s i o n s a b o u t rift s t r u c t u r e a n d d e v e l o p m e n t . I o w e a s p e c i a l d e b t o f g r a t i t u d e to R i c h a r d L e a k e y , whose

help

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REFERENCES

Bailey D.K. 1972. Uplift, rifting and magmatism in continental plates. J. Earth Sci., 8: 225-239. Barberi, E, Santacroce, R and Varet, J. 1982. Chemical aspects of rift magmatism. In: G. Palmason (Editor), Continental and Oceanic Rifts. Am. Geophys. Union, Geodyn. Ser. A, 223258. Baker, B.H., 1986. Tectonics and volcanism of the southern Kenya Rift Valley and its influence on rift sedimentation. In: L.E. Frostick, R.W. Renaut, I. Reid and J.-J. Tiercelin (Edtiors), Sedimentation in the African Rift. Spec. Publ. Geol. Soc., 25: 45-58. Bally, A.W., 1982. Musing over sedimentary basin evolution. Phil. Trans. R. Soc. London, Ser. A, 305: 325-328. Black, R., 1984. The Pan-African event in the geological framework of Africa. Pangea, 2: 6-16. Blair, T.C., 1987. Tectonic and hydrologic controls on cyclic alluvial fan, fluvial and lacustrine rift-basin sedimentation, Jurassic-lowermost Cretaceous Totos Santos Formation, Chiapas, Mexico. J. Sediment. Petrol., 57: 845-62. Bosworth, W., 1987. Off-axis volcanism in the Gregory Rift, East Africa: implications for models of continental rifting. Geology, 15: 397-400. Bosworth, W., 1989. Basin and range style tectonics in East Africa. J. Afr. Earth Sci., 8: 191-202. Bowen, B.E. and Vondra, C.F., 1973. Stratigraphicai relationships of the Plio-Pleistocene deposits, East Rudolf, Kenya. Nature, 242:391-393. Burgis, M. and Gaudet, J.J., 1981. The Ecology and Utilization of African Inland Waters. United Nations Environment Programme, Nairobi. Butzer, K.A., 1980. The Holocene lake plain of North Rudolf, East Africa. Phys. Geogr., 1: 42-58. Cerling, C.E. and Brown, EH., 1982. Tuffaceous marker horizons in the Koobi Fora region and the lower Omo valley. Nature, 299:216-221. Cerling, T.E. and Powers, P.W., 1977. Palaeorifting between the Gregory and Ethiopian Rifts. Geology, 5:44 1-444. Chapman, G.R., Lippard, S.J. and Martyn, J.E., 1978. The stratigraphy and structure of the Kamasia range, Kenya Rift Valley. J. Geol. Soc., 135: 265-281. Chen, W.P. and Molnar, P., 1983. Focal depths of intracontinental and intra-plate earthquakes and their implications for the thermal and mechanical properties of the lithosphere. J. Geophys. Res., 88:4183-4214. Chorowitz, J., 1989. Transfer and transform fault zones in continental rifts: examples in the Afro-Arabian Rift System. Implications of crust breaking. J. Afr. Earth Sci., 8:203-214. Chorowitz, J., 1990. Dynamics of different basin-types in the East African Rift. J. Afr. Earth Sci., 10:271-282. Coppens, Y., Howell, EC., Isaac, G.L. and Leakey, R.E.E 1976. Earliest Man and Environments in the Lake Rudolf Basin. University of Chicago Press, Chicago. Cohen, A., Ferguson, D.S., Gram, P.H., Hubler, S.L. and Sims, K.W., 1986. The distribution of coarse-grained sediments in

modern Lake Turkana, Kenya: implications for clastic sedimentation models of rift lakes. In: L.E. Frostick, R.W. Renaut, I. Reid and J.J. Tiercelin (Editors), Sedimentation in the African Rifts. Geol. Soc. Spec. Publ., 25: 127-139. Eugster, H.P., 1980. Lake Magadi, Kenya and its precursors. In: A. Nissenbaum (Editor), Hypersaline Brines and Evaporites. Dev. Sedimentol., 28: 195-232. Eugster, H.P., 1986. Lake Magadi, Kenya: a model for rift valley hydrochemistry and sedimentation? In: L.E. Frostick, R.W. Renaut, I. Reid and J.J. Tiercelin (Editors), Sedimentation in the African Rift. Spec. Publ. Geol. Soc., 25: 177-189. Fairhead, J.D., 1986. Geophysical controls on sedimentation within the African Rift System. In: L.E. Frostick, R.W. Renaut, I. Reid and J.J. Tiercelin (Editors), Sedimentation in the African Rift. Spec. Publ. Geol. Soc., 25. Fairhead, J.D. and Green, C.M., 1989. Controls on rifting in Africa and the regional tectonic model for the Niger and East Niger rift basins. J. Afr. Earth Sci., 8:231-250. Fairhead, J.D. and Henderson, N.B., 1977. The seismicity of southern Africa and incipient rifting. Tectonophysics, 41: 1926.

Fairhead, J.D. and Reeves, C.V., 1977. Teleseismic delay times, Bouguer anomolies and inferred thickness of the African lithosphere. Earth Planet. Sci. Lett., 36: 63-76. Fairhead, J.D. and Stuart, G.W., 1983. The seismicity of the African Rift System in comparison with other continental rifts. Geodyn. Ser. A, 8: 41-61. Frostick, L.E. and Reid, I., 1977. The origin of horizontal laminae in ephemeral stream channel-fill. Sedimentology, 24: 19. Frostick, L.E. and Reid, I., 1980. Sorting mechanisms in coarsegrained alluvial sediment:fresh evidence from a basalt plateau gravel, Kenya. J. Geol. Soc., 137:431-441. Frostick, L.E. and Reid, I., 1985. Beach orientation, bar morphology and the concentration of metalliferous placer deposits: a case study, lake Turkana, Northern Kenya. J. Geol. Soc., 142: 837-848. Frostick, L.E. and Reid, I., 1986. Evolution and sedimentary character of lake deltas fed by ephemeral rivers in the Turkana basin, northern Kenya. In: L.E. Frostick, R.W. Renaut, I. Reid and J.J. Tiercelin (Editors), Sedimentation in the African Rifts. Geol. Soc. Spec. Publ., 25:113-125. Frostick, L.E. and Reid, I., 1987. Tectonic control of desert sediments in rift basins ancient and modern. In: L.E. Frostick and I. Reid (Editors), Desert Sediments: Ancient and Modern. Geol. Soc. Spec. Publ., 35: 53-68. Frostick, L.E. and Reid, I., 1989. Is structure the main control on river drainage and sedimentation in rifts? J. Afr. Earth Sci., 8: 165-182. Frostick, L.E. and Reid, I., 1990. Structural control of sedimentation patterns and implications for the economic potential of the East African Rift basins. J. Afr. Earth Sci., 10: 307-318. Frostick, L.E. and Steel, R.J. 1993. Sedimentation in divergent plate-margin basins. In: L.E. Frostick and R.J. Steel (Editors), Tectonic Controls and Signatures in Sedimentary Successions. Int. Assoc. Sedimentol. Spec. Publ., 20:111-128. Gawthorpe, R.L. and Hurst, J.M., 1993. Transfer zones in extensional basins: their structural style and influence on drainage development and stratigraphy. J. Geol. Soc., 150:1137-1152. Gibbs, A.D., 1984. Structural evolution of extensional basin margins. J. Geol. Soc., 141: 609-620. Grove, A.T., 1986. Geomorphology of the African Rift System. In: L.E. Frostick, R.W. Renaut, I. Reid and J.J. Tiercelin (Editors), Sedimentation in the African Rifts. Geol. Soc. Spec. Publ., 25: 9-16. Grove, A.T., Street, EA. and Goudie, A.S., 1975. Former lake levels and climatic change in the Rift Valley of Southern Ethiopia. Geogr. J., 141: 177-202.

208 Haberyan, K.A. and Hecky, R.E., 1987. The Late Pleistocene and Holocene stratigraphy and palaeolimnology of lakes Kivu and Tanganyika. Paleogeogr. Paleoclimatol. Paleoecol., 61: 169197. Hecky, R.E. and Degens, E.T., 1973. Late Pleistocene-Holocene chemical stratigraphy and palaeolimnology of the rift valley lakes of central Africa. Woods Hole Oceanogr. Inst., Tech. Rep., pp. 73-28. Hohnel, L. Von, 1894. Discovery of Lakes Rudolf and Stephanie, 397 pp. Johnson, G.D., 1974. Cainozoic lacustrine stromatolites from hominid-bearing sediments east of Lake Rudolf, Kenya. Nature, 247: 520-523. Johnson, T.C., Halfman, J.D., Rosendahl, B.R. and Lister, G.S., 1987. Climatic and tectonic effects on sedimentation in a rift valley lake: evidence from high resolution seismic profiles. Bull. Geol. Soc. Am., 98: 439-447. Karson, J.A. and Curtis, P.C., 1989. Tectonic and magmatic processes in the Eastern Branch of the East African Rift and implications for magmatically active continental rifts. J. Afr. Earth Sci., 8:431-454. Khan, A., Maguire, P., Henry, B. and Higham, M., 1986. KRISP86 ~ An international seismic investigation of the Kenya rift. Geol. Today, 139-144. Kreuser, T., Wopfner, H., Kaaya, C.Z., Markwort, S., Semkiwa, P.M. and Aslanidis, P., 1990. Depositional evolution of PermoTriassic Karoo basins in Tanzania with reference to their economic potential. J. Afr. Earth Sci., 10: 151-167. King, B.C. 1978. Structural and volcanic evolution of the Gregory Rift Valley. In: W.W. Bishop (Editor), Geological Background to Fossil Man. Scottish Academic Press, Edinburgh, pp. 29-54. Kuznir, N.J. and Park, R.G., 1987. The extensional strength of the continental lithosphere: its dependence on geothermal gradient, and crustal composition and thickness. In: M.P. Coward, J.F, Dewey and P.L. Hancock (Editors), Continental Extensional Tectonics. Geol. Soc. Spec. Publ., 28: 35-52. Lambiase, J.J., 1989. The framework of African rifting during the Phanerozoic. J. Afr. Earth Sci., 8: 183-190. Laronne, J.B., Reid, I., Yitshak, Y. and Frostick, L.E., 1994. The non-layering of gravel streambeds under ephemeral flood regimes. J. Hydrol., 159: 353-363. Larsen, P.H. 1988. Relay structures in a Lower Permian basement-involved extensional system, East Greenland. J. Struct. Geol., 10: 3-8. Leakey, R.E.F. and Leakey, M.D. 1970. New hominid remains and early artifacts from Northern Kenya. Nature, 226: 223. Leakey, R.E.F., 1973. Evidence for an advanced Plio-Pleistocene hominid from Lake Rudolf, Kenya. Nature, 242: 447-450. Leakey, R.E.F., 1974. Further evidence of the Lower Pleistocene hominids from East Rudolf, Northern Kenya. Nature, 248: 653-656. Leakey, M.G. and Leakey, R.E., 1978. Koobi Fora Research Project 1. The Fossil Hominids and an Introduction to their Context. Clarendon Press, Oxford, 191 pp. Loupekine, I.S., 1971. R61e de la g6ochimie dans le recherche d'6nergie g6othermique. Application au T.F.A.I. Th~se 3i~me cycle, Universite de Paris, Paris. Macdonald, R., 1987. Quaternary peralkaline silicic rocks and caldera volcanoes of Kenya. In: J.G. Fitton and B.G.J. Upton (Editors), Alkaline Igneous Rocks. Geol. Soc. Spec. Publ., 30: 313-333. McKenzie, D.P., 1978. Some remarks on the development of sedimentary basins. Earth Planet. Sci. Lett., 40: 25-32. McKenzie, D.P. and Jackson J. 1986. A block model of distributed deformation by faulting. J. Geol. Soc., 143: 349-353. McKenzie, D.P., Davies, D., and Molnar, P., 1970. Plate tectonics of the Red Sea and East Africa. Nature, 226: 243-248.

L.E. F R O S T I C K Maguire, P.K.H. and Long, R.E., 1976. The structure of the western flank of the Gregory rift (Kenya), part 1: the crust. Geophys. J. R. Astron. Soc., 661-675. Miall, A.D., 1991. Stratigraphic sequences and their chronostratigraphic correlation. J. Sediment. Petrol., 71: 497-505. Mohr, P., 1983. Volcanotectonic aspects of Ethiopian rift evolution. Bull. Centr. Rech. Explor.-Prod. Elf Aquitaine, 7: 175189. Morley, C.K., Patton, T.L. and Munn, S.G. 1990. Transfer zones in the East African Rift System and their relevance to hydrocarbon exploration in rifts. Am. Assoc. Pet. Geol. Bull., 74: 1234-1253. Oxburgh, E.R., 1978. Rifting in East Africa and large scale tectonic processes. In: W.W. Bishop (Editor), Geological Background to Fossil Man. Scottish Academic Press and University of Toronto Press, pp. 7-18. Peacock, D.C.P. and Sanderson, D.J., 1991. Displacements, segment linkage and relay ramps in normal fault zones. J. Struct. Geol., 13: 721-733. Posamentier, H.W. and Vail, P.R., 1988. Eustatic controls on clastic deposition, II m sequences and systems tract models. In: C.K. Wilgus, B.S. Hastings, C.G. Kendall, H.W. Posamentier, C.A. Ross and J.C. Wagoner (Editors), Sea-Level Research: an Integrated Approach. SEPM Spec. Publ., 42: 125-154. Roberts, G.P., Gawthorpe, R.L. and Stewart, I., 1993. Surface faulting within active normal fault zones: examples from the Gulf of Corinth fault system, central Greece. Z. Geomorphol. Rogers, J.J.W. and Rosendahl, B.R., 1989. Perceptions and issues in continental rifting. J. Afr. Earth Sci., 8: 137-142. Rosendahl, B.R., 1987. Architecture of continental rifts with special reference to East Africa. Ann. Rev. Earth Planet. Sci., 15: 445-503. Rosendahl, B.R. and Livingstone, D.A., 1983. Rift lakes of East Africa: new seismic data and implications for future research. Episodes, 1: 14-19. Sander, S. and Rosendahl, B.R., 1989. The geometry of rifting in Lake Tanganyika, East Africa. J. Afr. Earth Sci., 8: 323-354. Searle, R.C., 1970. Evidence from gravity anomolies for thinning of the lithosphere beneath the rift valley of Kenya. Geophys. J. R. Astron. Soc., 21 : 13-31. Specht, T.D. and Rosendahl, B.R., 1989. Architecture of the Lake Malawi Rift, East Africa. J. Afr. Earth Sci., 8: 323-354. Talbot, M.R., 1988. The origins of lacustrine oil source rocks: evidence from lakes in tropical Africa. In: A.J. Fleet, K. Kelts and M.R. Talbot (Editors), Lacustrine Petroleum Source Rocks. Geol. Soc. Spec. Publ., 40: 29-43. Tallon, P.W.J., 1978. Geological setting of the hominid fossils and Acheulian artifacts from the Kapthurin Formation, Baringo district, Kenya. In: W.W. Bishop (Editor), Geological Background to Fossil Man. Scottish Academic Press, Edinburgh, pp. 361-374. Thomson, J., 1887. Through Masai Land. 364 pp. Tiercelin, J.J., 1990. Rift basin sedimentation: responses to climate, tectonism, and volcanism. Examples of the East African Rift. J. Afr. Earth Sci., 10: 283-305. Truckle, P.H., 1976. Geology and late Cainozoic lake sediments of the Suguta Trough, Kenya. Nature, 263: 380-383. Vail, P.R., Mitchum, R.M., Todd, R.G., Widmier, J.M., Thomson, S., Sangree, J.B., Bubb, J.N. and Hatlelid, W.G., 1977. Seismic stratigraphy and global changes in sea-level. In: C.E. Payton (Editor), Seismic Stratigraphy m Applications to Hydrocarbon Exploration. AAPG Mem., 26:49-212. Vincens, A. and Casanova, J., 1987. Modern background of Natron-Magadi basin (Tanzania-Kenya): Physiography, climate, hydrology and vegetation. Sci. Geol. Bull., 40:9-21. Walsh, J.J. and Watterson, J. 1987. Distributions of cumulative displacement and seismic slip on a single normal fault surface. J. Struct. Geol., 9:1039-1046.

T H E EAST A F R I C A N RIFT BASINS Williams, L.A.J. and Chapman, G.R., 1986. Relationships between major structures, salic volcanism and sedimentation in the kenya Rift from the equator northwards to Lake Turkana.

209 In: L.E. Frostick, R.W. Renaut, I. Reid and J.J. Tiercelin (Editors), Sedimentation in the African Rifts. Geol. Soc. Spec. Publ., 25: 59-74.

Chapter 10

The Coastal Basins of Somalia, Kenya and Tanzania

E.I. MBEDE and A. DUALEH

INTRODUCTION

The western Indian Ocean seaboard is an Atlantic type of continental margin, hence the sedimentary basins involved are typical pull-apart basins of Klemme (1980). The margin has, however been subjected to transform movements through Late Jurassic to Late Cretaceous times when Madagascar was moving southwards relative to Africa. It is thus referred to as a transform continental margin by, among others, Bosellin (1986) and Mascle et al. (1987). The margin seems to have been an emergent and stable block during the whole of Palaeozoic time. Sedimentation starts on top of a peneplaned Precambrian basement surface made up of highly metamorphosed rocks. These crop out to the west of coastal sediments in what is called the Mozambiquan belt in Tanzania and Kenya (Figs. 6 and 9). Further north, in Somalia, basement rocks crop out as oval shaped areas in southern Somalia and along the northern Somali main escarpment (Fig. 2). Here they contain low grade (Inda series) and medium to high grade (Old Formation) metamorphic rocks related to the Pan-African tectonothermal event that terminated with the intrusion of granites, granodiorites, synites and abundant dikes (D'Amico et al., 1982). Basement exposures along the northern Somali coast, and along the escarpment are attributed to the separation of Arabia from Africa during the Miocene. Whereas exposures in the Bur-Acaba and Nogal areas are considered to be reactivated Mesozoic structures, similar to other structural uplifts recorded further south. At least four sedimentary cycles, each of which was probably associated with a major tectonic event, have been recorded on the east African margin during Phanerozoic time. First is the precursory tectonism of the opening of Indian Ocean and the fragmentation of East Gondwanaland (Late Carboniferous to Early Permian). This lead to the formation of continental rifts which subsequently hosted continental deposits. Marine connections during this time appear

at different ages in Tanzania, Kenya and Somalia. These reflect physical limitations to marine incursion in the isolated fault separated restricted troughs of the proto "Malgash Gulf". Salt diapirs in Kenya and Somalia, and the Mandawa salt basin of Tanzania, are considered to be of this cycle. The second phase started with the deposition of basal arenaceous sandstone unit overlain by marine beds. This phase is related to the major faulting phase during the Early Jurassic, which indicates reactivation of basement before the deposition of the arkosic Ngerengere Beds in Tanzania and their equivalents, the Manzeras Sandstones in Kenya, and the Adigrat Formation in Somalia and Ethiopia. A strong marine transgression, which probably culminated during the Middle Jurassic (Bathonian), affected the whole of East Africa. Wholly marine conditions were already established in N.E. Kenya and Somalia by the Early Jurassic, whereas further south, in SE Kenya and Tanzania, this did not take place until the Middle Jurassic. This indicates that the sea was slowly encroaching the margin from the north and east, where open marine conditions already existed (Fig. 13). The southward drift of Madagascar relative to Africa occurred along a short spreading centre and along the Davie fracture zone, a transform fault believed to have been active from Late Jurassic (156 m.y.) to Early Cretaceous (130 m.y.) time. A Late Jurassic transgression was associated with tectonism related to this event. Another major transgression flooded most of East Africa in Aptian and Late Palaeocene to Early Eocene times. These two event could be related to the development of the Owen fracture zone and to the widening of the Indian Ocean at the close of Early Cretaceous and Early Tertiary to Late Eocene times. The fourth cycle includes tectonism related to the drift of southern Arabia away from Africa (northern Somalia) and the development of the East African rift system during Oligocene and Miocene times. This resulted in the total withdrawal of the sea

African Basins. Sedimentary Basins of the World, 3 edited by R.C. Selley (Series Editor: K.J. Hsti), pp. 211-233. 9 1997 Elsevier Science B.V., Amsterdam. All rights reserved.

212

E.I. MBEDE and A. DUALEH

from the present areas of northern Somalia (pre-drift doming), and in the reactivation of sedimentation in the areas bordering the Indian Ocean coastal belt. It also caused the formation of the present Gulf of Aden and the main physiographic features of northern Somalia. A Late Eocene to Oligocene regression occurs along the east African margin. This paper attempts to synthesize the geology and stratigraphic evolution of the coastal basins of the east Africa margin (Fig. 1). The principal data used include well logs, geophysical, geochemical data and literature available. The geology of each country is discussed separately so as to give emphasis to the local stratigraphic variations within the region, even though the margin evolved as one unit. The structural evolution is then looked at as one unit, and finally there is a section on economic considerations. Besides gas discovered in Tanzania, no commercial petroleum reserves have been reported so far in the region. Oil and gas shows have been reported everywhere, and the basin is considered to be a low potential gas prone province (Chatellier and Slevin, 1988), but we think that the conclusion is too premature 0~

10 ~

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with the present intensity of data available. Other geological resources, including gypsum, common salt, kaolinite, limestones and other building materials, are exploited at the moment, while heavy mineral beach sands, as well as palaeo-placers are reported to be abundant. The basins discussed in this paper include only those bordering the present Indian Ocean. They include the Somali Embayment, the Somali Coastal Basin, the Luug-Mandera Basin, the Kenya Coastal Basin, the Selous-Ruvu-Tanga Basin and Lindi Rift Basin to the south. These basins are separated by basement highs, either cropping out, or concealed beneath a thin sedimentary cover.

REVIEW OF THE GEOLOGY OF THE SOMALI COASTAL BASIN Introduction

The first epeirogenic movement to affect the region formed a series of intersecting basins separated by structural highs. The latter include the 40 ~

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Bur-Acaba uplift of southern Somalia, and the Nogal and Hargesya-Ergavo uplifts of northern Somalia (Fig. 2). These highs formed the boundaries of the

present Somali sub-basins, and controlled the pattern of sedimentation throughout. Two main sedimentary basins, separated by a basement ridge (the Bur-

214

E.I. MBEDE and A. DUALEH

Acaba uplift), occur in southern Somalia. These are the Mesozoic Luug Mandera Basin, which trends N N E - S S W and the Somali Coastal Basin trending parallel to the Indian Ocean coast. This contains both Mesozoic and Tertiary sediments. The northern part of the Somali Coastal Basin is overprinted by the E - W elongated Somali Embayment, which also contains both Mesozoic and Tertiary sediments. A tentative correlative chart of these three basins is shown in Fig. 5. Pre-Jurassic rock do not crop out in southern Somalia and none of the wells drilled in it reaches the basement. The older rock penetrated are of Triassic to Early Jurassic age, and were recorded in Brava-1 well drilled in the coastal basin (Fig. 2). The Hol-1 well, drilled at the axial part of Luug-Mandera Basin, bottomed in a fluvio-deltaic sequence of Hettangian-Toarcian age. Marine sedimentary rocks predating the first wide spread regional transgression were recorded in the Obbia-1 well of the Somali Embayment. Lower Jurassic neritic carbonates, the Didimut Beds, were recorded in the western margin of the Luug-Mandera Basin (Beltrand and Pyre, 1973). This depositional event is regarded to be linked to a phase of continental rifting. The first widespread transgression covered the horn during Early to Middle Jurassic time. Below is the description of stratigraphy and sedimentology of the Somali basins bordering the present Indian Ocean.

containing Middle Jurassic to Early Cretaceous ammonites were penetrated in the Brava-1 well (Fig. 2). This sequence, now called the Brava Formation, is correlatable with the Upper Jurassic shales that crop out in coastal Kenya (Beltrand and Pyre, 1973). The Brava Formation is unconformably overlain by the Gumburo Group (Upper Cretaceous). It contains over 1000 m of light gray shales, sandstones and siltstones, with rare limestone interbeds deposited in an inner middle neritic environment with strong deltaic influence. The Barren Beds, of Late Palaeocene to Middle Eocene age, include a maximum thickness of about 3000 m of predominantly shallow marine sandstones partially intercalated with siltstones and shales. Lignitic shales occur in upper part of this unit indicating inner neritic near shore depositional environment with deltaic influence. The Somali Merca Formation (Miocene to Pliocene)consists of mainly medium to fine grained sandstones and white to cream microcrystalline limestones, with thin shale and anhydritic interbeds of littoral to shallow marine environment. These older formations do not crop out because they are covered by Recent alluvial and eolian sediments along the coastal belt. The description of the formations is therefore based on subsurface data.

The Luug-Mandera Basin The existence of thick Pre-Jurassic sediments in the axial part of Luug-Mandera Basin is indicated by the discrepancy between section inferred from geophysical data and those measured in penetrated sections (Beltrand and Pyre, 1973). Little is known

Somali Coastal Basin (Fig. 3) Over 2000 m of predominantly greenish shales, occasionally interbedded with limestone bands, and

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'~ogadishu

THE COASTAL BASIN OF SOMALIA, KENYA AND TANZANIA from this section because of a lack of direct complete evidence, however, it is believed that it starts with continental conglomerates and sandstones comparable with the Karroo Series of northeastern Kenya, followed by an evaporitic sequence. Surface sections cropping out along the southwestern flank of the Luug Mandera Basin were studied thoroughly by, among others, Beltrand and Pyre (1972), and was later elaborated by Buscaglione and Fazzuoli (1987). On the basis of these works, the Middle Jurassic to Early Cretaceous section of this basin has been divided into four formations. The Baidabo Formation (Pleinsbachian to Bathonian) consists of over 800 m of predominantly thick bedded oolitic and algal limestones, interbedded with varicolored shales, resting unconformably on the Bur-Acaba crystalline basement. The basal 20 m of this formation is a coarse grained quartzitic sandstone (the Deleb Member) that passes laterally and vertically into varicolored shales (the Uanai Member). This is followed by detrital, generally oolitic, limestone with shale interbeds and abundant shell fragments (the Baidabo and Goloda members). These were deposited in continental to shallow shelf environments. The Baidabo Formation is unconformably overlain by the Anole Formation (Callovian to Oxfordian), consisting of gray calcareous shales, marly limestones and fossiliferous calcilutites, with abundant ammonites and belemnites. The Anole Formation corresponds to the maximum extension of the sea in this basin and its depositional environment ranges from shallow to deep shelf environments. The Uegit Formation (Kimmeridgian to Portlandian) consists of cyclically repeated oolitic, oncolitic, and bioclastic limestones, marls and ferruginous sandstones. It is thought to have been deposited on a restricted platform. It unconformably overlies the Anole Formation, and has a thickness of about 350 m. The base of the unit is marked by cross-bedded sandstones, indicative of the beginning of a regressive phase. Towards the end of the Jurassic the LuugMandera Basin became isolated from the Somali Coastal Basin. The Bur-Acaba region became totally emergent and acted as a physical barrier between the two main basins (Beltrand and Pyre, 1973). By Late Jurassic to Early Cretaceous time an evaporitic basin was established in the low areas of the basin, resulting to the deposition of the lagoonal Garbaharre Formation (Portlandian to Lower Cretaceous), while continental sandstones were deposited on the flanks. The Garbaharre Formation includes two members. The lower Busul Member (tentatively referred to the Late Portlandian) consists of about 300 m of yellowish bioclastic packstones and grainstones, gray bioclastic wackestones, quartzose sandstones and yellowish dolomites. The upper Mao Member consists of alternating gypsum and anhy-

215

drite levels with thin beds of shales, calcarenites, dolomites and cross-bedded sandstone, and is about 310 m thick. In the southwestern part of the basin predominantly reddish coarse to medium grained quartzose sandstones (the Amber Beds) crop out. These are about 180 m thick, consisting of nearshore to continental deposits. No index fossil had been found in it, and it has tentatively been referred to as Early Cretaceous. From then onwards the basin ceased to subside.

The Somali Embayment (Fig. 4) None of the wells drilled in the coastal part of the basin reached the basement, however, the Garade-More-1 well drilled on the northern flank of the basin reached the basement. The Adigrat Formation is the basal formation of Somalia. Its age is controversial because it is contains only poor or non diagnostic fossils. In the Garade-More-1 well the basal 80 m of the Adigrat Formation consists of fine to very fine grained carbonate-cemented sandstone interbedded with continental shales, overlain by 40 m of interbedded dark marl and very fine-grained argillaceous tidal flat dolomites. The Adigrat Formation is regionally diachronous, and its upper limit lies within Lower to Middle Jurassic (Dualeh, 1986). The Hamenlei Formation of Bathonian to Oxfordian age, described at the type section in the eastern Ogaden basin, is a white well bedded, mainly oolitic, fossiliferous limestones having thickness of about 210 m (Barnes, 1976). In the Garade-More-I well it is about 1093 m thick, consisting of mainly, oolitic and pseudo-oolitic packstones and grainstones interbedded with silty mudstones and dolomites. In the Obbia- 1 well 2175 m of dark gray shales and gray argillaceous limestones were penetrated without reaching the base. It is difficult at the moment to define the lower limit of the unit because it is transgressive. The depositional environment is back reef shoal in the Ogaden, tidal flat in the GaradeMore-1 area and basinal in the Obbia-I well, where the lowermost part penetrated is believed to belong to the Pleinsbachian-Toarcian stages. The Urandab Formation (Kimmeridgian) unconformably overlies the Hamenlei Formation. It crops out also in the eastern Ogaden where it consists of 55 m of gray to green gypsum bearing shales with intercalations of argillaceous limestones (Barnes, 1976). In the Garade-More-1 well the Urandab Formation consists of 95 m of gray to green slightly laminated glauconitic and pyritic marls with the lower 18 m consisting of oolitic interclasts. The maximum thicknesses recorded for this formation is in the subsurface of the eastern part of the Somali Embayment, where it is made up of 500-700 m of basinal dark gray shales and gray, marly limestone beds.

216

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THE COASTAL BASIN OF SOMALIA, KENYA AND TANZANIA The formation was deposited in a tidal flat complex in the Garade-More-1 well, is basinal in the mid Somali Embayment and of shallow marine origin in the eastern Ogaden. The Gabredarre Formation (Tithonian) was deposited in the Somali Embayment as a basinal sequence. In the Obbia-1 well it consists of 343 m of dark gray and dark brown shale with some gray fine crystalline limestones. This contrasts with the 410 m in the type section in east Ogaden, where it consists of partially oolitic yellowish, fine crystalline limestones with interbedded marly gypsum and limestones (Barnes, 1976). A regression began throughout the region during the Tithonian. In the south and central Somalia and southern Ogaden an evaporitic formation (Main Gypsum Formation) is generally regarded as Tithonian to Albian in age on the basis of the stratigraphic position. Also the Gabredarre Formation is totally missing from the Garade-More-1 well on the northern flank of the Somali Embayment, this was probably due to epeirogenic movements. The type locality of the Main Gypsum Formation is in the southern Ogaden. It consists of around 200 m of gypsum with marls and intercalation of calcareous lagoonal shales. Eastwards it grades laterally into the Cotton Formation (Lower Cretaceous). The type section is in the Cotton-1 well where it consists of fore-reef limestones and middle neritic shales. In the Somali Embayment, the Cotton Formation is a deep marine limestone. The Cotton Formation is overlain by the Gumbro Group (Cenomanian to Maastrichtian). In the Obbia well of the Somali Embayment this consists of 640 m of the light coloured, fossiliferous and porous limestones, with a few beds of dark gray shales. In the eastern Ogaden this unit consists of fossiliferous, lignitic and pyritic dark gray shales deposited in deep quiet waters, with restricted circulation (Clift, 1956). In the southern Ogaden and southern central Somalia, the Gumbro Series has been divided into four formations, each with a distinct characteristic lithology. The Mustahil Formation in the Ogaden is about 200 m thick, consisting of alternating white to yellow lenticular marly limestones and marls with gypsum at the top. The Ferfer Formation type section is also in the Ogaden, where it consists of about 200 m of gypsum with calcareous marly and shaly intercalations. The Beletutuen Formation type section is at Beletutuen in south central Somalia where it is about 415 m thick, mainly limestone bearing gypsum with shale and sandstone beds. The type section of the Jesomma Formation is in south-central Somalia and consists of 350-400 m of fine to coarse grained continental sandstones with local gypsum beds at the base. This formation marks the end of the Cretaceous cycle in Somalia. Tertiary sedimentation is believed to have started with the deposition of the Anrado Forma-

217

tion during the Early Eocene. At the type section in Nogal (northern Somalia), the Anrado Formation consists of 550 m of fine crystalline compact, light brown limestones with local thin gray, shaly beds of shallow marine environment. The Taleh Formation (Middle Eocene) was deposited in an evaporitic basin in Nogal, central Somalia and northwestern Somalia. At its type section in Nogal, the unit is 450 m thick and consists mainly of gypsum, anhydrite and shales with intercalations of limestone and cherty, marly beds. Its facies changes from being evaporitic in Nogal eastwards to deep marine fine clastics offshore. The Late Eocene section is called the Karkar Formation. In the Nogal area this is about 400 m thick, consisting of chalky limestone with intercalations of paper shales and gypsum (near shore facies with a hypersaline lagoonal episode). Eastwards, it changes to deep water facies, where in the Obbia-1 well 360 m of interbedded gray micaceous shales of Late Eocene age occur. The Karkar Formation marks the last major transgression in the region. Subsequently marine sedimentation was restricted to the present coast and offshore areas. About 731 m of coarse grained friable sandstones and greenish silty shales with chalky porous limestones have been penetrated in the Obbia-1 well, these are referred to as undifferentiated Miocene deposits.

GEOLOGICAL REVIEW OF THE KENYA COASTAL BASIN Introduction

The Kenyan Coastal Basin is dissected by a number of reactivated Mesozoic structures (Fig. 6). The basin is considered by Reeves et al., (1987) as the southern arm of end Jurassic/Early Cretaceous rifting which lead to the opening up of Indian Ocean. The Somali Coastal Basin is the northern arm, and the Anza Graben to the northwest failed to open. This is when Madagascar is considered to have migrated southwards. The stratigraphy of the basin has recently been reviewed by among others Cannon et al., (1981) and Rais-Assa (1986, 1988). Figure 8 compares the stratigraphic terminology for the basin according to different authors. The oldest and most extensive outcropping sedimentary rocks locally called the Duruma Series or Duruma Sandstones are of Karroo age, faulted to the west against the basement, while to the east they either disappear under the cover of post-Karroo rocks, or are faulted against them. Salt diapirs detected by geophysical data offshore (Figs. 6 and 7) are thought to be of this age and equivalent to the Mandawa salt basin of southern Tanzania. To the north, Karroo rocks thin beneath the cover of non-marine Neogene, un-

218 derlying the Cretaceous and Tertiary sediments in the southern part of Anza Graben. They reappear again north of Wajir in the Luug-Mandera Basin discussed later. Karroo rocks are overlain by limestones of Middle Jurassic age (the Kambe Formation) which marks the beginning of the post-Karroo marine phase. Karroo

The basal part of Duruma Series, referred to as the Taru Grits (Caswell, 1953, 1956) or the Taru Formation (Cannon, 1981) is made up of coarse grained, poorly sorted, feldspathic, fluvial sandstones. The formation reaches a maximum thickness of 2700 m, and has been dated as Upper Carboniferous to Permian on the basis of the fresh water bivalve Paleonodante fischeri known only from beds of Upper Carboniferous to Late Permian age (Walter and Linton, 1973). The Taru Formation is overlain by the Maji ya Chumvi Formation, consisting of fine grained laminated gray, silty shales and flaggy sandstones with ripple marking, cross-bedding and frequent sun cracking and rain pitting, often intercalated with carbonaceous beds. The Maji ya Chumvi Formation is considered by Walter and Linton (1973) to represent a swampy to lagoonal depositional environment between Late Permian and Early Triassic age (Rais-Assa, 1988). The overlying Mariakani Formation, though much coarser than the former, is also made up of fine grained to medium grained flaggy sandstones, micaceous siltstones and silty shales arranged in upward-coarsening deltaic sequences. Together with the overlying highly micaceous, black shale with coaly beds intercalations, the Mariakani Formation is dated as Middle to Late Triassic. The Manzeras Sandstone Formation of the uppermost Duruma Series is composed of coarse grained and cross-bedded sandstones. The lower part shows an upward-coarsening sequence of probably deltaic origin while the upper part is eolian. The contact between the Manzeras Sandstone and the underlying Matolani Formation is quite distinct to the south. Here it is described as unconformable with evidence of active faulting and fracturing at the top of the Manzeras Formation (Rais-Assa 1988). The Upper Manzeras shows progressive discordance with many synsedimentary faults and synsedimentary submeridinal synclinal structures. These movements, also recorded at the end of Karroo deposition of the Ngerengere Beds, are said to have originated from epeirogenic uplift of the margin related to the initial stage of continental rifting which was followed by an erosive phase. This is probably the End Jurassic to Early Cretaceous rifting described by Reeves et al. (1987) before Madagascar started drift-

E.I. MBEDE and A. DUALEH ing southwards. The Manzeras Sandstones, which mark the end of deposition in the Karroo Basin in Kenya have not yet been dated, but Caswell (1953) attributes them to the Upper Triassic, while Cannon et al. (1981) dates them as Lower Jurassic. Post Karroo

Apart from the Manzeras sandstones, no Lower Jurassic rocks have been described from coastal Kenya. Their absence has been interpreted as either due to continued deposition of the Manzeras, or a break in sedimentation (Rais-Assa, 1988). The Kambe Formation (Middle Jurassic) unconformably overlies the Manzeras Formation. It is faulted in some parts to the west against the underlying Karroo rocks. Lithologically it consists of dark gray reefal to lagoonal, pisolitic and oolitic limestones interbedded with the calcareous Posidonia shales of Westermann (1975), reaching up to 65 m. They are interpreted as deposited well below the photic zone in quiet poorly aerated waters. The limestones are characteristic of shallow water environment. The Kibiongoni Beds are grouped together with the Kambe Limestone as the Kambe Formation by recent workers (Rais-Assa, 1988). They lie unconformably on top of the Kambe Limestone. They contain boulders of both the Manzeras sandstones and the Kambe limestones, so they are clearly younger. They are also lithologically distinct from the Kambe limestones. They start with a basal conglomerate followed by unfossiliferous sandstone bearing ripple marks, rain pits and, on top, are the silty shales of restricted, presumably estuarine origin (Westermann, 1975). The Kambe Limestones have been dated as Bajocian/Bathonian, but the age is still questionable. The Kibiongoni beds are believed to be of Bathonian/Callovian age. The Kambe Formation grades upwards into Upper Jurassic shales including the Miritin, Rabai and Changamwe shales of Caswell (1953). These shales have also been referred to by Cannon et al. (1981) and Rais-Assa (1988) as the Mto Mkuu Formation. The basal part of the Mto Mkuu Formation described by Walter and Linton (1973) contains pebbles of both the Manzeras and Kambe Formations. It is cross-bedded. The formation is thought to be deltaic in origin. Apart from a small patch of Freretown limestone north of Mombasa no Cretaceous outcrop has been reported along coastal Kenya (Haughton, 1963). More than 4000 m of Cretaceous sediments are reported at the depocentre of the Lamu Embayment (Walter and Linton, 1972) and a thick sequence of Cretaceous sediments was encountered at Deep Sea Drilling Project (DSDP) site 241 east of the Lamu depocentre. The Lower Cretaceous clastics of coastal Kenya are made up of series of quartzites of Neo-

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THE COASTAL BASIN OF SOMALIA, KENYA AND TANZANIA containing planktonic foraminifera, deposited in a deep quiet water environment in the depocentre, while along the margin limestones containing a very sparse benthonic fauna are reported. Tertiary sequences are well developed in the Lamu Embayment being quite thick at the depocentre. Leg 25 of DSDP found an almost complete Tertiary sequence, with only part of the Oligocene and Upper Eocene section missing. Palaeocene sediments do not crop out in coastal Kenya, but are present at depth, being quite thick in the central part of the basin wedging out northwards and westwards, to where Middle Eocene sediments lie directly on top of Cretaceous beds. Lithologically the Middle Eocene sediments consist of dense micritic limestones interbedded with dark gray to brown shales deposited in a shallow marine environment. The Middle Eocene to Oligocene sediments crop out along coastal Kenya as a belt of low lying hills running parallel to the present coast. South of the Lamu Embayment they are argillaceous and feldspathic sandstones, poorly sorted and friable, with variable development of limestones, deposited in fluviatile, littoral and deltaic environments. Further north evidence of marine environment disappears and a sequence of variegated unfossiliferous red green mudstones and poorly sorted sandstones appears. The Upper Eocene rocks reported in the Anza and Bahati wells (Fig. 6) are barren continental beds, similar to those described in the Somalia Coastal Basin described earlier. Miocene beds are well represented in wells drilled in the Lamu Embayment. At the depocentre they are mainly limestones, dolomitic in some parts, with veins of anhydrites. Traced northwards and westwards they become sandy grading into variegated mudstones. They represent a transgressive phase, which began at the end of Oligocene. Pliocene to Quaternary deposits are mainly unfossiliferous laid down in fluvial and eolian environment. Along the coast the succession is of reefal limestone of Palaeogene to Neogene age, while offshore marine deposition has been taking place. Figure 7 is the section across the Kenyan coast.

THE GEOLOGICAL REVIEW OF TANZANIACOASTAL BASIN Introduction The main sub-basins distinguished in coastal Tanzania include the Selous-Ruvu-Tanga rift basin, along which a NNW-SSW fault trend is predominant, and the Lindi Rift Basin, to the south of which the Mandawa salt basin forms part, containing a NNW-SSE Karroo fault trend (Fig. 13). Both basins are Mesozoic to Tertiary in age and are crossed

221

by a number of north-south structural highs, while E - W is another remarkable fault trend in the basins. The major structural features are shown in Fig. 9. The rifting process, which began during early stages of the eastern Gondwanaland break up, continued progressively into the Early Jurassic. Periodic movements along bounding faults lead to the cyclic deposition of continental, fluviatile and lacustrine sequences, with periodic marine incursions in areas of high subsidence rates. These resulted in the development of restricted marine basins or gulfs where black shales and evaporites were deposited. At the close of the Early Jurassic tectonism along bounding faults became less intense, and a shallow marine transgression followed during the Middle Jurassic as a result of continued subsidence and tilting. An overall regression, associated with a number of tectonically influenced sea level fluctuations is remarkable from the early part of Cretaceous. A major regional unconformity is recognized at the base of the Aptian and Lower Albian. Activities at this time may be related to the continuous wrench faulting which lead to the southward movement of Madagascar. The Late Cretaceous and Early Tertiary was a period of tectonic stability, while significant tectonic reactivation began in the Late Palaeogene and continued into the Neogene. These movements were related to the development of the modern East Africa Rift system, which resulted in massive structural inversions and intensified eastward tilting of the present onshore areas, where they were accompanied by rapid subsidence and deposition. The development of offshore basins such as the Mafia and Zanzibar channel is believed to have taken place at this time. Selous-Ruvu-Tanga Basin This basin lies within the NNE-SSW trending rift (Tanga fault trend), the area described includes the area from 11~ S northwards to the Kenyan border together with the offshore areas (Fig. 9). The Karroo sediments in this basin consist of a complex series of continental fluviatile systems developed along the downthrown flanks of normal faults of Late Carboniferous to Early Mid Jurassic age. Movements along the faults were periodic and initiated erosive phases which generated fluvial systems which deposited upward-fining megacycles with basal conglomerates. These pass upwards into high energy braided stream deposits, and are succeeded by low energy meandering stream, flood plain and deltaic to swampy deposits. At least two Karroo megacycles are recognized within this basin. The older Karroo sediments include the Lower Permian Hatambulo Formation, described by Hankel (1987) around Stigler's Gorge. These are composed of tightly cemented feldspathic sandstones of deltaic and la-

222

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with red interbeds and gray silty claystones. Sediments in this area are considered to be lacustrine to braided stream deposits, while evidence of occasional marine incursions, coaly beds and deltaic influence exists. Geophysical data reveal thicknesses of above 3000 m (Kent et al., 1971 and Kent and Pyre, 1973), while recent seismic interpretations indicate that significant amounts of Triassic sediments must have been removed by Jurassic erosion. Seismic interpretation in this area also revealed complex sequences within the Ngerengere Beds, separated by disconformities, indicating different cycles of deposition in response to tectonic control by the bounding faults. In the central part of the basin at Ngerengere, the Ngerengere Beds consist of arkosic sandstones, with occasional limestones, and shales reflecting a high energy environment, becoming much quieter with time. There is evidence for a marine incursion

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224 before the full marine environment was established during Bajocian/Bathonian time. The intermittent faulting that influenced the cyclic deposition of the Ngerengere beds almost ceased during the Bajocian. Continued subsidence lead to the development of a shallow epicontinental sea by the end of the Bajocian and left the Karroo grabens along coastal Tanzania flooded with marine waters overstepping the basement. On the crests of basement ridges reefs developed in Bathonian and Callovian times. Around Tanga, the Ngerengere Beds are overlain by 340 m of compact well bedded oolitic to pisolitic limestones (the Amboni Limestone), of Bajocian/Bathonian age. Around Msata, Lugoba and Msolwa reefal limestone, or reworked rubbly limestone, directly overly basement, while at Kidugalo and Kidunda calcareous sandstones and sandy oolitic limestone of Bajocian to Callovian age are recorded. Between Msata and Msolwa, however, dark gray Posidonia shales, considered to be the lateral equivalent of Middle Jurassic limestones, directly overly the basement. Posidonia shales have also been recorded at depth in the wells drilled in the north of this area. The oolitic and argillaceous reefal limestones mentioned above indicate a near shore environment with a limited clastic supply during the Middle Jurassic. Deep drilling indicates that the limestone is replaced basinwards by deep marine shales. The dark gray Posidonia shales described above were deposited in a restricted back reef environment. The Middle Jurassic transgression continued into the Late Jurassic, but this was a rather quiet phase, and deposition was mainly marine with predominantly clastic sediments. These show an upward decrease in grain size from sand to clay throughout the Kimmeridgian along coastal Tanzania (Fig. 11). This low energy environment prevailed in most parts of coastal Tanzania for the whole of the Late Jurassic, while fluviatile deposition was taking place in some parts of the basin. Around Tanga, 700 m of Oxfordian to Kimmeridgian interbedded shales, sands and marls are observed, while over 1000 m of sandstones, limestones and mudstones are exposed along the Wami River to the south. The PuguMusanga high is believed to have been emergent at this time, and marginal Upper Jurassic coastal facies have been recorded in the Kisangire-1 and Kisarawe-1 wells drilled on top of this high (Figs. 9 and 10). To the west restricted water circulation lead to the deposition of neritic to bathyal deposits around Wangiyongo (Kajato, 1982). Further south, in the Selous Rift, a new fluvial system developed and deltaic deposition prevailed on the Callovian shelf. This delta continued to prograde northwards during Late Jurassic into Neocomian time. In general there is an overall shallowing of the sea from the

E.I. MBEDE and A. DUALEH central part of the basin northwards and southwards, while to the east deep marine deposition continued. The Late Jurassic transgression was followed by a period of regression starting in the Neocomian. This regression resulted in the deposition of fluviatile sandstones that form a major reservoir along coastal Tanzania. In the Kisarawe-1 well 500 m of Neocomian sandstones were encountered, and Wangiyongo borehole found more than 700 m of deltaic bituminous sandstones interbedded with clay that graded downward into Upper Jurassic marine shales. Along the Bagamoyo road septarian limestones and conglomeratic sandstones crop out, indicating some tectonic activity at this time. The Neocomian regression was followed by a period of fluctuating water level, open marine conditions prevailed in response to a gradual subsidence and rising of the sea level. This Middle Cretaceous deposition developed into a full marine transgression during the Late Cretaceous. Tectonic activity at this time may have resulted from the influence of both continued wrench faulting, as Madagascar continued to move southwards, and the effects of the break-up of West Gondwanland (the separation of South America from Africa). In this basin the AIbian to Cenomanian sequence overlaps Neocomian sediments, are themselves transgressed by Senonian strata, indicating locally complex disconformities. A distinct regional unconformity related to the worldwide to Austrian unconformity can be observed at the base of the Aptian and Lower Albian section. In the Kisarawe-1 well, Aptian siltstones and interbedded shales, indicate deposition in an open marine to outer shelf environment. To the north the Aptian sediments consist of upward-fining sequences of gray calcareous claystones, with minor siltstones and limestones, deposited in a marine mid-shelf environment. Aptian sediments are overlain by Albian limestones, marls and mudstones. These extend from Kisarawe to Musanga in the south, and to Wami in the north, where 900 m of marine clays, sandstones and limestones of Aptian to Turonian age crop out along the river. Eastwards, in the Pemba-5 well more than 800 m of Upper Senonian silty mudstones and rare limestones were deposited in an outer shelf environment. Upper Cretaceous outer shelf sediments are also recorded in the Zanzibar-1 well, Rasimachuis-1 (Campanian-Maastrichtian) and Ruaruke- 1 (Turonian-Campanian) and also exposed east of Msata. 900 m of Albian sediments were recorded in the Chalinze borehole. In the Maneromango area to the south Neocomian sandstones are overlain by a complex series of lithofacies deposited in inner shelf environment. The basal conglomerate is Aptian in Agwe. To the east, more than 1000 m of Albian to Maastrichtian sediments unconformably overlie the Neocomian strata. They are coarse grained, indicat-

THE COASTAL BASIN OF SOMALIA, KENYA AND TANZANIA

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ing a shallowing of the sea in this area towards the end of the Cretaceous. In general the Mid to Upper Cretaceous sediments in this basin indicate deep marine conditions (Kent, 1971). It is likely that the Upper Cretaceous sea transgressed westwards beyond the existing rift boundaries. To the south, a sequence of gently dipping sandstone crops out extensively across the Selous rift from the Rufiji River southwards to the Mozambiquan boarder. It oversteps Karroo sediments to directly overlie basement and is overlain by Miocene sediments. It has been correlated the Aptian Makonde Beds on lithological similarities (Spence, 1957). Other workers place these rocks at the top of the Karroo Series, while others refer to them as Middle to Upper Jurassic in age. The Upper Cretaceous transgression continued into the Early Tertiary with a minor regional hiatus at the Palaeocene-Late Cretaceous boundary. Palaeocene sediments are mainly deep marine dark gray to green clays with occasional sandy and silty intercalations. Gas shows were reported in Palaeocene beds of the Zanzibar-l, Pemba-5 and Mafia-1 wells. These sediments indicate shallowing of the sea. An ensuing regressive phase started during the Mid Eocene and continued into the Oligocene, whose strata are largely absent along the Tanzanian coast. This regression may have resulted from isostatic adjustment due to active sea floor spreading related to development of the Owen fracture and widening of the Indian Ocean. It reflects building up of broad continental shelf and slope at a faster rate than the subsidence which was probably continuous from the Late Cretaceous. Though there are

abundant instances of lateral facies variation over a short distance, clastic deposition was the most dominant mode of deposition. Miocene sediments overstep older beds, indicating renewed transgression and tectonic activity contemporaneous with the development of the modem East African rift system. Uplift and basinward tilting of hinterlands, together with intensive erosion and rapid basinal subsidence, resulted in the deposition of large volumes of sediments and in the development of a large prograding delta across the Dar-es-Salaam Embayment and Zanzibar channel. This is accompanied by synsedimentary faults similar to those observed on a seismic section across the Zanzibar channel (Fig. 10). In the Pemba-5 well to the north, there is evidence of deltaic deposition in Oligocene to Lower Miocene sediments. The Middle to Upper Miocene sediments consist of marine limestones, silty mudstones and sands indicate distance from the delta. The E - W trending Rufiji depression is believed to have formed at this time. The final regression resuited in a widespread unconformity over Miocene deposits. Marine Pliocene deposition was limited to the present day offshore areas. Onshore Neocomian to Miocene beds are overlain by regressive estuarine fluviatile Plio-Pleistocene sands in existing depression such as the Ruvu valley and the Rufiji depression. The Lindi Rift Basin The Lindi rift is a huge basin located in southeast Tanzania, of which the Mandawa salt basin forms

226

E.I. MBEDE and A. DUALEH

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part (Fig. 9). A dominant N-S structural trend can clearly be seen. In the Mandawa Basin the Pande and Kizimbani highs and the Mandawa anticline are examples. The prevalent Karroo fault direction here is N N W - S S E (Lindi fault trend) while NNE-SSW and E - W trends are also observed. The Karroo litho-

facies in this basin are more distinct and more varied due to differences in the subsidence rates within half grabens. On the eastern flank of the Matumbi hills crop out 3000 m of highly faulted Karroo proximal fanglomerates, massive red to brown sandstones, siltstones and mudstones attributable to braided and meandering stream deposits. Nodular limestones and sandstones indicate occasional marine incursions. To the south at Mandawa evidence of marine incursion is more obvious because early Karroo clastics are followed by transitional continental to marine beds then lagoonal shales and evaporitic sequence of Triassic age. Thicknesses and the distribution of the Ngerengere Beds in this basin are complicated by extensive faulting. Around the Matumbi hills, Middle Jurassic oolitic and reefal limestones were deposited on top of the basement, indicating shallow near shore marine environments, similar to those described earlier from the Ruvu Selous Basin to the north. Southwards, in Mandawa, there are Middle Jurassic marine marls with subordinate oolitic limestones. A local intra Mid-Jurassic unconformity occurs between Bajocian and Bathonian sediments indicating early salt movement in the basin. The salt pillow left by these movement controlled Late Jurassic and Cretaceous deposition in southern Tanzania. A change to littoral and eventually neritic environment can be observed. In Mandawa shales intercalated with evaporites were first deposited, grading into marls, sandstones and then into continental sandstones. Lagoonal shales were still deposited in a restricted marine basin west of the anticline. Northwards, around Matumbi, marine and deltaic buff sandstones and clays were deposited during the Upper Jurassic. Basinwards, deposition of deep marine clays took place. Around Songosongo Upper Jurassic marine clays are reported to have probably produced gas from Cretaceous reservoirs (Kajato, 1982). The End Jurassic to Early Cretaceous left lateral movement that split the Kizimbani and Pande highs are thought to have also caused salt flow which resulted in the formation of the N-S trending Mandawa anticline. These movements are thought to have been related to the southward drift of Madagascar. They also resulted in a local transgression to the west during the Early Cretaceous when Neocornian sandstones, siltstones and occasionally limestones, were deposited. The Neocomian transgression became more widespread during the Lower Aptian with initial deposition of conglomeratic sandstones and siltstones southwards and westwards of the Mandawa anticline. These are the Makonde beds. Reefal Orbitolina limestones were deposited east and west of the Kiturika hinge line east of Mandawa, these limestones extend southwards to Lindi, and pass laterally into Upper Aptian Kihuluhulu and Lower Albian Kigongo marls (Fig.

THE COASTAL BASIN OF SOMALIA, KENYA AND TANZANIA

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11). South of Mandawa there are Upper Aptian and Albian clays with septarian nodules. The widespread Mid-Cretaceous transgression continued into the Late Cretaceous, facilitated by the eastward tilting and subsidence of the basin. Deep marine environments were established just east of the Mandawa anticline where argillaceous facies intercalated with bands of arranaceous beds were deposited. Palaeocene sediments in this basin indicate the last phase of the Cretaceous transgression, with a noticeable unconformity at the Late Cretaceous/ Palaeocene boundary. Sediments deposited during Palaeocene are mainly clays, marls, siltstones, sands, reefal and algal limestones changing laterally into marls and clays basinward. The beginning of the Eocene Period is associated with tectonic activities and marks the onset of a regression that continued into the Oligocene. Mid and Upper Eocene sediments reflect continental slope build up probably caused by basement faulting. By the Late Eocene a continental shelf was established where platform carbonates were deposited. Gas shows are reported from Eocene limestone reservoir in Songosongo. An angular unconformity occurs at the Middle and Upper Eocene boundary at Mandawa. Oligocene sediments are limited in their distribution and were probably eroded prior to Miocene faulting related to the development of the modem East African rift system. Between Lindi and Kilwa Miocene sediments are well developed and are strongly transgressive, extending to the N.E. of Mandawa. The Pliocene was another phase of local transgression, which deposited the continental Mikindani Beds. These are mainly sandstones, gravels and clays laterally merging with marine reefal limestones and sandy clays.

STRUCTURAL EVOLUTION

The structural development of the eastern margin of Africa is related to the fragmentation of eastern

227

Gondwanaland and the birth of Indian Ocean. This is demonstrated by the correlation of sedimentary units from deep sea drilling well locations offshore and those cropping out onshore (Wolfgang et al., 1974). Faulting controlled subsidence and sedimentation, while flexure dominated the region through out the time. The Indian Ocean, however, is one of the least understood seaways in the world. The western Indian Ocean basin, the so called Somali Basin, has recently been intensively studied by among others Beltrand and Pyre, (1973), Schlich et al., (1974), Kent, (1972), Mascle et al., (1987), Rabinowtz et al., (1982, 1983), Bosellin, (1986) and Dualeh and Naim, (in press). The story of its development can be summarized as follows. The first stage of Indian Ocean development is believed to have taken place from the Late Jurassic to Mid Cretaceous when Madagascar is believed to have migrated southwards (Rabinowtz et al., 1983). Oceanic crust created by this movement must have formed the floor of the proto-Indian ocean at that time. The ocean that existed in the region during Early to Middle Jurassic time, the "Neotethys sea", is considered to have been an epicontinental sea (Cannon et al, 1981). The second stage took place in two phase, the first during the Late Cretaceous, when a rift separated Mascarene from Madagascar. During the second phase in the Palaeocene a new branch of the southern rift separated the Indian subcontinent from the Seychelles microcontinent along the Carlsberg Ridge, leading to the closure of the "Neotethys sea". It is the first stage of Indian Ocean development that affected most of the area discussed in this paper, however, the impact of earlier and later events can be seen in sedimentary units. The development of the eastern margin of Africa began with the formation of an intracontinental rift during the early stages of the eastem Gondwanaland fragmentation. Initial faulting is believed to have taken place during the Late Carboniferous if not Early Permian. Rocks of this age have been recorded in Kenya and Tanzania, while in Somalia Karroo rocks are questionably dated as Triassic. Early tectonism resulted into the formation of NNE-SSW, NNW-SSE and in some places E - W trending faultbounded basins and sub-basins. Continental clastic sediments were deposited in these constantly subsiding grabens and half-grabens while continued subsidence lead to partial marine incursion in the fast subsiding sub-basins. Marine influence increased throughout the Jurassic leading to the establishment of full marine conditions in the region during the Middle Jurassic. Magnetic anomalies in the Gulf of Somalia (Rabinowtz et al., 1983) show that Madagascar started to move southwards from East Africa 156 m.y. ago (M 25). Figure 13 is a palaeogeographic recon-

228

E.I. MBEDE and A. DUALEH

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ANTACTICA

Fig. 13. Tectonic and palaeogeographic reconstruction of east African continental margin during the late Late Jurassic to Early Cretaceous time.

struction of Madagascar during Late Jurassic-Early Cretaceous times. This phase was associated with Upper Jurassic to Early Cretaceous marine shales recorded almost everywhere in the region. Marine deposition terminated in the failed rift basins such as the Luug-Mandera Basin. N-S trending gravity highs in Tanzania, as well as in Kenya, are thought to indicate strong crustal changes caused by early transform movement (Rabinowtz, 1983). Madagascar is believed to have attained its final position 130 m.y. ago (M 9). Figure 14 shows the position of Madagascar relative to other major physiographic features in the Cretaceous Period. This when the continental margin was fully formed, and the first expression of the Owen fracture zone ap-

peared. Progressive eastwards shift of the depocentre is observed from the Cretaceous into the Tertiary as a result of a combination of continued regional sag associated with sea floor spreading of Indian ocean. By the Palaeocene to Early Eocene India is believed to have started moving northwards. This marks the last widespread marine episode. The Mid Eocene to Oligocene is a regressive period in the region and this is the time when the "Neotethys Sea" is believed to have been subducted. This regional regression has been related to the pre-rift doming of the Afro-Arabian shield. The last phase of tectonic activity in the region began in the Miocene and continues to the present day. This phase is related to the opening up of the Gulf of Aden and the Red Sea, and the

THE COASTAL BASIN OF SOMALIA, KENYA AND TANZANIA

!

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establishment of the East African rift system to the west. Major structural patterns related to this tectonic history are shown in Figs. 6, 9 and 12. No compressive events have been recorded in the region so far. Some anticlines, like Mandawa, have been related to salt movement, while other structural features, like the NNE-SSW trending Sengt anticline, and Garbaharry anticline of the Luug-Mandera Basin,

were formerly related to either salt movements, or to movement of Upper Jurassic shales (Pyre, 1972). According to Carmingnani et al. (1983), however, the anticlines are thought to be caused by wrenching associated with a transpressional component of fault movement. This is probably the case for other anticlines observed elsewhere in the region.

230 ECONOMIC GEOLOGY

Hydrocarbon prospectivity No systematic source rock studies have been done so far in the region, however, seeps are known in some parts of the basin (Tabajar, Kenya and Wangiyongo, Tanzanzia). A gas discovery, together with oil shows, has been reported from Songosongo and Mnazi Bay, Tanzania, and gas shows have been recorded in Tertiary reservoirs in most of the wells drilled in the region. These all indicate the existence of mature source rocks in the region. Exploration done so far has been inexhaustive. The few wells that have been drilled mainly tested structural traps that can be considered risky in this region due to updip migration during fault reactivation. Very little testing of stratigraphic traps has been done if any. In the early 1980s Robertson Research Group collected samples from deep wells in the region for geochemical studies, their results have not been published. The mainly lacustrine, fluvial, deltaic to swampy deposits of the early Karroo sequence can be expected to have locally developed source beds rich in either type I or type III kerogen. Whereas the late Karroo sequence contains restricted marine deposits in some parts can be expected to have well developed rich source beds. The source character of Karroo beds in the region is very little known, because the formation has not been reached in the deep wells drilled so far. Permian samples studied by BEICIP (1984) in Kenyan Karroos gave TOC values of up to 0.9% while Kamen-Kaye and Barnes (1979), described shales of this age in Somalia and N.E. Kenya Karroo sections, and suggested that they may have produced the gas found in the neighbouring Ethiopia. Outcrop samples described by Kreuser, (1984) from Tanga, Rufiji, Mikumi and Nyakatitu sections gave TOC values between 0.3-2.4% of structureless vitrinitic material, while east of Rufiji at depth TOC values of above 2% are recorded in Upper Triassic shales. In Mandawa marine shales of Karroo age contain sapropelic oil prone kerogen with a fair amount of extractable oil and TOC values of above 1%. The thick post-Karroo sequence is said to have developed in total marine environment accompanied with a series of transgressions and regressions which could have been favorable for good marine source bed development, with either paralic type II kerogen or type III kerogen where deltaic conditions existed. The Middle Jurassic Posidonia marine shales and equivalent interfingering shallow water limestones, have good amorphic algal material with sapropelic oil prone kerogen and good TOC values. Upper Jurassic shales from wells drilled in central part of Tanzania give average TOC values of 0.5%, with

E.I. MBEDE and A. DUALEH humic gas prone kerogen in some parts. They are believed to have generated the gas in the Songosongo gas field (Kajato, 1982). In Kenya Upper Jurassic shales have been penetrated in only one well (Garissa-1) where they contain type IV kerogen. Surface samples studied by different oil companies between 1974-1983 reveal that the shales of this age have very low TOC values, containing immature amorphic and herbaceous materials with type II oil prone kerogen and abundant coaly type III kerogen. The thick marine fossiliferous calcareous shales of the Brava Formation (Jurassic to Cretaceous) in Somalia exhibit favourable source character. They have organic contents of between 0.26-0.72% in the Jurassic section, and 1.2% in the Cretaceous section. This is in the form of oil prone kerogen with an average vitrinite reflectivity of 0.66Ro. Albian to Palaeocene clays and shales overlain by a thick Neogene sequence are well developed in Zanzibar and Pemba, and to some extent in the Mafia channels (Fig. 10). Their source character is not reported, but gas shows have been recorded from Tertiary reservoirs in some wells. The carbonaceous shales and lignites described in the thick pile of deltaic Miocene sediments of the Zanzibar channel could provide another source for this gas. The Lower Cretaceous section is not a source rock along coastal Kenya, though it might be basinward where the Upper Cretaceous sediments have source characteristics similar to Tertiary sections, with mainly detrital Type IV organic matter associated with Type III humic Kerogen. The latter limits their potential for liquid hydrocarbon generation (Mbede, 1987). In Somalia Cretaceous and Tertiary shales offer favourable source beds in the vicinity of potential reservoirs. The petrophysical character of sedimentary formations in the region is not yet well established. The stratigraphic evolution of the basins and sub-basins, however, suggest that several lithostratigraphic units could be significant hydrocarbon reservoirs under favorable conditions of generation and migration. At outcrop Karroo sandstones show very poor reservoir character, being mainly feldspathic and poorly sorted. Great burial depths before Tertiary inversions might have contributed to the destruction of porosity and permeability within Karroo sediments. The Upper Karroo section, however, is expected to be of good reservoir potential. The Adigrat Formation provides a gas reservoir in Ethiopia, while the Manzeras sandstones of Kenya are expected to have good reservoir character, being partially eolian. The Ngerengere Beds of Tanzania around Tanga give good neutron porosities in some wells, while in Mandawa they are described by Kajato (1982) to be of good reservoir potential. Middle Jurassic depositional environments are suggestive of good potential reservoir development

THE COASTAL BASIN OF SOMALIA, KENYA AND TANZANIA in most parts of East Africa, being mainly reefal, oolitic and fossiliferous limestones. The Lower Cretaceous sediments are known to be regressive and, in Tanzania, Neocomian fluviatile and deltaic sandstones are potential reservoirs. In the Songosongo field gas is found in a deltaic Albian sand which unconformably overlies extensive Neocomian continental sands (Kajato, 1982). In the northern part of Kenya Lower Cretaceous quartzitic sandstones are expected to have good reservoir character (Mbede, 1987), whereas in Somalia the Cretaceous section is mainly argillaceous, though locally some potential reservoir beds can be observed within the Gumburo Group of the Somali Embayment. Mid-Cretaceous lithofacies were deposited under fluctuating sea level prior to the regional Late Cretaceous transgression. They are expected to have stratigraphic traps within the Albian-Aptian section. Though the Upper Cretaceous is known to be transgressive almost everywhere in the region, significant Cenomanian and Turonian regressions deposited potential reservoirs. Kajato (1982) suggests that Upper Cretaceous sands are potential reservoir within the Mafia channel structure. There are several potential Tertiary reservoirs in the region, mainly along and seaward of the present coastline. Continental barren beds of northeastern Kenya and Somalia can prove to be of good reservoir character, while a gas blow-out in 1981 was reported from Eocene limestones in the Songosongo gas field, Tanzania. Palaeogene reservoirs have been reported in Zanzibar 1, Pemba-5 and Mafia-1 wells. The Mnazi Bay gas discovery is in Miocene sands, while gas shows have been recorded in Miocene sands of the Rasimachuis-1, Zanzibar-1, Pemba-5, Mafia- 1 and Kisarawe- 1 wells. Tertiary prospects are also widely reported in Somalia and Kenya, but their development is limited compared to that in Tanzania (Figs. 3, 4, 7 and 10). The tectonic development of the basins in this region leaves no doubt that both structural and stratigraphic traps must be well developed. Structural traps will mainly be related to the tensional forces related to early rifting processes. The later opening up of the Indian Ocean generated structures related to compression associated with wrench movements. Stratigraphic traps will be mainly related to the post-Karroo transgressive and regressive sequences. Middle Jurassic evaporites and shales can provide sealing within restricted Karroo basins, while Upper Jurassic shales can provide sealing for Lower and Middle Jurassic reservoirs. Upper Cretaceous shales can provide sealing for Lower and Middle Cretaceous reservoirs as exemplified in the Songosongo gas field. Widespread deep water deposits exist within the Tertiary sequence, including Miocene shales capping the Lower Miocene gas reservoir in Mnazi Bay.

231

Tensional forces have also occurred throughout the history of the region. Although partly good source rocks are at least developed locally in the region and reservoir potential are recognized, no commercial oil has been reported so far, despite drilling most major structures. The high geothermal gradients normally associated with rifting are expected to have speeded up maturation of potential source beds within Karroo formations. Reservoir potential is poor in these rocks, but traps are abundant, though sealing is a problem. The thick post-rift sequence on top of the Karroo sediments might have contributed to the overmaturity of Karroo source beds and to the destruction of porosity and permeability. Petroleum may have migrated into early traps, however, and then have remigrated into younger traps during later movements. Maturation within the post-Karroo sequence is believed to have been slow, because of a low geothermal gradient. This idea is supported by the immature Upper Jurassic outcrop sample collected in Kenya. Maturity may be expected to increase basinward. Jurassic source beds have not been reached in most wells drilled. In Kenya Lower Cretaceous source rocks have been found to be overmature, while Mid Cretaceous source rocks are just at the oil window (Mbede, 1987). Tertiary shales are immature in most wells except in Tanzania, where Tertiary development is extraordinary (Fig. 10). Here a burial curve reconstruction has proved that the maturity of source beds within the Pugu-Musanga structural axis is higher than one would expect for normal subsidence to their present depths. This is suggestive of Mid Miocene inversion which is believed to have resulted in the removal of several thousands of metres of section. Tertiary beds are suspected to have generated oil which migrated into traps within the vicinity, but Miocene inversion might have caused significant redistribution and eventual loss of hydrocarbons by erosion or leakage. Hence structural traps in this region are considered to be more risky than stratigraphic traps which need detailed geophysical mapping and are expensive to explore (Nagati, 1996).

Other mineral deposits Exploitation of other mineral resources is active around the region. Early Jurassic to Recent limestones are used by cement factories. Each country has at least one cement factory. Limestones, together with shales, are also used locally for the production of facing stones, aggregates and other building materials. High grade gypsum, interbedded with Triassic to Jurassic shales occur in the Mandawa Basin in south Tanzania. No exploitation of this deposit has ever taken place, but near Malindi (Kenya) gypsum is exploited on a small scale for the cement in-

232

dustry. The cement industry in Tanzania is supplied with gypsum from Recent lake bed deposits found together with clastic and limy materials. Kaolinite suitable for paper and rubber manufacture is produced from Late Miocene kaolinitic sandstones in the Pugu hills (Tanzania). Clean sands obtained as by-product of this kaolin industry is suitable for use as glass sands and is expected to be used by the Mbagala glass industry in Dar-es-Salaam. Beach sands are considered to be not so good for glass industry due to less of the 30-70 mesh, and a high proportion of impurities such as carbonate grains and refractory minerals (zircon and kyanite). A number of heavy mineral deposits exist but they have received little attention by prospectors so far. Workable amounts of ilmenite, rutile, zircon and leucoxene are reported. No exploitation of heavy mineral is going on at the moment, though reconnaissance exploration is known to have been carried out. More than 72 million tons of proved reserves were estimated (Halse, 1980) in Kenya and a mining license has been awarded. In Tanzania a pilot plant for heavy mineral concentration was established by STAMICO (the State Mining Corporation) in the 1970's around Dar-es-Salaam, but was later abandoned. Production of salt from sea water by solar evaporation is the major source of common salt within the countries, while the widespread sedimentary basins also provide aquifers for groundwater resources especially in the arid areas of northeastern Kenya and Somalia.

REFERENCES

Barnes, S.U., 1976. Geology and oil prospects of Somalia, East Africa, Am. Assoc. Petrol. Geol. Bull., 60(3): 389-413 Beltrand, M.D. and Pyre, A., 1973. Geological evolution of Southwest Somalia, in: G. Blant (Editor), Sedimentary Basins of African Coasts, part 2, South and East coast, Paris Assoc. Serv. Geol. Afr., pp. 159-191. Bosellin, A., 1986. East Africa continental margin, Geology, 14( 1): 76-78. Buscaglione, L. and Fazzuoli, M., 1987. Jurassic carbonate microfacies of Somalia, Dept. Di Sienze della Terra Univ. di Firenze, Firenze, Italy (unpublished). Cannon, R.T., Simiyu Siambi, W.M.N. and Karanjani, F.M., 1981. The proto Indian Ocean and probable Paleozoic/ Mesozoic triradial rift system in East Africa, Earth Planet. Sci. Lett., 52:419-426. Carmingnani, L., Aili Kassim, M. and Fantozzi, P., 1983. Nota preliminarie sul Rilevamento Della Regione Di Gedo (Alta valle Del Giuba), Somalia Meridionale, Geol. Somalia, 6: 85109. Caswell, P.V., 1953. Geology of Mombasa Kwale area, Rep. Geol. Surv. Kenya, Nairobi, 24 pp. Caswell, P.V., 1956. Geology of Kilifi-Manzera area, Rep. Geol. Surv. Kenya, Nairobi, 34 pp. Chaltellier, J. and Slevin, A., 1988. Review of African petroleum and gas deposits, J. Afr. Earth Sci., 1(3): 561-578. Clift, W.O., 1956. Sedimentary history Ogaden district, Ethiopia, In: Symp. Sorbe Petroleo & Gas, XX Int. Geol. Congr., pp.

E.I. MBEDE and A. DUALEH 89-107. D'Amico, C., Ibrahim, H. and Sassi, EE, 1982. Outline of the Somalia basement, Quad. Geol. Somalia Univ. Naz. Somalia, Mogadishu, 6: 15-32. Dualeh, A., 1986. Geological and stratigraphic evolution of the Northeast Somalia continental margins and adjacent areas. M.Sc. Thesis, Univ. South Carolina (USA), 129 pp. (unpublished). Dualeh, A. and Nairn, A. The continuing story of the fragmentation of Gondwana: A contribution from Somalia, In: Geology of Somalia and surrounding regions, proceedings (in press). Hankel, O., 1987. Lithostratigraphic subdivisions of the Karroo rocks of the Luwegu Basin (Tanzania) and their biostratigraphic classification based on microfloras, macrofloras, fossil woods and vertebrates, Geol. Rundschau, Stuttgart, Band 76(2): 539-565. Hasel, J.E.E, 1980. Heavy mineral beach sand deposits, initial exploration report. Haughton, S.H., 1963. The stratigraphic history of Africa South of Sahara, Oliver & Boyd Ltd, London, 365 pp. Kajato, H.K., 1982. Gas strike spurs: Search for oil in Tanzania, Oil and Gas J., March 15th, pp. 123-130. Kamen-Kaye, and Barnes, S.U., 1979. Exploration geology of North Eastern Africa-Sychelles Basin, J. Petrol. Geol., 2(1): 23-45. Kent, EE., 1972. Continental margin of East Africa: A region of vertical movements, in: C.A. Burke and C.L. Drake (Editors), Geology of Continental Margins, pp. 313-320. Kent, EE., Hunt, J.A. and Johnstone, D.W., 1971. The geology and geophysics of coastal Tanzania, Inst. Geol. Sci. London Inst. Geophys. Sci., Geophysics, Paper No. 6, 101 pp. Kent, P.E. and Pyre, J.T.O., 1973. The development of Indian Ocean margin in Tanzania, in: G. Blant (Editor), Sedimentary basins of African coasts, part 2, South and East coasts, Paris Assoc. Serv. Geol. Afr., pp. 113-13 i. Klemme, H.D., 1980. Petroliferous basins classification and characteristics, J. Petr. Geol. 3(2): 157-207. Kreuser, T., 1984. Karroo basins in Tanzania, in: Geologie Africaine, J. Kleerkx and J. Michot (Editors), Musee Royal de Belgique, Tervuren, pp. 231-245. Mascle, J., Mougenot, D., Blarez, E., Marinho, M. and Virlogeux, P., 1987. African transform continental margins: example from Guinea, Ivory Coast and Mozambique, Geol. J., 22: 53-56, pl 68. Mbede, E.I., 1987. A review of the hydrocarbon potential of Kenya, J. Afr. Earth Sci., 6(3): 313-322. Nagati, M., 1996. Tanzania wildcats to evaluate Jurassic Mandawa salt basin, Oil and Gas J., Oct. 7, pp. 109-112. Rabinowtz, P.D., Coffin, M.E and Falvey, D., 1982. Salt diapirs bordering the continental margin of Northern Kenya and Southern Somalia, Science 215: 663-665. Rabinowtz, P.D., Coffin, M.E and Falvey, D., 1983. The separation of Madagascar and Africa, Science, 220: 67-69. Rais-Assa, R., 1986. Geologie et evolution du bassin Permocarbonifere a Cretace de Mombasa (Kenya), Travaux des Laboratoires della Terra st Jerome Marseille ser X, 76, 25 PP. Rais-Assa, R., 1988. Stratigraphy and geodynamics of the Mombasa Basin (Kenya) in relation to the Genesis of proto Indian Ocean, Geol. Mag. London, 125(2): 141-147. Reeves, C.V., Karanjani, EM. and MacLeod, I.N., 1987. Geophysical evidence for a failed Jurassic rift and triple Junctions in Kenya, Earth Planet. Sci. lett., 81:299-311. Schlich, R., Simpson, E.S.W. and Vallier, T.L, 1974. Regional aspects of Deep Sea Drilling in the western Indian Ocean, Inst. Rep. Deep Sea Drilling Proj., 25: 743-759. Spence J., 1957. The geology of Eastern Province of Tanganyika, Bull. Geol. Surv. Tanganyika, 28 pp.

THE C O A S T A L BASIN OF S O M A L I A , KENYA AND T A N Z A N I A Walter, R. and Linton, R.E., 1973. The sedimentary basins of coastal Kenya, In: G. Blant (Editor), Sedimentary basins of African coasts, Part 2, South and East coasts; Paris Assoc. Serv. Geol. Afr., pp. 133-158 Westermann, G.E.G., 1975. Bajocian Ammonoid Fauna of

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Tethyan affinities from Kammbe Limestone Series of Kenya and implication to plate tectonics, Newslett. Stratigr., 4(1): 23-48. Wolfgang, S., Branson, J.C. and Turpie, A., 1974. A petroleum potential of the deep water regions of the Indian ocean.

C h a p t e r 11

The Owambo Basin of Northern Namibia

R. M c G . M I L L E R

The geology of the Owambo Basin is known from outcrops along its margins, from interpretation of seismic, aeromagnetic and gravity surveys and from a few widely spaced wells. The Owambo Basin is floored by mid-Proterozoic crustal rocks of the Congo Craton and contains possibly as much as 8000 m of sedimentary rocks of the Nosib, Otavi and Mulden Groups of the late-Proterozoic Damara Sequence, 360 m of Karoo rocks and a blanket of semi-consolidated to unconsolidated Cretaceous to Recent Kalahari Sequence sediments up to 600 m thick. Deposition of Nosib fluviatile sandstones began about 900 Ma ago during intracontinental rifting to the south and west and ended with local glacial deposition. As rifting evolved to spreading, the region of the Owambo Basin became a stable platform (Northern Platform of the Damara Orogen) marginal to oceans to the south and west. Carbonates of the Otavi Group were deposited between about 730 and 700 Ma on this platform. Two cycles of carbonate deposition (Abenab and Tsumeb Subgroups), each starting with quiet, relatively deep water (stromatolite-poor, laminated rocks) that became progressive shallower with time (stromatolites, oolites, evaporite(?) minerals), were separated by a widespread glacial episode (Chuos Formation). Reversal of plate motion in the adjoining oceans culminated in subduction and continental collision in the branches of the Damara orogen to the south and west. Erosion products of Dl uplift in the west were deposited as an upward-fining molasse (Mulden Group) in the Owambo Basin between 650 and 600 Ma. D2 deformation folded Mulden and Otavi rocks together and produced the uplifted, lblded western and southern margins of the Owambo Basin that give it its present form. After extensive erosion, a Lower Permian ice sheet north of and within the Owambo Basin deposited the Dwyka Formation tillite. As this ice sheet retreated, Lower Permian basin plain and fluvio-glacial sediments and low-grade coals of the Prince Albert Formation accumulated. Middle Permian to Lower Triassic sediments are absent from the Karoo succession and, rather than having been removed by intra-Karoo erosion, may never have been deposited in this region. Prince Albert rocks are overlain in the Nanzi well by Upper Triassic aeolian sandstone of the Etjo Formation, the only known occurrence in the Owambo Basin. Basalts, probably correlates of the Kaikrand Formation of southern Namibia, occur only in the far eastern parts of the basin. Continental Cretaceous to Recent aeolian sands and lacustrine clays of the Kalahari Sequence form a thick blanket over all older rocks in the Owambo Basin. The lacustrine clays with their associated fluviatile silts and sands may have been transported by endorheic rivers flowing from the northwest before being deposited in shallow lakes in the southern half of the basin in much the same way that the Okawango Swamps of Botswana are being fed today by the Okawango River. The present-day Etosha Pan developed as a result of pan-edge retreat.

INTRODUCTION T h e O w a m b o B a s i n is l o c a t e d on the C o n g o Craton b e t w e e n 14~ to 18~ and b e t w e e n the n o r t h e r n b o r d e r of N a m i b i a to 19~ (Fig. 1). It e x t e n d s n o r t h w a r d s into s o u t h e r n A n g o l a and c o u l d continue into w e s t e r n Z a m b i a . T h e e a s t e r n b o u n d a r y in N a m i b i a m a y be a buried b a s e m e n t swell t r e n d i n g in a n o r t h e a s t d i r e c t i o n from G r o o t f o n t e i n to M a s h a r i (Fig. 2). This swell is b r e a c h e d a b o u t h a l f w a y a l o n g its length by a s o u t h e a s t - t r e n d i n g valley a b o u t 350 m d e e p (Fig. 2). T h e r e g i o n is also referred to as the E t o s h a B a s i n but as such is too readily linked with the area a r o u n d the E t o s h a Pan w h i c h is only a rela-

tively small part of the g r e a t e r O w a m b o Basin. T h e n a m e E t o s h a B a s i n is thus s o m e w h a t m i s l e a d i n g . A l t h o u g h h a v i n g f o r m e d initially as a late-Prot e r o z o i c s e d i m e n t a r y basin, it w a s also a c e n t r e o f d e p o s i t i o n in the P e r m o - J u r a s s i c and a g a i n d u r i n g the late C r e t a c e o u s to Tertiary. T h e p r e s e n t f o r m is still basin shaped. T h e hilly to m o u n t a i n o u s rim on three sides of the basin r e a c h e s e l e v a t i o n s of 1 5 0 0 1700 m in the north in A n g o l a and 1 4 0 0 - 2 1 0 0 m in the w e s t and south. F r o m the l o w e s t p o i n t in the basin, the E t o s h a Pan, h a v i n g an a v e r a g e e l e v a t i o n o f 1084 m, the surface of the O w a m b o B a s i n g r a d u a l l y rises to 1150 m a b o v e the G r o o t f o n t e i n - M a s h a r i swell in the s o u t h w e s t , to 1250 m at the b a s e of the

African Basins. Sedimentary Basins of the World, 3 edited by R.C. Selley (Series Editor: K.J. Hsti), pp. 237-268. 9 1997 Elsevier Science B.V., Amsterdam. All rights reserved.

238

R. McG. MILLER ran rocks define the basin margins, are regionally the most extensive and are believed to underlie the whole basin. A relatively thin succession of Karoo rocks of limited extent overlies the Damara Sequence unconformably in the centre of the basin and locally along the margins. Cretaceous(?) to Tertiary deposits of the Kalahari Sequence form the final extensive sedimentary blanket that covers all but the rim exposures (Fig. 4).

The Damara Sequence

Fig. 1. Location of the Owambo Basin in western Central Africa.

rim in the west and to 1200 m at the base of the rim in the south. Very little is known about the detailed stratigraphy of most of the basin due to a thick cover of Kalahari Sequence sediments. Available information has been gleaned from mapping along the margins, from a few widely spaced wells (Fig. 2) and from seismic surveys. This paper therefore deals mainly with the late-Proterozoic Damaran rocks but also presents what is known of the younger rocks. It draws heavily on reports by Hedberg (1979) and Hugo (1969). Seismic surveys (Hedberg, 1979) suggest that the top of the granitic basement in the centre of the Owambo Basin may be as much as 7.6 km below the present surface (Fig. 3a). The Bouguer anomaly map (Fig. 3b) suggests a deep east-west axis in the southern part of the basin where the total thickness of sediments and sedimentary rocks may be even greater than 7.6 km.

The Nosib Group forms an arkosic, largely fluviatile base to the Damara Sequence. This is overlain, both conformably and unconformably, by platform carbonates of the Otavi Group within which occurs a marker mixtite, the Chuos Formation. The Mulden Group is a syntectonic molasse (Miller, 1983) that overlies the Otavi Group unconformably to paraconformably. The Damara Sequence rocks of the Owambo Basin were deposited on the stable northern platform of the Damara Orogen during phases of intracontinental rifting, spreading and continental collision (Miller, 1983) in two mobile branches of the orogen to the south and west (Fig. 1). Each major lithostratigraphic unit in the Owambo Basin has equivalents in the adjacent mobile belts (Miller, 1983). The following sedimentological descriptions are based extensively on the work of (Hedberg, 1979) but also on that of S6hnge (1957) and Martin (1965) in the Otavi Mountainland.

The Nosib Group The Nosib Group was deposited during intracontinental rifting that initiated the development of the northeast-trending branch of the Damara Orogen (Miller, 1983). The north-trending mobile belt to the west, the Kaoko Belt, is believed to have been a deep-water ocean by the end of Nosib times (Miller et al., 1983). Deposition of the Nosib Group may have began at about 900 Ma and was completed by about 730 Ma (Miller, 1983); to the south of the Owambo Basin a syenite with an age of 840 Ma intrudes Nosib sediments (Kr6ner, 1982). Two formations are distinguished, the laterally extensive Nabis Formation and the overlying, very local Varianto Formation.

REGIONAL STRATIGRAPHY

The stratigraphy and surface geology of the Owambo Basin is presented in Table 1 and Fig. 2. The Pan-African Damara Sequence, the oldest basinal unit, rests on a gneissic and granitic basement containing in-folded mid-Proterozoic cover rocks that are intruded by granites ranging in age from 1870 to 1730 Ma (Burger et al., 1976). The Dama-

The Nabis Formation. The Nabis Formation is broadly an upward-fining succession of conglomerate and feldspathic sandstone with local lenses of phyllite. A laterally extensive massive, largely clast-supported, basal conglomerate reaches 280 m in thickness. It contains subangular to subrounded pebbles and cobbles of quartzite, granite and gneiss set in a

,..] 9

9

7~ 9 7: 9

7:

to Fig. 2. Surface geology of the Owambo Basin and its elevated margins (from Miller, 1992a).

240

R. McG. MILLER

Fig. 3. (a) Contour map in metres below sea level of the top of the pre-Damara basement in the southwestern and central Owambo Basin (after compilation by Hedberg, 1979, from stacked vibroseis records). (b) Bouguer anomaly map of the Owambo basin. Contours in 10 milligal intervals. After Geological Map of South West Africa/Namibia (1980). grey to maroon feldspathic, gritty to coarse-grained sandy matrix. In the northwest, inclusion size decreases upwards. In most places, the basal conglomerate grades upwards through an interbedded transition zone into medium- to coarse-grained, moderately to poorly sorted buff to light reddish brown sandstone that is up to 1000 m thick (e.g. at Otjovazandu). Local grit and conglomerate lenses make up as much as 20% of the sandstone section and every gradation from sandstone and grit to conglomerate occurs. These e

lenses decrease in thickness and quantity upwards; pebbles in these conglomerates, mainly quartzite, also decrease in size upwards. The sandstones range from lithic arenite and feldspathic greywacke to arkose, feldspathic sandstone and quartzite. They are characterised by thin to massive and lenticular bedding, cross-bedding and by floating pebbles (mainly quartzite, minor granite and schist) that make up one per cent of the rock but decrease in quantity upwards. Bedding planes and foresets are often defined by heavy mineral laminae.

241

T H E O W A M B O BASIN OF N O R T H E R N N A M I B I A Table 1 Stratigraphy of the Owambo Basin Era

Sequence Group

Recent to Tertiary

Kalahari Sequence

Andoni Olukonda Beiseb

Karoo Sequence

Juras.-U. Trias.

Etjo

Aeolian sandstone

Lower Permian

Prince Albert Carbonaceous shale, sandstone, siltstone, low-grade coal Dwyka Tillite

Cretaceous Jurassic

Late Proterozoic Damara Sequence

Mulden Group

Subgroup Formation

Maximum thickness (m)

White sand, light green clayey sand, green clay Red sand and sticky red clay Conglomeratic, gritty, calcrete-cemented sand

274 175 50

Ombalantu

Red, semiconsolidated clay, minor sandstone

102

Kalkrand

Basalt

Owambo Kombat Tschudi

Htittenberg Tsumeb Otavi Subgroup Group (Otavi Mountainland) Elandshoek

Maieberg

Chuos

Auros Abenab Subgroup Gauss

Berg Aukas

Nosib Group

Lithology

Varianto

Nabis

Mid-Proterozoic Epupa, Huab and Grootfontein Metamorphic Complexes, Khoabendus Group, Fransfontein Granitic Suite

The sandstone becomes medium to very fine grained and better sorted upwards; bedding becomes thin and well defined as grain size decreases. In thin section, quartz grains are subangular to subrounded

90+ 137 221 158

Varicoloured shale, sandstone, siltstone, dolomite, limestone Grey to black shale, siltstone, sandstone Sandstone, siltstone, minor conglomerate, shale

2600

Thin bedded, dark grey dolomite; dark grey to black cherts, often oolitic; dark grey limestone; overlain by medium to light grey dolomites with silicified oolitic to pisolitic marker dolomites near the top Massive, light grey dolomite, brecciated and silicified in upper parts, overlain by bedded silicified dolomite, in turn overlain by bedded dolomite with silicified stromatolites Laminated, light to dark grey, argillaceous limestone overlain by light to medium grey, laminated dolomite; contorted laminations in upper part Tillite, pebbly shale and siltstone, local siliceous iron formation; thin layers of shale, siltstone, shale, dolomite Up to four cycles of thin bedded, dark grey limestone and/or tan shale overlain by grey dolomite, distinctive marker stromatolites in the dolomites Light grey buff and white massive dolomite, colloform textures common, local arkosic interbeds, black limestone and abundant chert and silified dolomite near base; brecciated, thick-bedded, medium grey, oolitic layers near top Medium to very light grey, laminated (locally massive) foetid dolomite, local concentrically laminated stromatolites; transition beds - - local shale with dark grey dolomite interbeds, overlain by dark grey limestone with local cross beds

1200

Reddish brown, poorly sorted, variously ferruginous mixtite with shaley, silty and sandy matrix; lenses of grit, sandstone, siltstone, shale Coarse- to fine-grained, thinly to massively bedded, brown-weathering, feldspathic sandstone with isolated pebbles and heavy mineral laminae; local conglomeratic and gritty lenses

823 1800

1800

2000

700

640

750

525

133

1250

Ortho- and paragneiss, metasediments, metavolcanics, granite

whereas feldspar, mainly K-feldspar, is subrounded. Fine-grained white mica is a minor constituent and magnetite and hematite make up between 1 and 10% of the rock.

242

R. McG. MILLER

Fig. 4. Geological cross section across the Owambo Basin (from Miller, 1992a).

The Varianto Formation. The Varianto Formation forms the uppermost 130 m of the Nosib Group in the Tsumeb-Kombat area. In places, there is a transitional change between the Nabis and the Varianto Formations, in others the Varianto transgresses across the Nabis to rest directly on basement. The lower part of the Varianto Formation is made up of very poorly sorted, pebbly shale and conglomerate with a sandy, highly ferruginous argillaceous matrix; better sorted interbeds of feldspathic grits, arkose, laminated shale and siltstone are present. The upper portion consists of reddish brown to maroon, ferruginous, poorly sorted conglomeratic, argillaceous siltstone and sandstone. Inclusions decrease from 5 to 2 cm in size from the base to the top and basic volcanic inclusions, probably derived from rift-related alkaline volcanism 10 km south of Kombat (Miller, 1983), become more abundant upwards. Hematite and minor magnetite constitute between 10 and 60% of the rock and thin lenses of banded iron formation occur in places. The sandstone is arkosic at the base but becomes a lithic greywacke at the top; K-spar is the dominant feldspar. The poor sorting of and variability within the Varianto Formation has been taken as evidence for a glacial origin (S6hnge, 1957) although Kr6ner and Rankama (1972) prefer to refer to the rock as diamictite.

Lithostratigraphically equivalent and possibly of similar origin are local conglomerates and conglomeratic sandstones at the top of the Nosib Group in the Ruacana area. Local rock types. The uppermost 100 m in the area northwest of the Kamanjab Inlier lacks the conglomerates of other areas. Instead, lensoid phyllite layers and a few interbedded limestone and dolomite layers suggest the zone may represent a transition to the overlying Otavi Group carbonates. In this region over 50 copper showings (chalcocite and supergene oxidation products) occur at the top of the Nosib Group or in the lower 100 m of the overlying Otavi carbonates. The setting appears to be similar to that of the Zambian Copperbelt. Depositional environment: Conglomerates were locally derived from a pre-Nosib basement of moderate relief. The upward fining suggests initial basement highs, a time-transgressive basal contact, a gradual moderation of relief and infilling of depressions. Continental conditions are suggested by the poor sorting, cross-bedding, heavy mineral laminae, the high labile content and the general absence of shales/phyllites and carbonates (Hedberg, 1979). A fluviatile environment is indicated with possible local late-Nosib glacial conditions in the regions of

THE OWAMBO BASIN OF NORTHERN NAMIBIA

243

the geanticlinal ridges (Tsumeb-Kombat and Ruacana regions) separating the Owambo Basin from the developing intracontinental rifts to the south and west. A possible transition to shallow marine conditions occurred in late Nosib times northwest of the Kamanjab Inlier.

and irregular masses up to one metre across of granular silica. Stromatolites and small-scale slump structures occur locally. Crackle brecciation of the upper part of the formation is common, the breccias being healed by sparry dolomite. A one metre thick oolite layer occurs commonly near the top.

The Otavi Group The Otavi Group was deposited on the stable northern platform of the Damara Orogen whilst the mobile belt to the south evolved from intracontinental rifting through a phase of spreading into a narrow deep-water ocean with a mid oceanic ridge. Spreading is believed to have lasted from about 730 Ma to 700 Ma (Miller, 1983). The subdivision of the Otavi Group into formations is based on detailed mapping in the Tsumeb area (Figs. 5, 8). Although certain lithological characteristics are remarkably continuous, major facies changes and the lack of detailed mapping in the west makes the correlation right across the basin, particularly within the Abenab Subgroup, tentative at best.

The Auros Formation. The Auros Formation starts with a shale or limestone at the base, is typically thin bedded and consists of up to four cycles of dark grey limestone and/or grey and tan to pink shales (up to 80 m thick) followed by light to medium grey dolomites. The basal 45 m of the lowest dolomite in the type section on the farm Auros contains irregular chert masses and bedded cherty layers that make up 20% of the section ("Jasperoid Zone" of S6hnge, 1957). This is followed by 43 m of partly silicified dolomite with chert layers in which an original oolitic texture is well preserved in the silicified zones. The uppermost 25 m of this dolomite contains preferentially silicified LLH-S, LLH-C and conophyton(?) stromatolites as well as stromatolitic domes one metre across ("Ringel" stromatolites of Krtiger, 1969). In the next of the cyclical dolomite units, clustered, vertical, unbranched, columnar stromatolites 20 cm in diameter and up to 6 m high occur ("columnar" stromatolites of Krtiger, 1969). The uppermost cyclical dolomite has very thin siliceous partings along bedding planes (silicified algal mats) overlain by cylindrical gas-escape structures 1 cm in diameter and up to 40 cm long. Many of these structures are preferentially silicified down the central core or in a circular zone around the central core ("quartz cluster" structures of the Tsumeb geologists). The proportion of limestone decreases progressively from 30% northeast of Tsumeb to zero on the western and southern edges of the Otavi Mountainland.

The Abenab Subgroup of the Otavi Mountainland The Abenab Subgroup is subdivided into three formations in the Tsumeb area and Otavi Mountainland (Table 1) where it rests with a slight angular unconformity on Nosib rocks in places. The Berg Aukas Formation. The characteristic rock type of the Berg Aukas Formation is medium to very light grey dolomite that is laminated within medium to thick bedded layers. These become thin to medium bedded at the top of the section. The dolomite at the base is foetid and contains 1-3 mm thick chert laminae. Stromatolites occur locally. In the Otavi Mountainland, the Berg Aukas Formation is extremely variable and can contain abundant terrigenous sediments which are often referred to as transition beds (S6hnge, 1957; Martin, 1965). Up to 350 m thick, these can consist of arenaceous rocks with interbedded argillaceous limestone, or arenaceous limestone that overlaps cleaner limestone onto a Nosib floor, or grey shales with dark grey pyritic dolomite interbeds overlain by very dark grey, cross-bedded limestone containing clastic sand-size grains of dolomite, limestone, chert, quartz and feldspar and rare inclusions of quartz up to 5 mm in diameter (e.g. Fig. 5). The Gauss Formation. The typical Gauss Formation dolomites are very light grey and massive. Colloform textures are common. The basal part of the formation contains up to 15% chert and the dolomite is extensively replaced by grains, patches

The Abenab Subgroup west of the Otavi Mountainland West of the Otavi Mountains, the basal Abenab begins either abruptly with detrital carbonates (eastern edge and northwest of the Kamanjab Inlier) or with a phyllite-rich transition zone (between the Kamanjab Inlier and the Otavi Mountainland, northern edge of the Kamanjab Inlier, Kaokoland). Terrigenous sediments, mainly grey and pink phyllites with fewer sandstones, quartzites, arkoses and sandy dolomites, dominate the succession where it adjoins basement highs in the extreme west but pink and grey dolomites with minor interbedded limestones, pink, tan and green shales and sandy carbonates are the main rock types further into the Owambo Basin (Fig. 5). The subgroup thins rapidly towards basement highs and is overstepped by Tsumeb Subgroup rocks

244

R. McG. MILLER

Fig. 5. Schematic stratigraphic section of the Abenab Subgroup along the rim of the Owambo Basin. Constructed from sections measured by Hedberg (1979). Not to scale (from Miller, 1992a).

along the margins of the Kamanjab Inlier and south of Ruacana. The isopach map of the Abenab Subgroup (Fig. 6) shows an axis of nondeposition extending from the northern edge of the Kamanjab Inlier (the Abenab may have been removed by pre-Chuos erosion) eastwards into a zone of minimum thickness in the northern Otavi Mountainland. In the west, this nondeposition or minimum-thickness axis bends northwestwards and follows approximately the geanticlinal ridge separating the Owambo Basin from the mobile belt to the west. The Abenab Subgroup pinches

out against basement and Nosib rocks west of Ruacana. The major depocentres were the Owambo Basin itself, central Kaokoland and the southern part of the Otavi Mountainland. Areas in which the succession is thinnest may have been regions of periodic uplift, and intraformational unconformities may be present. Depositional environment: The thickness variations, basal clastic carbonate beds, local slump structures and the pinching out of basal formations and indeed of the whole subgroup indicate (1) considerable post-Nosib relief, particularly along the margins

THE OWAMBO BASIN OF NORTHERN NAMIBIA

245

Fig. 6. Palaeoisopach map (in metres) of the Abenab Subgroup(after Hedberg, 1979). of the Owambo Basin, (2) marine transgression over an irregular surface, and (3) a time-transgressive base to the Abenab Subgroup. The Kamanjab and Kopermyn Inliers may never have been completely covered by the Abenab succession. Periodic uplift in the west is suggested by the abundance of terrigenous sedimentary components along the western margin of the Owambo Basin (Fig. 5; Hedberg, 1979). Rapid facies changes, particularly at the base of the succession, point to a variety of local environments. Initially, open marine or brackish conditions may have prevailed during which limestone was deposited. Deeper water along the margins of the basin is suggested by the fine clastic material. The laminations and thin bedding in dolomites at and near the base of the succession coupled with the paucity of stromatolites suggests quiet, saline water of intermediate depth. The Gauss Formation may have been deposited under highly saline conditions and the characteristic colloform textures may have formed after post-depositional solution of evaporite minerals. Shallow water towards the end of the Gauss period is indicated by the oolite layer. The thin bedding, oolitic textures and diagnostic stromatolites of the Auros Formation suggest quiet, shallow conditions over extensive areas towards the end of Abenab deposition. Interbedding of shale, limestone and dolomite could possibly be ascribed to salinity changes related to alternating open and restricted basin conditions. The columnar stromatolites may have formed in somewhat deeper water than the underlying LLH stromatolites.

The Tsumeb Subgroup The Tsumeb Subgroup is subdivided into four formations in the Otavi Mountainland, namely, from the base upwards, the Chuos, Maieberg, Elandshoek and Htittenberg Formations (Table 1).

The Chuos Formation. The glaciogenic Chuos Formation lies unconformably on an irregular surface of Abenab dolomites in the western Otavi Mountainland and in central and northeastern Kaokoland. South of Ruacana, the unconformity cuts deep into the Abenab rocks and in central Kaokoland, Chuos sediments fill fractures in the underlying dolomites. It is absent between the Kamanjab Inlier and the Otavi Mountainland and east of the latter. The Chuos Formation consists largely of tillite and boulder or pebble shales. In the Otavi Mountainland, clasts are commonly 1 m in size and granite boulders can be up to 2 m across. Faceted and striated clasts occur. In this region, 95% of the clasts are coarse- to fine-grained quartzites and Abenab carbonates and cherts; the remainder are granite and gneiss. The matrix is a grey, unsorted, argillaceous to silty sandstone with calcitic, siliceous and ferruginous cements. Pyrite and grains of magnetite and hematite are common. The matrix changes rapidly from schistose to arenaceous or calcareous over short distances. Local rock types in the Otavi Mountainland are a 15 m thick shale at the base, fine-grained sandstone lenses and a 1.5 m thick intraformational layer of clast-free micritic dolomite. A typical section from central Kaokoland is the following: 106 m (top): tillite with a brown to dark grey ma-

246 trix of argillaceous sandstone; subangular boulders up to 25 cm across, mainly quartzite, lesser amounts of gneiss, schist, granite, lava, dolomite; some well rounded pebbles; rock approaches pebbly siltstone at base; 196 m: maroon, argillaceous to sandy siltstone with up to 5% by volume of clasts up to 5 cm across; 35 m: green, clast-free siltstone, shale lenses; 181 m: greenish grey shale and siltstone with sandstone layers, all have floating clasts up to 5 cm across; 183 m (base): thickly to thinly bedded tillite with an argillaceous to silty matrix; unsorted, angular to subangular clasts that decrease in size upwards, 50 cm in diameter at base; proportion of Abenabderived clasts decreases upwards; local moderate sorting. Interbedded iron formation is common from the northern edge of the Kamanjab Inlier to Ruacana. Although highly ferruginous zones occur at all levels in the Chuos Formation, the best development of low-grade iron ore occurs interbedded with shale, siltstone and sandstone at the base of the formation at Ongaba in western central Kaokoland where 156 million tons grading 37% Fe and 6% Mn are present (Martin, 1965). The isopach map of the Chuos Formation (Fig. 7) shows nondeposition in the region of basement exposures on the eastern edge of the Otavi Mountainland and along an axis extending in a northeasterly direction from the eastern tip of the Kamanjab Inlier. In this region, the southern part of the Otavi Mountainland was a major depocentre. Further westwards then northwards, the Kamanjab lnlier as well as the

R. McG. MILLER basement to the west of the Owambo Basin must have been exposed to glacial erosion. Much of this debris was deposited along the western edge of the Owambo Basin. Depositional environment: A glaciogenic origin for the Chuos Formation has been strongly favoured by many authors (Gevers, 1931; S6hnge, 1957; Martin, 1965; Miller, 1980) and questioned by others (Schermerhorn, 1974, 1975). Recent detailed mapping in the central Damara Orogen has revealed abundant evidence for a glaciogenic origin in the form of unequivocal dropstones and extensive fluvio-glacial and glacio-marine reworking (Henry et al., 1986; Badenhorst, 1988). Along the margins of the Owambo Basin, the features that point to a glaciogenic origin for the Chuos Formation are: Faceted and striated pebbles, lack of sorting, massive bedding, an abundant, poorly sorted, argillaceous to sandy matrix made up mainly of angular grains, limitation of the formation to the margins of basement highs (i.e. not present in the area between the Otavi Mountainland and the Kamanjab Inlier) and, the greater abundance of the more argillaceous type of matrix deeper into the basin. The Maieberg Formation of the Otavi Mountainland. The Maieberg Formation is characterised largely by laminated argillaceous limestone overlain by a relatively thin dolomite that is laminated in the east. Local dolomite and clastic units occur. Southwards, across the southern margin of the Owambo Basin, the proportion of dolomite to limestone increases until it reaches 100% on the edge of the mobile belt.

Fig. 7. Palaeoisopach map (in metres) of the Chuos Formation (after Hedberg, 1979).

THE OWAMBO BASIN OF NORTHERN NAMIBIA

247

Fig. 8. Schematic stratigraphic section of the Tsumeb Subgroup (excluding the Chuos Formation) along the rim of the Owambo Basin. Compiled from sections measured by Hedberg (1979). Not to scale; see Fig. 4 for legend (from Miller, 1992a).

In the Otavi Mountainland, the succession consists of grey, laminated, thin bedded limestone (Zone 2 of S6hnge, 1957) capped by thin, light grey dolomite (Zone 3 of S6hnge, 1957) laminated within massive beds (Fig. 8). Characteristically, the laminations in the dolomite are contorted. The limestone contains pyritic clay and silt laminae which constitute up to 15% of the rock. Some pale green to tan beds are present and stylolites are common. Detrital carbonate grains occur at the top of the limestone.

The Maieberg Formation west of the Otavi Mountainland. A local arkosic carbonate member occurs at the base of the formation on the northern edge of the Kamanjab Inlier. This consists of a 10 m basal conglomerate and arkose followed by

laminated dark grey dolomites, in turn followed by dolomite with arkose interbeds. The lower units of the latter are cross-bedded. The arkosic carbonate unit is followed by a sequence of four members that extend from the northern edge of the Kamanjab Inlier to central Kaokoland (Fig. 8). The Rasthof Member, the lowest of these members, forms the base of the Maieberg Formation where the arkosic carbonate member is absent and consists of grey dolomite, laminated within massive beds that become medium bedded towards the top. The laminations are contorted and form 3-4 m folds that decrease in amplitude upwards. Hedberg (1979) ascribes these structures to slumping. There are local autoclastic, oolitic layers and coarsely detrital layers.

248

R. McG. MILLER

Fig. 9. Palaeoisopach map (in metres) of the Maieberg Formation (after Hedberg, 1979). LLH-C stromatolites and pillow-like algal growths occur. The dolomite is strongly foetid. The Gruis Dolomite Member is a light grey to pink and tan, massive to medium bedded dolomite that is laminated in parts and contains local well sorted clastic dolomite and limestone beds in its upper and lower parts. The main grey limestone member follows on the Gruis Dolomite and is very similar to the main limestone in the Otavi Mountainland. Quartz grains constitute up to 35% of the basal part of the limestone along the edge of the Kamanjab Inlier and this can be correlated with a 60 m thick protoquartzite, sandstone, siltstone and shale unit at the base of the limestone in central Kaokoland. The limestones contain on average between 12 and 15% argillaceous material but all gradations from pure limestone to calcareous shale occur. Some beds contain medium to coarse-grained carbonate intraclasts; autobreccias occur locally. Pink, tan and maroon beds are present in the section and become more abundant towards the top. Pseudomorphs of hematite after pyrite occur throughout. Vertical stromatolites occur near the base of the section and algal mats higher up. The thin upper grey dolomite member is laminated. Detrital dolomite grains are present and disseminated blebs of cryptocrystalline quartz up to 1 cm across occur locally. Rapid facies changes take place in the northwest. The basal limestone in the southern limb of the east-west trending synform 30 km north of Opuwa (Fig. 2) is totally replaced by dolomite in the northern limb. These carbonates are overlain by shales that grade rapidly eastwards into arkoses and felds-

pathic sandstones with thin shale and sandy, oolitic limestone interbeds. These are overlain by pink, massive to thickly bedded limestone and dolomite which, in turn, are overlain by the upper dolomite member. The isopach map of the Maieberg Formation (Fig. 9) shows two approximately east-trending axes of minimum deposition in the northern and southern Otavi Mountainland with an intervening depocentre as well as depocentres to the north and south. The northernmost of these two axes extends in a southwesterly direction to the eastern tip of the Kamanjab Inlier. A short, parallel axis of minimum deposition occurs 30 km to the northwest. In Kaokoland, the Maieberg Formation thins westwards from the Owambo Basin towards the exposed basement. The Elandshoek Formation of the Otavi Mountainland. The Elandshoek Formation consists of three units and distinguishes itself from the underlying bedded Maieberg rocks in the Otavi Mountainland by a lack of laminations, by having significantly more chert beds, by being extensively silicified and by its predominantly light grey colour. The basal massive dolomite member (Zone 4 of S0hnge, 1957) is light grey in colour. Towards the top it is brecciated, fragments being up to 30 cm in size. The breccia is cemented by secondary quartz which increases in amount upwards to a maximum of 10% of the rock as the intensity of brecciation increases. In the Otavi Mountainland, detrital, subangular to rounded, moderately well sorted grains of dolomite, chert and quartz (the latter making up to 40% of the rock in places) are

THE OWAMBO BASIN OF NORTHERN NAMIBIA readily discernible. Local rock types in the west are light grey and pink breccias, syndepositional and tectonic breccias, lenticular beds of limestone, and stromatolites. A bedded silicified dolomite member follows the basal unit (Zone 5 of St~hnge, 1957). This is distinguished by its bedded nature and by thin interbeds of light grey to reddish brown cherts and silicifled dolomites. The latter weather a reddish brown colour. The chert beds are often oolitic, vary from a few centimetres to 10 m in thickness and commonly contain detrital grains. Cherts and secondary quartz make up about 30% of this unit. In the Otavi Mountainland, the dolomites are dark grey and laminated at the base in the east but westwards the whole unit becomes massive and indistinguishable from the underlying massive dolomite member. Slumping is common throughout the unit but decreases upwards; slumped and planar beds are interbedded in the upper part. The upper 200 m are syndepositional autobreccias with fragments up to 15 cm across; this zone decreases in thickness westwards where fragments reach only 2.5 cm in size. Clay partings occur in the dolomite in the west. Algal mats are the main form of organic structure. The uppermost unit is a light grey, bedded, silicifled and cherty dolomite with silicified stromatolites (Zone 6 of S~3hnge, 1957; included in the Hiattenberg Formation by SACS, 1980). lnterbedded, medium grey dolomites that are locally foetid become increasingly more abundant towards the top. Most of the dolomite contains well sorted, detrital carbonate grains and some quartz grains. The base of this unit is a 2 m thick stromatolitic marker bed. Algal mats in this bed pass upwards into LLH-C stromatolites that are only branched in the upper 10 cm and have been identified as the form-genus conophyton (Hedberg, 1979). Three metres higher up is another highly silicified stromatolitic dolomite that grades into chert and contains slightly flattened, concentrically layered and zoned, silicified concretions that Kriiger (1969) referred to as "algal buns". Many of the cherts higher in the unit contain conophyton and other types of stromatolites. Ripple marks are present in some cherts. Cherts and silicified dolomites make up between 10 and 30% of this unit. Argillaceous, quartzose and hematitic beddingplane partings, some with muscovite, are common. Stromatolites and oolitic beds are abundant in the Kombat area; local pisolite beds and intraformational conglomerates occur in the west and northwest of the Otavi Mountainland where cherts are less abundant.

The Elandshoek Formation west of the Otavi Mountainland. West of the Otavi Mountainland, the general character of the formation is still distin-

249 guishable but the lower two members become very thin and the formation is made up mainly of the bedded dolomite member with silicified stromatolites which reaches its greatest thickness along the northern margin of the Kamanjab Inlier. Here the basal massive dolomite can be absent altogether. The detrital nature of the formation is readily apparent in this region and a hematite coating to grains gives some of the dolomites in the basal part a pink colour. One metre thick chert beds occur approximately every 5 m in these lower dolomites. LLH-S and LLH-C stromatolites are relatively common in the lower and upper parts of the formation but decrease in abundance from the inlier basinwards. The 5-1A well in the central Owambo Basin terminated 95 m below the top of the Elandshoek Formation, the rocks encountered being similar to the upper Elandshoek exposed to the south and west (Hedberg, 1979). Along the western margin of the Owambo Basin, bedding, colour, autoclastic and slump breccias, and distinctive overlying and underlying rock types enable the Elandshoek Formation to be identified. Massive bedding at the base in places suggests correlation with the basal massive dolomite member. Thinning and interbedding of terrigenous clastics in the Ruacana area and southwest thereof indicate basin margin conditions along a basement high. "Algal buns" occur at approximately the same stratigraphic level as in the type section. Local deposition in deep water is indicated by a thick succession 80 km south of Ruacana (Fig. 10) with only few stromatolites and oolite beds, minor chert, and a 150 m thick slump breccia that may be a mass flow deposit. The isopach map of the Elandshoek Formation (Fig. 10) shows the greatest thicknesses to the north and south of a sinuous east-west axis of minimum deposition through the central Otavi Mountainland. This axis extends westward to the Kamanjab Inlier, follows the inlier then bends northwestwards to follow the geanticlinal ridge separating the Owambo Basin from the mobile belt to the west. North of this region, the area west of Ruacana is shown to have been a basement high and, as with the Maieberg Formation, the region 80-100 km south of Ruacana was a major depocentre. The Elandshoek Formation thins significantly over the Kopermyn Inlier.

The Hiittenberg Formation in the Otavi Mountainland. The Htittenberg Formation in the type area is divided into a lower limestone-beating member and an upper dolomite member. The latter contains three distinctive, highly silicified marker oolitic to pisolitic dolomites near the top. The base of the formation is marked by the change in colour from the light grey of the Elandshoek Formation to dark grey. In addition, black-weathering cherts occur ei-

250

R. McG. MILLER ANGOLA (~ Ombalantu

NAMIBIA

On _t_(~)Oluk ondo

0\

Kookolond

|

Okankolo

+

OWAMBO

[~90

(~ Nonzi

(~

BASIN

OPO-I|

(~)ST-J

~ Andonl

Okosnanokano

|

o

o

i

/ i

i

i

i

5o

'

Ko=o.job1"Tg~,m"--Z. 5oo . . . .

A

J km

|

Fig. 10. Palaeoisopach map (in metres) of the Elandshoek Formation (after Hedberg, 1979).

ther at the base (northern edge of Kamanjab Inlier), up to 200 m above the base (central Kaokoland) or, very locally, just below the base (Kopermyn Inlier). In the western Otavi Mountainland, the contact between the Elandshoek and Htittenberg Formations is transitional. The lower member (Zone 7 of S6hnge, 1957) starts with a 65 m thick basal unit of dark grey dolomite with ferruginous pelitic and siliceous bedding-plane partings that weather a red or orange colour. A distinctive nodular chert bed occurs about 25 m above the base and cherts form about 5% of the section above this. Black-weathering chert occurs near the top of this dolomite. A 16 m thick transition zone follows in which the same dolomite contains thin interbeds of dark grey limestone with detrital carbonate grains and an indistinct oolitic texture. Dark grey dolomitic limestone 104 m thick overlies this transition zone. Interbedded with the limestone are a few dark grey dolomites, very dark grey chert layers and thin layers of tan to light grey, dolomitic, laminated shale and siltstone (dark grey to black in the Tsumeb Mine). The limestone is capped by 3 m of black pisolitic chert which, in turn, is overlain by light brown dolomite with interbedded dark grey limestone layers. The upper member (Zone 8 of S6hnge, 1957) starts with 75 m of medium grey, medium bedded, foetid dolomite that contains irregular chert masses and thin chert beds. This is followed by 50 m of medium to light grey dolomite with rounded dolomite fragments up to 6 mm across; algal mats occur at the base of this section. Detrital components become increasingly more abundant in the remain-

der of the upper member which consists of the same medium to light grey dolomite. The three marker beds of weather-resistant, oolitic dolomite occur in this section. They consist of coarse-grained silicified oolites set in a matrix of light grey to white, finegrained, partly silicified dolomite; the upper layer is cross-bedded. The uppermost 10 m of the Htittenberg Formation consists of thinly bedded layers of medium grey and dark grey dolomite. The medium grey dolomite is laminated and shows the detrital texture clearly. Local features in the Htittenberg Formation in the Otavi Mountainland are interbedded argillaceous sandstones near the base and top of the lower member, carbon specks in the basal foetid dolomite, an "augen limestone" in the limestone section in the Tsumeb Mine and south of Tsumeb that contains scattered radiating clusters of elongated, black sparry calcite crystals which may be secondary after anhydrite (Tsumeb Corporation, 1986), ripple marks and possible desiccation cracks in the upper member (S6hnge, 1957; Grobler, 1961), and cross-bedding that indicates currents from the north and east (S6hnge, 1957) towards a deeper basin in the south and west (Fig. 11).

The Hiittenberg Formation west of the Otavi Mountainland. Between the Otavi Mountainland and the Kamanjab Inlier, the lower member retains its character. The limestones are clearly detrital and contain fragments of lighter grey dolomites. Southwards the amount of limestone decreases significantly, a few thin brown limestone beds appear in the upper part of the formation and a 5 m thick

THE OWAMBO BASIN OF NORTHERN NAMIBIA

251

ANGOLA (~ Ombalontu

NAMIBIA

..f.(~]Olukondo

OWAMBO ~) Nonzi

~) Okankolo

BASIN

OPO-I ~)

( (~ST-I

Okasnonokana

~) Andoni

(~ Beiseb

| ~5,A

+ o

I

0

50

Kamonjab Inlier

I00 km

|

Fig. 11. Palaeoisopach map (in metres) of the Hiittenberg Formation (after Hedberg, 1979).

autoclastic dolomite forms a prominent local marker bed near the top of the formation. Many of the typical Hi.ittenberg features are present in the area north of the Kamanjab Inlier but interfingering of lithologies makes the subdivision into the units of the type area more tenuous than is implied by the schematic representation in Fig. 8. The basal, dark grey foetid dolomite with the black, generally oolitic cherts is readily distinguishable. This is followed by interbedded light to dark grey, thin to medium bedded dolomite containing carbonate clasts, oolites and up to 5% quartz grains. Light to dark grey cherts make up between 10 and 20% of the section. Limestone is very minor and pinches out entirely westwards. Stromatolites, especially the LLH-S type, become abundant in these dolomites towards the west. The lighter coloured upper dolomite and chert succession contains local pink dolomites, quartzites (with 95% quartz and chert) and limestones. Some of the latter are interbedded with the quartzites and contain coarse-grained quartz, rock fragments and pisolites. Intraformational conglomerates with subrounded clasts up to 10 cm across occur at various levels in the formation. The dark grey dolomites of the lower Htittenberg Formation and the thin to medium bedded, light and medium grey beds of the upper Hi.ittenberg Formation persist northwards into Kaokoland. The proportions of medium- to coarse-grained carbonate detritus and quartz grains increase upwards in the succession. Limestones occur only locally near the top of the formation. Stromatolites, in chert and highly silicified dolomite beds, are common in the

south but become distinctly less abundant further north. The first black cherts occur 200 m above the base of the formation in southern Kaokoland. In well 5-1A in the Owambo Basin, the Hi.ittenberg Formation was penetrated completely. It is very similar to nearby exposures on the northern edge of the Kamanjab Inlier. The isopach map of the Htittenberg Formation (Fig. 11) shows an axis of minimum thickness extending from Tsumeb to the eastern end of the Kamanjab Inlier. Thereafter the axis follows that for the Elandshoek Formation very closely. Again, the area west of Ruacana was a basement high during Htittenberg times. The formation is also somewhat thinner around the Kopermyn Inlier and in the region of well 5-1A. Depositional environment of the Tsumeb carbonates: Local conglomerates, clastic grains and beds and the contorted laminae suggest sufficient vertical relief to have facilitated early submarine erosion and slumping. The algal mats in the Rasthof Member point to shallow water initially but the laminated bedding and the paucity of stromatolites at the base of the succession (Maieberg Formation) suggest deep, quiet water initially with local algal growth in marginal regions (Rasthof Member). The high content of argillaceous material in the limestones is compatible with quiet-water deposition in a basin distal from a terrigenous source. The pyrite in the limestones points to restricted basin conditions. The greater number of algal mats and stromatolites, clastic grains and oolitic layers in the Elandshoek Formation suggest a progressive decrease in water depth with time throughout the basin, the ex-

252

R. McG. MILLER

ception being in the area 80-100 km south of Ruacana where the water remained relatively deep. In the Otavi Mountainland, the middle member of the Elandshoek Formation was deposited on a southward sloping sea floor (S6hnge, 1957; Hedberg, 1979). The abundance of stromatolites in the lower parts of the Htittenberg Formation and oolite beds in the upper parts indicate that this shallowing continued through to the end of Otavi deposition. The limestones and the thin shales in the upper parts suggest varying salinities and possibly local lagoonal conditions.

Total thickness of the Otavi Group. Figure 12 shows the greatest combined thickness of the Nosib and Otavi Groups within the Owambo Basin to be 4000 m in the southwest. This decreases to 2100 m near the centre of the basin. Along the basin rim, the Otavi Group alone is considerably thicker, reaching 6000 m between Ruacana and the Kamanjab Inlier. Minimum thicknesses in the basin rim follow an axis extending from the Tsumeb area to the Kamanjab Inlier and then from there in a northwesterly direction. This coincides more or less with the minimum thickness axes for individual formations and must have been a region for maximum syndepositional erosion, intraformational disconformities and possibly palaeokarsting. As such it could well be an important target for Mississippi-Valley-type mineralisation. This may explain why more than half the mineral deposits in the Otavi Mountainland occur in the zone where the total thickness of the Otavi rocks is less than 3000 m (600 Pb-Zn and Pb-Zn-Cu deposits known in the Otavi Mountainland).

"%:!

I

R

The Mulden Group Spreading in the Damara Orogen was followed by reversal of plate motion, ocean closure, subduction of the African plates beneath the South American plate (Kaoko Belt) and of the Kalahari Craton beneath the Congo Craton (northeast-trending belt), and finally by continental collision (Miller, 1983). The Mulden Group is a northern syntectonic molasse deposited between the Dl and D2 phases of deformation in the Damara Orogen. Derived largely from the Kaoko Belt to the west following D~ uplift (post-Dl granites dated at 650 M a Miller, 1983), the Mulden Group was folded together with the underlying Otavi rocks during the D2 event, the final major deformation event in this region. This is dated by the syn-D2 emplacement of the Tsumeb polymetallic sulphide pipe at 600 Ma (Welke et al., 1983). Proximal in the west and distal in the east, the Mulden Group lies unconformably to paraconformably on the Otavi rocks in the Owambo Basin, fills synclinal structures along the basin margins and consists of a basal Tschudi Formation overlain a middle Kombat Formation and an upper Owambo Formation. White micas from Mulden rocks in well 5-1A near the margin of the basin give a maximum age of about 540 Ma which suggests a total resetting of detrital white micas and probable growth of new white mica during the peak of Damara metamorphism (Clauer and Kr6ner, 1979). Cooling ages of 460 Ma for other mica fractions from the same samples are identical to mica cooling ages throughout the Damara Orogen (Miller, 1983).

%

ANGOLA (~ Ombalontu N A M I B I A

/

OW

BASIN ST-I ~) Andonl (~ Beiseb

;nanakona

9,o~ ~oS~a

9

,

.

~o

,oo~

Kamanjab

InUer

-ZOo0~

++

~

~

..~r

~

-,,qO00

Fig. 12. Palaeoisopach map (in metres) of the Otavi and Nosib Groups within the Owambo Basin (after Hedberg, 1979) and of the Otavi Group along the rim of the basin (compiled from Figs. 5, 6, 8, 9 and 10).

THE OWAMBO BASIN OF NORTHERN NAMIBIA

253

Table 2 Average modal composition of Tschudi Formation sandstones Stratigraphic level and location

Upper Tschudi

Lower Tschudi

West ofOtavi Mountainland Borehole 5-1A N. edge of Kamanjab Inlier, between 25 and 200 m above base N. edge of Kamanjab Inlier, basal 25 m N. edge of Kamanjab Inlier Borehole 5-1A West of Otavi Mountainland

Quartz

Rock fragments

Feldspar K-spar

Plag.

67 52

16 + - - 23

11

61

13 2

19 42 55 55

Chert

Dolomite calcite

Siltstone shale schist

4

6 ~--- 15 -----~

6

3

2

4

1

46

23

3

+--- 20 ---+ 19 12

56 5

2 ~--- 17 ------~

Accessory minerals and/or matrix

Number of samples in average

-

3

6

2

11

11

6

2

-

3 3

1

11 3

After Hedberg (1979).

The Tschudi Formation. This is only poorly exposed in synclines around the margin of the Owambo Basin and is known largely from well information. The lower Tschudi in the east consists of 30 m of dark grey shales, slates, feldspathic greywackes, marls, quartzite and, in the Tsumeb area, two 2 m thick conglomerate layers consisting of angular, light grey dolomite fragments set in a sericitic to marly sandstone matrix. In the sandstones, the main rock fragments are siltstone, shale or schist, chert and quartzite; carbonate rock fragments as found in the southwest are noticeably absent (Table 2). This unit pinches out at the western edge of the Otavi Mountainland. On the northern edge of the Kamanjab Inlier, the lower Tschudi is represented by a 14 m thick, clast-supported, chert-pebble conglomerate consisting of subangular to subrounded pebbles of Otavi cherts and carbonates (the latter represented in outcrop by voids) set in a sparse matrix of chert and quartz. Northwards into the basin, the conglomerate is underlain by maroon (locally green), argillaceous, calcareous to dolomitic siltstones and shales that thicken basinwards to 280 m in the outcrops furthest north of the inlier and to 442 m in well 5-1A where shale predominates. A few thin quartz arenite beds containing about 55% chert grains are interbedded. Rock fragments in the siltstones are predominantly chert, dolomite and limestone suggesting an extensive Otavi source in early Mulden times. Feldspar is conspicuously absent. The conglomerate thins basinwards and there is a concomitant increase in quartzite pebbles particularly towards the top of the layer. This indicates removal of the Otavi cover with time and exposure of the Nosib and pre-Damara basement. The upper Tschudi consists of light grey arkoses, feldspathic sandstones and feldspathic greywackes that reach 800 m(?) in thickness in the Otavi Moun-

tainland and almost 1000 m in well 5-1A. Locally, the rocks are red, brown or green in colour, particularly in the west. On the northern margin of the Kamanjab Inlier, bedding is thin to intermediate close to the inlier but thick to massive further into the basin. Coarse grained at the base (only the lowest 50 m north of the Kamanjab Inlier), the rocks become, somewhat irregularly, medium to fine grained upwards. Interbedded siltstones occur in the upper 100-170 m in wells 2-1 and 5-1A. Sorting is moderate. Grains are subangular to subrounded, feldspars being better rounded than quartz. K-feldspar predominates over plagioclase. Muscovite and biotite are detrital components. Quartz cement is present and is slightly more abundant in the west than in the east. Rock fragments are siltstone, shale, schist, chert and quartzite. Pebbles of Otavi chert up to 1 cm across occur in the lowest 50 m in the Tsumeb area. Since the Tschudi Formation is so poorly exposed, it is not possible to construct an isopach map for the basin margin. Seismic surveys within the central Owambo Basin show the thickest development of Tschudi rocks to be in the southwestern comer of the basin (Hedberg, 1979). The Kombat Formation. The Kombat Formation of the type area in the Otavi Valley near the southern edge of the Otavi Mountainland is a series of tightly folded, dark grey phyllites and silty to sandy phyllites that overlie the Tschudi Formation. In the southwestern part of the Owambo Basin, the formation is subdivided into three members within which are four upward-fining cycles of largely finegrained terrigenous clastic sedimentary rocks that are not know in outcrop but which were drilled though in wells 5-1A and 1-1 (Fig. 13). Seismic data indicates that the formation occurs throughout the basin but that it is largely absent from the

254

R. McG. MILLER m 1800

1600

j-j+ 5 - IA

ST,-I 9 Red &

,_-__.-_-~: -..'7....--.,~.~...--..

\grey

. ._-.:_.

\ Grey to black

1400

Owambo Formation

\

Shale: vari-coloured, dolomitic, calcareous, pyritic, micaceous, interbeds of vari-coloured, finegrained, feldspathic sandstone and grey, argillaceous, pyritic, slightly vuggy limestone and dolomite

. . . . . . . . .

. . . . . . . . . . .

'.-.-'.-----7.'.-~--

Shale: grey to black, dolomitic, calcareous, pyritic, .'-" --:-__:--_-"_-':-." micaceous~ minor intert)eds of grey, fine_"~L.L__--Z____~ grained, argillaceous, calcareous sandstone " - " - "-~" " - - ~ "

X

7 7.7.-- --.7. .

.

.

.

.

.

.

.

.

.

. . . . . . . . . . . . . . . . . .

n :D o rv tO, z LI.I El ..J :3 :E

-:- -_--2-:z: : .'.---.=-'-'.-2 Shale" vari-coloured as above with interbeds "/:-_:-."' of brown, fine-grained, argillaceous, =-'------'--micaceous, pyritic, feldspathic sandstone i.-ZL~.--~ .--~_~. and siltstone .

800

~:~ '

1

.

.

.

,,.,. m

---.--'-------." Shale and siltstone: grey to block, fissile, U- _5--_ 5 _ _ : calcareous, pyritic

Kombat Formation

"= " : - - _ L

400

".-'_.":

: ::7":-".-'.""-" BLACKHALE [

,l - - - - - - - J

"EMBER~I

-.

I

,nor ~

Dolomitic shale, argillaceous dolomite

Sandstone, siltstone

0

i1

Shale Silty, sandy shale Argillaceous sandstone, siltstone

Fig. 13. Stratigraphy of the Kombat and Owambo Formations. Left hand column compiled from well 1-1" ST-1 located 170 km east of 1-1 (after Hedberg, 1979). Facies sequences indicated.

crest of anticlines near the basin margins (Hedberg, 1979). The contact to the underlying Tschudi rocks is transitional within the Owambo Basin. The lower member is 330 m thick and consists largely of light to medium grey siltstones (Fig. 13). Sandstones form 50% of the section at the base. Minor grey to grey-brown shales are present. The colour darkens in the uppermost 86 m where some of the siltstones are also highly micaceous. The lower member can be distinguished from the underlying Tschudi rocks on the basis of neutron and gamma ray logs. The black shale member is 93 m thick in well 1-1. It has a distinctive electric log trace, is a prominent seismic reflector and is an excellent marker unit throughout the Owambo Basin. The lower 49 m consist of black to very dark grey, almost silt-free shales. Resistivity in this shale is close to zero on all electric logs run. This unit forms the top of the first upward-fining cycle (Fig. 13). The following 44 m are made up of 26 m of light to medium

grey siltstone (moderate resistivity), 7 m of dark to medium grey shale (low resistivity), 8 m of siltstone as below and an upper 3 m thick dark grey shale which also has a low resistivity. Totalling 400 m in thickness in the 1-1 well, the upper member is made up of four upward-fining cycles of grey, grey-green and green sandstone, siltstone and shale, respectively 70 m, 105 m, 75 m and 150 m thick (Fig. 13). Traces of a red colouration are present in the shales of the upper two cycles. Resistivity readings decrease upwards to a minimum at the top of the shale in each cycle. The Owambo Formation. The Owambo Formation consists of three main sedimentary units in the ST-1 well, the upper and lower containing red and grey sedimentary rocks, the middle unit being grey to black in colour (Fig. 13. The upper red and grey unit is not present in wells 5-1A and 1-1. The base of the formation is placed at the base of the first sandstone containing numerous red layers. The Owambo

255

THE OWAMBO BASIN OF NORTHERN NAMIBIA

Formation is 908 m thick in the 1-1 well and 1490 m thick in the ST-1 well. Six facies sequences make up the succession in the 1-1 well. These sequences consist of fining-upwards clastic rocks and four are capped by dolomite or limestone (Fig. 13, Table 3). Most of the wells drilled in the central Owambo Basin penetrated only the uppermost Owambo Formation (Hugo, 1969) but well ST-1 is believed to have been drilled completely through the Owambo Formation into the top of the Kombat Formation (Fig. 13). In this region, the succession consists mainly of varicoloured pyritic, micaceous shales, more or less dolomitic or calcareous, with lesser amounts of siltstone, fine-grained sandstone and occasional dolomite beds. Red to reddish brown beds are numerous in the upper half of the formation here. Ubiquitous rhombohedral casts up to 2 cm across, either void or filled with calcite and/or specular hematite, suggest the former presence of evaporite minerals (gypsum?). Figure 13 shows a tentative correlation of the Kombat and Owambo Formation over a distance of 160 km between wells 1-1 and ST-1 based on colour changes and seismic results which suggest that the Black Shale Member of the Kombat Formation could be several hundred metres below TD in well ST-1. Correlation of the individual facies sequences is not possible between these two wells. As with the upper Kombat Formation, seismic data suggest that the Owambo Formation is absent from the crests of anticlines near the basin margins

Table 3 Stratigraphy of the middle and lower units of the Owambo Formation in the 1-1 well Middle grey unit:

Sequence 6

101 m

60 m

Sequence 5

6m 26 m 45m

Sequence 4

10 m 42m

16m Sequence 3

96 m 59m

Lower red and grey unit."

Sequence 2

40 m

52m 9m

Sequence !

313 m

(Hedberg, 1979). The total thickness of the Mulden Group as interpreted from seismic surveys is given in Fig. 14. .

26 m

%

ANGOLA N A M I B I A

%

,%

Okankolo .3400 Kaokoland

4-

Andom

9~

o

so

4-

BASIN (osnanakana

I

grey dolomite and argillaceous dolomite, interbedded dark grey shale and siltstone, reddish brown and greyish green shale, grey to green, very fine-grained sandstone; red, grey and green shale, minor interbedded siltstone in the lower half, red and green, fine-grained sandstone.

After Hedberg (1979).

(~ Ombalantu

o, . . . .

light to dark grey, slightly argillaceous dolomite with minor pyrite and shale, dark grey, silty, calcareous and dolomitic shale with beds of grey to dark grey dolomite; grey, sandy limestone, dark grey, silty, dolomitic shale with beds of grey to dark grey dolomite, grey, very fine-grained, dolomitic sandstone; grey dolomite, grey to dark grey, slightly dolomitic shale with minor siltstone, grey, fine-grained, argillaceous slightly dolomitic sandstone; grey to dark grey shale and siltstone, some black shale, grey fine- to medium-grained sandstone.

~J~

. . . .

,OOkm

j

Komonjob Inlier

i

Fig. 14. Palaeoisopach map (in metres) of the Mulden Group (compiled from Hedberg, 1979).

~) Be,seb

256

R. McG. MILLER %quartz

200m

0 |

50 i

.

I00 0

%chert 50 = |

%felspors 0 50

U. Tschudi

(D .e0 L

E

i

0 C O U

!

0 r 0 9* ' -

Pre-Otavi source

IO0 m

> o

I

0

I

u l-

I I I

o .e-

.In "10

l o,

-0mI

Tschudi

-1

I

1

Fig. 15. Modal composition of the Tschudi Formation sandstones north of the Kamanjab Inlier. Progressive stripping of the Otavi cover to expose pre-Otavi rocks in the source area produces a distinct change in modal composition about 25 m above the top of the lower Tschudi (after Hedberg, 1979).

Values decrease from a maximum of 4200 m in the west to 1600 m in the south.

Depositional environment of the Mulden Group: The Mulden deposits are the product of uplift and erosion of the Kaoko Belt to the west. Some detritus was derived from the south. Figure 15 shows abundant chert and scarcely any feldspar in the lower Tschudi. The abundant carbonate and chert rock fragments in conglomerate in the west coupled with the predominance of similar lithic grains in underlying siltstones indicates that the lower Tschudi must have been derived largely from Otavi rocks. The predominance of siltstones and shales in the lower Tschudi suggest either limited uplift or distant sources. Increased rates of uplift and denudation and exposure of granitic basement to produce coarser feldspathic detritus are reflected in the arkoses at the base of the upper Tschudi. The progressive upward fining of this unit is interpreted by Hedberg (1979) as representing a period of partial stabilisation resulting either from progressive reduction of relief or greater transport distances. The Kombat and Owambo Formations appear to be a continuation of Tschudi sedimentation but the finer overall grain size suggests either less rapid erosion of the source area or still greater transport

distances. Short-lived changes in the source area or the transportation network would account for the upward-fining depositional cycles. Deposition appears to have taken place in a continental basin with reducing conditions developing at a fairly early stage (dark grey to black colours) and being succeeded by alternating oxidising and reducing conditions (varicoloured rocks). Reducing conditions set in again as the supply of terrigenous material decreased towards the end of this period and carbonates were deposited. This return to more reducing conditions may have occurred sooner in the centre of the basin than along the margins (Hedberg, 1979). Evaporitic minerals may have been deposited with the carbonates. Structure of the iate-Proterozoic rocks of the Owambo Basin

DI deformation in the branches of the Damara Orogen to the south and west at about 650 Ma (Miller, 1983) produced large-scale uplift of the western edge of the Owambo Basin and the region west thereof. Although difficult to identify, Fl folds were probably produced along and more or less parallel to the margins of the basin. F~ structures have been recognised south (Miller, 1980) and west (Guj, 1970) of the basin and in the Tsumeb Mine (Lombaard et al., 1986). Pre-Mulden erosion incised deeply into folded Otavi rocks and even through them into the underlying pre-Damara basement west of the basin (Martin, 1965) and south of the Kamanjab Inlier (Frets, 1969). D2 deformation at about 600 Ma occurred after deposition of the Mulden Group, folded the Mulden Group rocks together with the underlying Otavi rocks and produced the main structural grain of the basin. Elongate, doubly plunging F2 anticlines and synclines trend mainly east-west along the southern margin of the basin and north-south to N W - S E along the western margin (Fig. 2). These structures are fairly tight along the basin margins but become more open further into the basin. The Karoo Sequence

In the central Owambo Basin, drilling has intersected, from the base upwards, glaciogenic rocks of the Dwyka Formation, shales and coals considered to belong to the Prince Albert Formation, and aeolian sandstone of the Etjo Formation. These rest on deeply weathered rocks of the Owambo Formation. Along the western rim of the basin, isolated outcrops of Dwyka tillite fill westward-flowing glacial valleys (Fig. 2). Basaltic lavas which may be equivalent to the Kalkrand Basalt Formation are shown by aeromagnetic surveys to occur in the southeastern corner of the Owambo Basin beneath the Kalahari succes-

THE OWAMBO BASIN OF NORTHERN NAMIBIA --\%,_

,l

,~,

257

ANGOLA

(~ Ornbolontu

,(-~Olukondo I ~

~,~

0\

Kookolond

0 an o o

~,

- :o~z:o/

+/,,-

oO," o

,

OWAMB f

...-too .

/( O~ST-I ~'41 : 220:0/ Okosnonokono ( ~ ~

14VV ..,rV V V V

/~) Andoni

Beiseb

/ ~ v v v v v v v v v v

O_

2-1~

^ 30q

,~vvL~a~,~vvv

v v ~/V V VVV V ~,-';-W'

VVV

(~I-I

0~ _

\VVVVVVVV' ~',.VV \§ V V V

f~

0 i

I . . . .

50 i

. . . .

I

I,// I00 V "~" 3 J km " ~ ~

Kamanjab

Inlier

Fig. 16. Palaeoisopach map (in metres) of the Karoo Sequence in the central Owambo Basin. Numbers below each well give the respective thicknesses in metres of the Dwyka Formation, the Prince Albert Formation and the Etjo Formation. Basalts occur only in the southeast. Compiled from Hugo (1969) and Hedberg (1979).

sion (Fig. 16). They have been intersected in wells drilled further east. Etendeka Formation basalts and quartz latites that outcrop beyond the basin margin in the southwest (Fig. 2) do not occur inside the Owambo Basin. The Dwyka tillite and the overlying Prince Albert shales are lower Permian in age (Anderson and Anderson, 1985). Kalkrand basalts in the south of the country have been dated at 184 Ma (Duncan et al., in press). Distribution and isopachs of the subsurface Karoo rocks are given in Fig. 16. In the south, aeromagnetic data suggest that the Karoo Sequence is truncated by a northeast-trending fault (Fig. 16). To the east, the Karoo rocks may be cut by a north-south trending fault. The Onimwandi dolerite dyke swarm (Fig. 16) is probably of Karoo age and extends from the Angolan border to about 19~ Seismic surveys do not reveal any vertical displacement of the Karoo rocks across the swarm. No work has been carried out on the basalts east of the Owambo Basin but thicknesses vary from 30 to 90 m. Maximum thickness of the sediments drilled was 357 m (Beiseb well). Seismic surveys suggest that the Karoo Sequence was deposited on an extremely uniform surface produced during post-Damaran erosion which, further south, removed most of the Damara mountainbelt before Karoo deposition began (Miller, 1983). This surface truncates structures in the underlying Mulden Group.

The Dwyka Formation. The formation consists largely of tillite; 158 m were intersected in the Beiseb well. Clasts consist of reddish brown and white quartzite, granite, gneiss, vein quartz, siltstone, light green shale, chert, diorite, and white and grey dolomite and limestone. Averaging 1-3 cm in size but reaching 10 cm across, these clasts are set in a matrix of grey, massive, sandy mudstone. Thin, dark grey to black, laminated pyritic shales, dolomitic siltstones and local limestone are interbedded with the tillite. Some of the shales are varved. Thin cross-bedded sandstones occur locally near the base. The upper 44 m of the tillite in the Beiseb well is dark grey to black. On the western edge of the Owambo Basin, some 8 km east of Opuwa, Dwyka tillite is overlain by 100 m of tan to green-grey, very thinly bedded shales and mudstones with a distinct and very regular 2 - 4 mm lamination reminiscent of varves. The Prince Albert Formation. Following conformably on the Dwyka tillite, the Prince Albert shales are divided into a lower shale member, a middle shale and sandstone member and an upper carbonaceous shale member (Hugo, 1969). Hedberg (1979) suggests that most of these sediments have been derived from the west. Microflora from each of the three members are listed in Table 4 . The basal 3 m of the lower shale member consists of black carbonaceous dolomitic and micaceous shale and siltstone either overlain by conglomerate (ST- l) or interbedded with limestone (Beiseb). These

258

R. McG. MILLER

Table 4 Microflora of Lower Permian Prince Albert Formation shales in the ST- 1 borehole Location in section

Drilled depth

Microflora

Base of upper carbonaceous member

348 m

Vittatina cf. subsaccata Protohaploxypinus amplus Nuskoisporites gondwanensis Nuskoisporites cf. rotatus Vestigisporites methoris Taeniaesporites sp. Cycadospites sp.

Middle of middle shale and sandstone member

401 m

Nuskoisporites gondwanensis Vittatina cf. subsaccata Protosacculina cf. multistriata Vesicaspora ovata Potoniesporites sp. Vestigisporites sp. Virratriradites sp. Neoraistrickia sp. Granulatisporites sp.

Middle of lower shale member

483 m

Neoraistrickia ramosa Verrucosisporites cf. naumavai Verrusosisporites pseudoreticulatus Apiculatrisporites cf. cornutus Cirratriradites sp. Granulatisporites sp. Cycadopites sp. Fragments of Nuskoisporites? Fragments of Vestigisporites?

From Hedberg (1979).

beds are overlain by dark grey to black micaceous, calcareous to dolomitic, carbonaceous, pyritic shales with minor siltstone and rare limestone interbeds. Local light and dark lamellae suggest varves. Maximum thickness was 119 m in the Okasnanakana well. The middle shale and sandstone member reaches 100 m in thickness in a well drilled very close to the location of ST-I (Coal Commission, 1961). It consists of alternating shales similar to those in the lower member and light to greenish grey, fine-grained, silty and argillaceous sandstones. Commonly cross-bedded, the sandstones make up between 10 and 30% of the unit and are most abundant in the lower part. The upper carbonaceous member reaches its greatest thickness of 65 m in the well described by the Coal Commission (1961) and consists mainly of light to dark grey and black shales, siltstones and sandstones. These rocks contain coal fragments and rare pebbles, are cross-bedded in places and have occasional small-scale channel structures. The shales particularly are pyritic. Low-grade coal forms interbedded lenses and beds up to 7 m thick. Analyses of the coal are presented in Table 5. Three thin zones in the coal of well ST-1 carry

up to 0.05% U308. This occurs as uraniferous strontianite that fills cracks in the coal. The Beiseb and Okasnanakana wells also carried traces of uranium. Depositional environment of the Dwyka and Prince Albert rocks: Dwyka tillite and Dwyka glacial valleys in the west show ice movement to have been from east to west in northern Kaokoland and from northeast to southwest in southern Kaokoland (Martin, 1961). In southern Namibia ice flow was from north to south. Martin (1961) has proposed a large continental ice sheet in the northeast encompassing the Owambo Basin. Since the Dwyka tillite contains granite and gneiss clasts, and seismic profiles of the Owambo Basin show only thick sediments and no intrusives, the ice sheet must have extended north of the Owambo Basin into eastern Angola where granitic rocks occur (Hedberg, 1979). Hedberg (1979) suggests that the post-Dwyka sediments were deposited when reworking of till in elevated areas became much more widespread. Basin plain fluvio-glacial and fluvio-deltaic conditions in a tundra-like environment may have prevailed much as in South Africa where it is suggested the coals formed from peat deposits (Le Blanc Smith and Eriksson, 1979; Falcon et al., 1984).

The Etjo Formation. Sandstones correlated with the Etjo Formation aeolian sandstone were intersected only in the Nanzi well where they directly overlie the deeply weathered suboutcrop of the Owambo Formation. The Etjo sandstone in the Nanzi well is 137 m thick, hard, light grey to buff in colour and is well-bedded (cross-bedded?). It occurs between the depths of 257 and 394 m in the well and must have formed an isolated inselberg on the preKalahari floor of the Owambo Basin (Fig. 19). An Upper Triassic age for this sandstone is assumed in accordance with the age of the lithostratigraphically equivalent Clarence Formation sandstone in South Africa. However, it must be noted that its age could range between Upper Triassic and lowest Cretaceous since the uppermost aeolian sandstones in the lower Huab River area of NW Namibia are interbedded with the basal Etendeka volcanic rocks (Martin, 1965; Frets, 1969; Miller, 1988; Horsthemke, 1992; Ledendecker, 1992; Horsthemke et al., 1990) which have an age between 128 and 135 Ma (Milner et al., 1995; Allsopp et al., 1984; Rene et al., 1992). The absence of Middle Permian to Lower Triassic rocks. Correlates of Middle Permian to Lower Triassic rocks have not been found in the Owambo Basin. Even the basalts which may be Jurassic in age are confined to the east. At the time that deposition of the Upper Triassic(?) Etjo sandstone began on the floor of Prince Albert rocks, the Owambo Basin was a major topo-

259

THE OWAMBO BASIN OF NORTHERN NAMIBIA

Table 5 Analyses of coals from the Prince Albert Formation in the Owambo Basin Seam thickness (m)

Drilled depth (m)

Coal Commission borehole (very close to ST-1) 331 0.82 332 0.04 333 0.56 334 0.33 335 0.66 336 0.84 337 0.10 - 337

6.26

0.41 0,89

_ 343 345 354

H20 (%)

Volatiles (%)

Ash (%)

Fixed carbon (%)

Sulphur (%)

11.3 13.1 14.3 9.4 8.8 9.3 9.9 9.4 8.0 9.6 11.3 9.7 9.8 8.2 12.2 8.6 9.2

18.0 20.9 22.9 15.5 17.4 15.9 18.3 16.0 16.5 16.2 19.2 18.8 17.8 17.4 19.4 15.5 52.2

48.4 37.6 31.2 56.3 52.6 53.9 48.8 54.9 57.8 52.6 46.1 45.1 45.0 51.5 39.1 53.2 15.9

22.3 28.4 31.6 18.8 21.2 20.9 23.0 19.7 17.7 21.6 23.3 26.4 27.4 22.9 29.3 22.7 22.7

1.6 2.4 1.7 0.9 2.0 1.7 1.4 1.9 1.8 1.8 1.6 1.3 1.6 4.5 1.2 1.6 0,9

Calorific value (lb/ib) 5.2 6.5 7.2 a 4.4 5.1 4.8 5.4 4.6 4.4 4.9 5.6 5.9 5.9 5.3 6.4 5.0 5.0

Borehole ST- 1 4.65 1.23 Composite of thin seams Float recovered at SG !,58 from above composite b

333-336 336--338 338-340 340--350

4.7 nd 4.7 nd

23.6 nd 21.6 nd

46.0 51.8 46.4 21

25.7 nd 27.3 nd

nd nd nd nd

nd nd nd nd

340-350

7.4

41.0

8.1

43.5

nd

10.4 a

Beiseb Borehole 0.97 0.61 0.92 0.08 0.59 1.15 0.46 1.89

332 335 347 348 353 356 357 358

10.8 9.1 8.0 7.6 7.6 7.7 8.3 8.4

30.0 26.8 23.2 31.6 21.4 21.1 20.7 22.7

10.9 29.3 37.7 33.2 47.1 46.7 45.4 39.7

40.2 34.8 31.1 27.6 23.9 24.5 25.6 29.2

nd nd nd nd nd nd nd nd

8.7 a 7.7 a nd nd nd nd nd nd

a Best samples. All analyses by Coal Research Institute, Pretoria. b Float yield = 73.9%. After Hugo (1969).

graphic depression with the top of the Prince Albert sediments in the Etosha Pan area being between 600 and 1200 m below the present-day elevations of the basin rim in the south and west. All sedimentary units of the Karoo Sequence could well have been deposited in the Owambo Basin but would have had to have been largely removed by subsequent post-Karoo erosion. If such was the case, erosion products could only have been transported southeastwards through the breach in the Grootfontein-Mashari ridge or northwards. There is no breach in the southern and western rim deep enough to have allowed transport of basinderived Karoo erosion products to the south or west. Thick clastic wedges containing abundant fragments of recognisable Karoo rocks do not occur either southeast of the Owambo Basin (Geological Survey, 1989) or to the north in Angola or Zaire (Furon,

1963; Haughton, 1963) and there does not appear to be much good evidence for such an erosional phase. However, the Stanleyville and Loia Stages of the Jurassic-Wealdian Continental Intercalaire of Zaire are likely to have had a Karoo source. An alternative and preferred possibility is that of nondeposition of these beds and/or intra-Karoo erosion along an Upper Permian to Jurassic swell running north-south down central Namibia and possibly extending through eastern Angola into eastern or central Zaire (Fig. 17). An almost complete section occurs west of the swell in NW Angola. East of the swell in Zambia, Zimbabwe and Botswana, most major stratigraphic subdivisions of the Karoo are also present but an unconformity separates the Upper Permian to Lower Triassic Beaufort equivalents from Middle and Upper Triassic rocks. This unconformity is marked in Katanga where Lower

260

R. McG. MILLER reported despite some detailed mapping. In Namibia, the units present in these sections appear to follow each other conformably. It is this lack of clear intra-Karoo unconformities where major units of the succession are missing, together with the absence of Upper Permian to Lower Triassic sediments, that suggests that the latter may never have been present along the swell. The Huab River and Waterberg sections owe their more complete successions to accumulation in, respectively, a broad pre-Karoo depression and a half graben, each southwest trending and about 50 km wide. It is suggested that the absence of basalts in the central Owambo Basin is also due to nondeposition. The only probable dykes recognised in the basin are those of the Onimwandi dyke swarm (Fig. 16).

The Kalahari Sequence This analysis of the post-Karoo succession in the Owambo Basin is based largely on a re-examination of the cores of the following wells drilled in the late 1960s, ST-1 (Hedberg, 1979), Nanzi (well no. 9074), Olukonda (9197), Ombalantu (9262), Beiseb Pan (9296) and Okasnanakana Pan (9563) (Hugo, 1969). The Kalahari Sequence is subdivided into four formations, a basal, red, fine-grained Ombalantu Formation, a conglomeratic Beiseb Formation, a red Olukonda Formation and an upper Andoni Formation (Miller, 1992b). White, sandy calcrete is abundant close to the Otavi rocks along the basin margins. The succession is more variable than suggested by Thomas (1988). It blankets everything in the Owambo Basin and may reach 600 m in thickness.

Fig. 17. Sections of the Karoo succession in western central Africa. Only Lower to Middle Permian rocks and very minor Triassic and Jurassic rocks are present in a region extending from southern Namibia to eastern Zaire. Karoo rocks may not have been deposited in this region between the Upper Permian and Lower Triassic due to relative uplift (sections compiled from Furon, 1963; Haughton, 1963; Hugo, 1969; Hodgson, 1970; Heath, 1972; SACS, 1980; Gold Fields, 1986).

Permian rocks (Ecca) are overlain by limited local occurrences of Upper Triassic rocks (Haughton, 1963). Over the swell itself, Upper Permian and Lower Triassic rocks are not represented, i.e. in the Owambo, Waterberg, Hardap and Aussenkjer sections. Yet in these latter sections, deeply incised intra-Karoo erosional unconformities have not been

The Ombalantu Formation. This unit was included in the aeolian Etjo Formation of the Karoo Sequence by Hugo (1969) and, with considerable circumspection, by Hedberg (1979, p. 300). SACS (1980) correlated the unit with the coarse red Triassic sediments of the Omingonde Formation (immediately pre-Etjo) of the Waterberg plateau to the south (see Fig. 17 for section). As is apparent from the description that follows, the sediments are not typically Karoo. The logs given by Hugo (1969) name the various red lithologies encountered immediately above definitive Karoo and Owamboland rocks as sandstone or siltstone with varying amounts of clay, shale, clay and unconsolidated sand. The log of the Okasnanakana Pan well (9563) describes the red shale as having a gritty appearance due to abundant spherical centres of silicification. These are between 1 and 2 mm in diameter and are a common feature of all the red cores described as "sandstone" or

THE OWAMBO BASIN OF NORTHERN NAMIBIA "siltstone". In some cores these silicification centres occur within a fine filigree latticework of silicification veinlets, in others only the filigree latticework is present. The so-called red sandstones, siltstones and shales are finely laminated in places with smallscale, well developed cross-bedding being common. In places there are a few, small, irregular patches up to 10 cm thick in which the red colouration is reduced to a light grey colour. Close examination shows, however, that most of the red "sandstones", "siltstones" and "shales" are in fact red semi-consolidated but friable, variably silicified mudstones consisting almost entirely of clay. Even the silicification centres can be ground down with the fingers to a clay-sized powder. Some of the mudstone contains variable amounts of silt and sand-sized grains. These mudstones are not sufficiently indurated to warrant use of the term shale. The lack of induration and the presence of interbedded red unconsolidated sands in the Ombalantu well which were washed out in abundance with the drilling water clearly indicate that this succession is younger and less lithified than the Karoo Sequence. It is referred to in this paper as the Ombalantu Formation and apart from the above unconsolidated sands and local interbedded units described below consists largely of red mudstone. The basal two metres are pebbly in the Beiseb Pan well. In most wells, sections of the mudstone contain scattered angular frequents up to 2 cm across of pink to white very fine-grained limestone or siltstone. Thin layers of light brown sandstone and siltstone and white nodular pan limestone up to 20 cm thick are present in places, particularly in the Beiseb Pan well. Irregularly shaped nodules of white calcrete up to 4 cm across are present in places. Gypsum crystals and casts of gypsum crystals occur in the upper part of the formation in the Ombalantu well. Not known from outcrop, it is suggested that the Ombalantu Formation forms the base of the Kalahari Sequence, is Cretaceous in age and is possibly equivalent to the red, cross-bedded Kwango sandstone beds of Zaire and Angola (Furon, 1963; Haughton, 1963). The Ombalantu Formation has a broad elongate suboutcrop extending from the Andoni-Beiseb area in the southeast to Ombalantu in the northwest (Fig. 18a). The beds may also occur west of the inselberg of Etjo sandstone in the Nanzi well because the overlying Kalahari succession also contains abundant clay in this region. Deposition of the Ombalantu Formation consisted mainly of the accumulation of fine clastics in a shallow, low-energy, deltaic environment in a restricted continental basin in which there was sufficient evaporation to produce gypsum. In the Beiseb Pan area, gritty and pebbly material was introduced at an early stage from the basin margins. Thick aeolian sands accumulated marginal to the lake in the northwest.

261 Beiseb Formation. The Beiseb Formation reaches a maximum thickness of 30 m, is widespread, was intersected in all wells and represents a period of rapid and extensive input of material from the basin margins into the basin. It is generally reddish in colour but light green to white in the Nanzi and Okasnanakana wells. The formation consists of well rounded clasts of brown and grey sandstone and mudstone and grey and black chert (some oolitic) up to 12 cm in diameter that are set in a matrix of fine- to medium-grained, argillaceous, calcareous to dolomitic sandstone which is very hard where well cemented by carbonates or silica. Dolomite layers are interbedded in the ST-1 well. The lowest 5 m of the formation in the Ombalantu well contain gypsum crystals up to 5 cm long. The calcrete-cemented basal parts of the Kalahari Sequence that outcrop in the T s u m k w e - G a m area and along the Weissrand of southern Namibia may be equivalent to the Beiseb Formation. Olukonda Formation. The Olukonda Formation is a friable, poorly consolidated, reddish brown, poorly sorted, massive sand and sandstone up to 120 m thick that contains a few thin gritty and pebbly layers. In the Ombalantu well, 35 m of dark red sticky clay that becomes progressively more sandy upwards overlies the red sands. The formation has only a limited distribution and, like the Ombalantu Formation, has a broad elongate suboutcrop extending from Beiseb in the southeast to Ombalantu in the northwest (Fig. 18a).

Andoni Formation. The Andoni Formation occurs throughout the Owambo Basin as a cover to all underlying units and consists of interbedded white medium-grained sand, light greenish clayey sand and green clay. The sand, in zones between l0 and 200 m thick, is unconsolidated, slightly pyritic or hematitic and, near the top of the section, contains numerous irregularly shaped dolcrete and calcrete nodules up to 30 cm across. Silcrete nodules occur in the east and become more abundant in the northeastern part of Namibia. Sorting improves upwards in the sequence. Polished and frosted, angular to subrounded grains of quartz make up 90% of the sand; chalcedony, feldspar and chert are minor components. Burrows occur in cemented sand of the Beiseb well. Calcrete lenses occurs locally at or near the top of the Andoni Formation. The clay layers interbedded in the sand are between a few centimetres and 155 m thick (Ombalantu well). They are often sandy or silty and calcareous and are generally pyritic. Thin limestone layers up to l0 cm thick, some of which are laminated, occur interbedded in the clays. Oolitic layers between 2 and 10 cm thick and ostracod shells and

262

R. McG. MILLER

Fig. 18. (a) Regional distribution of the Ombalantu and Olukonda Formations. (b) Lithofacies of the Kalahari Sequence in the Owambo Basin. The area that contains more than 25% clay in the section is outlined. Compiled from Hugo (1969), Hedberg (1979) and Geological Survey records (1989). impressions occur in the clays of the Beiseb and ST-1 wells. Unidentified bone fragments were found in clay from the Nanzi well. More than half of the Andoni Formation consists of light green clay or sandy clay over a broad region that extends due south of Ombalantu for some 200 km. The section underlying the present-day Etosha Pan contains more than 25% clay (Fig. 18b).

A thin cover of reddish brown aeolian sand in the west and southeast may be Recent in age (SACS, 1980). The Etosha Pan. Coveting 7000 km 2, the Etosha Pan is the largest of the numerous pans that occur in the Owambo Basin. As a result of scarp retreat during several periods of erosion, several pans co-

THE OWAMBO BASIN OF NORTHERN NAMIBIA alesced to form the super Etosha Pan (Rust, 1984, 1985). The pan is set in a continental planation surface covered by an extensive red palaeosol containing three pedogenic calcrete horizons. According to Rust (1985), this surface and the palaeosol mark a continental unconformity within the Kalahari succession. The palaeosol has been reworked by aeolian and pluvial processes and the corresponding sediments occur in places. The exposed calcretes have karst features. The palaeosol falls within the region of "brown and reddish brown soils of arid and semi-arid regions" shown by d'Hoore (1964). The present pan floor is made up of evaporitic calcareous sandstones overlain by a thin layer of salt-bearing chalk. An 8 m high relict pan succession rising above the present pan floor in the east consists of basal green clay with halite layers followed successively by 1 m of oomicrite, about 1 m of oosparite with oncolites, 4 m of oomicrite and 1 m of intrasparrudite with calcrete layers; the lower oomicrite and the upper two units both have laminated and columnar LLH-C, SH-V and SS-C stromatolites (Smith, 1980). Stromatolites also occur on the present floor of the pan (Rust, 1984). Subterranean water is highly saline and the Kalahari and Karoo Sequences and the upper part of the Owambo Formation form a huge brine reservoir, particularly in the region just north of the Etosha Pan (Hugo, 1969). Analyses of brines from various depths are given in Table 6. Many pans in the Owambo Basin contain halite. A few contain thenardite and trona as well. The two Otjivalunda pans have a combined total of 190,000 tons of halite and 6.5 million tons of thenardite and trona. The ~4C dates of 33,000-28,000, 22,000-18,000 and 10,500-9,000 B P for the three pedogenic calcretes cementing the surface palaeosol (Rust, 1985) are unlikely to be the true ages and must be assessed with circumspection in view of the fact that slight solution and redeposition of exposed calcrete takes place repeatedly. Netterberg (1978) believes that calcretes contain carbonates of various ages. The oldest calcrete could be considerably older than 33,000 BE Depositional environment: The Kalahari Sequence was deposited largely under arid to semi-arid conditions. Seasonal flood waters flowing in from the north fed an extensive but shallow semi-permanent lake in the southern half of the basin. Deposition began with the introduction of red clays and silts of the Ombalantu Formation into a shallow, restricted central depression in which there was sufficient evaporation to produce gypsum. Elsewhere, a basal conglomerate was deposited and thick aeolian sands accumulated marginal to the lake. The latter appear to have been incorporated into the lake

263 deposits as fluviatile sands and silts where they are interbedded with the clays. Oxidising conditions prevailed throughout deposition of the Ombalantu Formation. A somewhat wetter period appears to have followed deposition of the Ombalantu Formation. This led to the introduction of coarse detritus from the basin margins to form the gritty to conglomeratic sandstone of the Beiseb Formation which extends across the whole basin. Extensive calcretisation of this conglomeratic sandstone followed its accumulation. Gypsum crystals up to 5 cm long in places in this unit and its calcretisation attest, however, to a continuation of the generally arid conditions that prevailed during deposition of the Ombalantu Formation. The bulk of the succession thereafter is probably aeolian in origin although the limited core recovered shows massive, structureless carbonate-cemented sand. Gritty and conglomeratic layers attest to periodic inwash of coarse detritus from the basin margins. Concomitant with deposition of the aeolian sands was the deposition of clays in a large, shallow interior lake that was fed by endorheic rivers. At times this lake covered large areas of the western and southern Owambo Basin (Fig. 18b). Subsequent to the deposition of the Ombalantu Formation, reducing conditions prevailed in the centre of the basin (greenish colour of the Beiseb Formation in the Nanzi, Okasnanakana and ST-1 wells and the absence of the red Olukonda Formation) but oxidising conditions still existed to the north, east and northwest. With time, reducing conditions spread throughout the whole Owambo Basin (Andoni Formation). Figures 18b and 19 suggest that the main drainage during deposition of the Kalahari Sequence may have been from the northwest, as it is today (Fig. 17), and the theory that the southeasterly flowing upstream portion of the Kunene River once fed into the Owambo Basin (Wellington, 1938), much as the present-day Okavango River flows into the Okavango Swamps of Botswana, is attractive. The Olukonda lake appears to have been located along the northern edge of the Ombalantu lake. However, during deposition of the green Andoni clays, the lake was much larger and was located further to the west and south, the southern part thereof being beneath the present-day Etosha Pan. River capture by the westerly flowing downstream portion of the Kunene River eventually diverted all the Kunene water away from the Owambo Basin leaving only smaller river systems flowing in from the northwest (Fig. 17). Seasonal influx of water that rapidly evaporated during the warm, dry winter and hot spring months will have produced relatively high concentrations of salts that are now present in surface pans or the underground brines.

I,,9 4~

Table 6 Analyses of brine sampled during drilling in the Owambo Basin (mg/l) Borehole

Nanzi

Nanzi

Olukonda

Olukonda

Okankolo

Okankolo

Ombalantu

Ombalantu

Okasnanakana

Okasnanakana

Okasnanakana

ST- 1

T.D.S. (180~ Sodium (as Na) Potassium (as K) Sulphate (as SO4) Chloride (as CI)

47,450 17,000 80 11,050 17,600

61,560 22,300 79 13,975 23,300

30,080 9,40 64 5,350 9,400

60,750 21,760 132 13,450 23,800

51,890 18,875 104 10,700 20,700

53,760 20,050 112 11,150 22,050

28,630 8,252 72 7,885 10,000

27,960 8252 86 7,825 9,500

101,800 40,100 420 13,700 41,250

75,905 29,750 320 8,700 34,000

42,410 16,150 152 4,350 19,800

35,480 14,000 138 3,900 16,200

915 290 139

1,340 35 21

2,400 130 51

760 1,295 1,046

1,240 85 41

1,465 250 186

250 4,750 2,048

275 4,000 1,860

11,250 72 38

8,000 15 7.5

2,750 35 13

3,525 28 10

151

14

79

249

41

64

2,702

2,140

34

7.5

22

18

140 CaCO3 CaSO4 125 MgCO3 MgSO4 664 Na2CO3 Na2SO4 16,343 28,902 NaCI 53 KCI

20

50

760 388

40

185

250 2,445

275 2,156

38 8

10

66

38

52 3,208

2,544

1,386 20,669 38,044 151

2,406 7,913 15,286 122

19,137 39,246 251

1,223 15,825 33,579 198

!,290 16,491 35,941 214

5,323 16,490 137

6,320 15,666 164

Total alkalinity (as CaCO3) Total hardness (as CaCO3) Calcium hardness (as CaCO3) Magnesium hardness (as CaCO3) _

Most probable combination of dissolved salts (mg/l)

297

31

7

13 17

11,834 20,262 67,393 801

8,454 12,867 55,588 610

2,877 6,434 32,423 290

3,733 5,768 26,714

_

Depth of borehole when sample was taken (m)

92

322

92

672

450

556

167

319

92

637

637

746

Stratigraphic unit at sample depth (formation)

Andoni

Ombalantu

Andoni

Owambo

Prince Albert

Owambo

Andoni

Ombalantu

Andoni

Owambo

Owambo

Owambo

Afterl2hr pumping test

Afterl2hr pumping test

Afterl2hr pumping test

After 60hr pumping test

Afterl0hr pumping test

Afterl2hr pumping test

Afterl2hr pumping test

Afterl2hr pumping test

Drilling water

After 10 hr pumping test

Artesian after completion

Artesian

Remarks

Analyses by Department of Water Affairs, Windhoek. After Hugo (1969).

t" t'-'

THE OWAMBO BASIN OF NORTHERN NAMIBIA

265

Fig. 19. Schematic section through the Kalahari Sequence based on the geology of the Beiseb, Okasnanakana, Nanzi and Ombalantu wells. Rust (1985) believes, however, that the present Etosha Pan does not owe its origin to inflow from the Kunene River. He postulates that the large pans developed on exposed calcrete surfaces by means of pan-edge or scarp retreat under the influence of periodic phases of pluvial erosion during seasonal rainfall. After formation of the extensive palaeosol that still covers much of the Owambo Basin, Rust (1984, 1985) advocates periods of pluvial erosion and pan-surface expansion alternating with three periods of stability during which pedogenic calcretes formed within the palaeosol and its reworked pluvial

and aeolian products. Sediments within the Etosha Pan indicate alternating subaqueous and subaerial regimes brought about by seasonal rainfall and subsequent evaporation (Smith, 1980). The cyanophytes that formed the stromatolites thrived in shallow, well lit, alkaline water.

SUMMARY

The Owambo Basin is an intracontinental sedimentary basin that is located on the northern bor-

266 der of Namibia and extends into Angola. It is floored by mid-Proterozoic granites, gneisses and infolded supracrustal rocks of the Congo Craton. The basin formed initially as the stable northern platform marginal to the evolving late-Proterozoic Damara Orogen and now contains at least 7600 rn and possibly 8800 m of sedimentary rocks and semi- to unconsolidated sediments. The Bouguer anomaly map of the basin suggests a deep east-west axis at about 18030'S. The oldest fill is fluviatile feldspathic sandstones of the early Damaran Nosib Group which was deposited during intracontinental rifting further south between 900 and 730 Ma. A brief glacial episode may have marked the end of Nosib deposition. During spreading in the Damara Orogen, between 730 and 700 Ma, up to 6000 m of platform carbonates of the Otavi Group were deposited in the basin. Initial, quiet, relatively deep-water conditions resulted in the formation of laminated dolomite with few stromatolites (Berg Aukas Formation). A gradual shallowing, possibly accompanied by the precipitation of evaporitic minerals (Gauss Formation) eventually enabled numerous stromatolites to form (Auros Formation). Changing salinities may have caused deposition of alternating limestones and dolomites. A pre-Otavi topography, particularly along the margins of the basin, was only gradually buried and is responsible for a time-transgressive base to the Otavi Group and for large-scale slumping within the basin. Major facies changes within the lower Otavi Group make it difficult to correlate across the basin. Widespread glacial conditions briefly interrupted carbonate deposition and tills, glacio-marine sediments, low-grade iron formation and thin, rare, clast-free micritic dolomites were deposited along the margins of pre-Damaran basement highs forming the edge of the basin. With the resumption of carbonate deposition during upper Otavi times, more uniform conditions developed right across the basin and many features of the type sections mapped in the east can be recognised in the west. As with the start of Otavi episode, upper Otavi deposition began in quiet, relatively deep water in a restricted basin that shallowed westwards (Maieberg Formation). General shallowing with time led to an increase in the number of stromatolites (Elandshoek Formation) and eventually to the extensive development of oolitic textures and possible local lagoonal conditions (Htittenberg Formation). A bottom topography, possibly still largely influenced by the marginal basement highs was responsible once again for extensive intrabasinal slumping. The spreading that led to the deposition of the Otavi Group and its deeper water equivalents in oceans to the south and west eventually ceased. Reversal of plate motion culminated in subduction and continental collision. The Mulden Group is a north-

R. McG. MILLER ern molasse to the Damara Orogen. It was derived from erosion of the Kaoko Belt to the west following D~ uplift and was deposited unconformably to paraconformably on the Otavi carbonates between 650 and 600 Ma. Proximal in the west and distal in the east and possibly more than 4000 m thick, the Mulden Group consists of a lower arkosic sandstone (Tschudi Formation) that fines upwards through siltstones and shales (Kombat Formation) into an upper shale and carbonate unit that contains gypsum casts (Owamboland Formation). The Otavi and Mulden rocks were folded together during the 600 Ma D2 deformation phase, the last major Damara deformation phase to affect the Owambo Basin. Metamorphic white micas from Mulden rocks near the basin margins formed at about 540 Ma and give cooling ages of 460 Ma. The basal Karoo sediments of the central Owambo Basin rest on a relative fiat post-Damara erosive surface. A major Lower Permian ice sheet to the north of the Owambo Basin and possibly coveting much of the Owambo Basin as well deposited tills in the basin and in westerly flowing, glacial valleys west of the basin (Dwyka Formation). Fluvio-glacial and fluvio-deltaic basin plain shales and peat deposits (Prince Albert Formation) overlie the Dwyka rocks. Upper Triassic aeolian sandstones of the Etjo Formation appear to be the only other Karoo sedimentary rocks to have been deposited in the Owambo Basin. Basalts which may be equivalent to the Jurassic Kalkrand Formation occur northeast of Tsumeb. The thickest Karoo section intersected in the basin was 360 m thick. Cretaceous to Recent terrestrial deposits of the Kalahari Sequence form the final filling to the Owambo Basin and accumulated under arid to semi-arid conditions. The upper Kalahari sediments blanket all other units with the exception of those forming the folded and elevated margins of the basin. The Ombalantu Formation is a succession of red, well-bedded, friable, partly silicified, semi-consolidated Cretaceous(?) mudstones that forms the base of the Kalahari Sequence in the centre of the Owambo Basin. This is followed by a calcrete-cemented conglomeratic sandstone (Beiseb Formation) which, in turn, is overlain by possibly as much as 500 m of aeolian sands with varying amounts of lacustrine clays (Olukonda and Andoni Formations) that were probably introduced by endorheic rivers flowing from the northwest into a large, shallow, semi-permanent lake similar to the present-day Okavango Swamps of Botswana. Evaporite minerals occur in places and the region north of the Etosha Pan is a large underground brine reservoir. The Etosha Pan is a super pan formed by fusion of several pans as a result of pan-edge retreat during periods of pluvial erosion of the present surface.

THE O W A M B O BASIN OF N O R T H E R N N A M I B I A

ACKNOWLEDGEMENTS

I would like to thank John Ward for numerous useful discussions and exchanges of ideas on the Mesozoic and Cainozoic rocks and sediments.

REFERENCES Allsopp, H.L., Bristow, J.W., Logan, C.T., Eales, H.V. and Erlank, A.J., 1984. Rb-Sr geochronology of three Karoo-related intrusive complexes. Spec. Publ. Geol. Soc. S. Afr., 13: 281287. Anderson, J.M. and Anderson, H.M., 1985. Palaeofiora of Southern Africa. Balkema, Rotterdam, 423 pp. Badenhorst, F.P., 1988. The stratigraphy of the Chuos mixtite in part of the southern Central Zone of the Damara Orogen, South West Africa. Commun. Geol. Surv. S.W. Afr./Namibia, 4:103-110. Burger, A.J., Clifford, T.N., and Miller, R. McG., 1976. Zircon U-Pb ages of the Fransfontein Granitic Suite, northern South West Africa. Precambr. Res., 3: 415-431. Clauer, N., and Kr/Sner, A., 1979. Strontium and argon isotopic homogenization of pelitic sediments during low-grade regional metamorphism: the Pan-African upper Damara Sequence of northern Namibia (SWA). Earth Planet. Sci. Lett., 43: 117131. Coal Commission, 1961. Interim report of the Coal Commission of South West Africa. Rep. Administration for South West Africa, Windhoek, 124 pp. (unpublished). Duncan, R.A., Hooper, P.L., Lehacek, J., Marsh, J.S. and Duncan, A.L., in press. The timing and duration of the Karoo igneous event, southern Gondwana. J. Geophys. Res. Falcon, R.M.S., Pinheiro, H.G. and Shepherd, P., 1984. The palynobiostratigraphy of the major coal seams in the Witbank Basin with lithostratigraphic, chronostratigraphic and palaeoclimatic implications. Q. News Bull. Geol. Soc. S. Afr., 28: 36-38. Frets, D.C., 1969. Geology and structure of the Huab-Welwitschia area, South West Africa. Bull. Precambr. Res. Unit, Univ. Cape Town, 5:235 pp. Furon, R., 1963. Geology of Africa. Oliver and Boyd, Edinburgh, 377 pp. Geological map of South West Africa/Namibia, 1980. Scale 1 : 1 000000. Geol. Surv. S.W. Afr./Namibia, Windhoek. Geological Survey, 1989. Well logs. Geol. Surv. S.W. Afr./ Namibia, Windhoek (unpublished). Gevers, T.W., 1931. An ancient tillite in South-West Africa. Trans. Geol. Soc. S. Afr., 34: 1-17. Gold Fields., 1986. Report on the diamond drilling on prospecting grants M46/3/1241, 1583, 1584 and 1585. Otjiwarongo Karoo Basin. Rep. Gold Fields, Namibia, Windhoek (unpublished). Grobler, N.J., 1961. The geology of the western Otavi Mountainland, South West Africa. M.Sc. Thesis, Univ. Orange Free State, 119 pp. Guj, P., 1970. The Damara mobile belt in the south-western Kaokoveld, South West Africa. Bull. Precambr. Res. Unit, Univ. Cape Town, 18:168 pp. Haughton, S.H., 1963. The Stratigraphic History of Africa South of the Sahara. Oliver and Boyd, London, 365 pp. Heath, D.C., 1972. Die Geologie van die Sisteem Karoo in die Gebied Mariental-Asab, Suidwes-Africa. Mem. Geol. Surv. S. Afr., 61:35 pp. Hedberg, R.M., 1979. Stratigraphy of the Ovamboland Basin, South West Africa. Bull. Precambr. Res. Unit, Univ. Cape Town, 24:325 pp.

267 Henry, G., Stanistreet, I.G. and Maiden, K.J., 1986. Preliminary results of a sedimentological study of the Chuos Formation in the Central Zone of the Damara Orogen; evidence for mass flow processes and glacial activity. Commun. Geol. Surv. S.W. Afr./Namibia, 2: 75-92. Hodgson, F.D.I., 1973. Petrography and evolution of the Brandberg intrusion, South West Africa. In: L.A. Lister (Editor), Symposium on Granites, Gneisses and Related Rocks. Spec. Publ. Geol. Soc. S. Afr., 3: 339-343. Horsthemke, E., 1992, Fazies der Karoosedimente in der HuabRegion, Damaraland, NW-Namibia. Gtittinger Arb. Geol. Pal~iont., 55:102 pp. Horsthemke, E., Ledendecker, S. and Porada, H., 1990. Depositional environments and stratigraphic correlation of the Karoo Sequence in northwestern Damaraland. Commun. Geol. Surv. Namibia, 6: 63-73. d'Hoore, J.L., 1964. Soil map of Africa. Scale 1 to 5 000000. Explanatory monograph. Commission for Technical Co-operation in Africa, Joint Project I 1,205 pp. Hugo, P.J., 1969. Report on the core-drilling programme in Owamboland 1967-1968. Rep. Geol. Surv. S.W. Afr., 46 pp. (unpublished). Krtiner, A., 1982. Rb-Sr geochronology and tectonic evolution of the Pan-African Damara belt of Namibia, south-western Africa. Am. J. Sci., 282: 1471-1507. Krtiner, A. and Rankama, K., 1972. Late Precambrian glaciogenic sedimentary rocks in southern Africa. Bull. Precambr. Res. Unit, Univ. Cape Town, 11:37 pp. KrUger, T.L., 1969. Stromatolites and oncolites in the Otavi Series, South West Africa. J. Sediment. Petrol., 39: 10461056. Le Blanc Smith, G. and Eriksson, K.A., 1979. A fluvioglacial and glaciolacustrine deltaic depositional model for PermoCarboniferous coals of the northeastern Karoo Basin, South Africa. Palaeogeogr., Palaeoclimatol., Palaeoecol., 27: 67-84. Ledendecker, S., 1992. Stratigraphie der Karoosedimente der Huabregion (NW-Namibia) and deren Korrelation mit zeit~iquivalenten Sedimenten des Paran~beckens (Stidamerika) und den Grossen Karoobeckens (Stidafrika) unter besonderer Berticksichtigung der iiberregionalen geodynamischen und klimatischen Entwicklung Westgondwanas. Gt~ttinger Arb. Geol. Pal~iont., 54:87 pp. Lombaard, A.F., Gtinzel, A., Innes, J. and KrUger, T.L., 1986. The Tsumeb lead-copper-zinc-silver deposit, South West Africa/Namibia. In: C.R. Anhaeusser and S. Maske (Editors), Mineral Deposits of Southern Africa, Vol. 2. Geol. Soc. S. Afr., Johannesburg, pp. 1761-1787. Martin, H., 1961. The hypothesis of continental drift in the light of recent advances of geological knowledge in Brazil and in South West Africa. Alex L. du Toit Mem. Lect. No. 7, Annex. Trans. Geol. Soc. S. Afr., 64:47 pp. Marin, H., 1965. The Precambrian Geology of South West Africa and Namaqualand. Precambr. Res. Unit, Univ. Cape Town, 159 pp. Miller, R. McG., 1980. Geology of a portion of central Damaraland, South West Africa/Namibia. Mem. Geol. Surv. S. Afr., S.W. Afr. Ser., 6:78 pp. Miller, R. McG., 1983. The Pan-African Damara Orogen of South West Africa/Namibia, In: R. McG. Miller (Editor), Evolution of the Damara Orogen of South West Africa/Namibia. Spec. Publ. Geol. Soc. S. Afr., 11: 431-515. Miller, R. McG., 1988. Geological map 2013, Cape Cross, scale 1:250000. Geol. Surv., S.W. Afr./Namibia, Windhoek. Miller, R. McG., 1992a. Hydrocarbons. In: Mineral Resources of Namibia. Geol. Surv. Namibia, Windhoek, pp. 7.3-1-7.3-19. Miller, R. McG., 1992b. The Etjo and Kalahari sediments of the Owambo Basin. Abstracts, Kalahari Symp., Geol. Soc. Namibia, Windhoek, pp. 42-51.

268 Miller, R. McG., Freyer, E.E. and H~ilbich, I.W., 1983. A turbidite succession equivalent to the entire Swakop Group. In: R. McG. Miller (Editor), Evolution of the Damara Orogen of South West Africa/Namibia. Spec. Publ. Geol. Soc. S. Afr., 11: 65-71. Milner, S.C., le Roex, A.P. and O'Connor, J.M., 1995. Age of Mesozoic igneous rocks in northwestern Namibia, and their relationship to continental breakup. J. Geol. Soc. London, 152: 97-104. Momper, J.A., 1982. The Etosha Basin re-examined. Oil Gas J., April 5th, pp. 262-287. Netterberg, F., 1978. Dating and correlation of calcretes and other pedocretes. Trans. Geol. Soc. Afr., 81:379-391. Rene, P.R., Ernesto, M., Pacca, I.G., Coe, R.A., Glen, J.M., Prrvot, M. and Perrin, M., 1992. The age of Paran~i flood volcanism, rifting of Gondwanaland, and the Jurassic-Cretaceous boundary. Science, 258: 975-979. Rust, U., 1984. Geomorphic evidence of Quaternary environmental changes in Etosha, South West Africa/Namibia. In: J.C. Vogel (Editor), Late Cainozoic Palaeoclimates of the Southern Hemisphere. Balkema, Rotterdam, pp. 279-286. Rust, U., 1985. Die Entstehung der Etoschapfanne im Rahmen der Landschaftsentwicklung des Etosha Nationalparks (nrrdliches Stidwestafrika/Namibia). Modoqua, 14: 197-266. SACS, South African Committee for Stratigraphy, 1980. L.E. Kent (Compiler), Stratigraphy of South Africa. Part 1. Lithostratigraphy of the Republic of South Africa, South West Africa/Namibia, and the Republics of Bophutatswana,

R. McG. M I L L E R Transkei and Venda. Handb. Geol. Surv. S. Afr., Vol. 8, 690 PP. Schermerhorn, L.J.G., 1974. Late Precambrian mixtites: glacial and/or nonglacial? Am. J. Sci., 274: 673-824. Schermerhorn, L.J.G., 1975. Tectonic framework of Late-Precambrian supposed glacials. In: A.E. Wright and F. Moseely (Editors), Ice Ages, Ancient and Modern. Spec. Issue Geol. J., 6:241-274. Smith, A.M., 1980. Lacustrine stromatolites of the Etosha Pan, S.W.A. (Namibia). M.Sc. Thesis, Univ. Natal (Durban), 140 pp. (unpublished). Srhnge, P.G., 1957. Revision of the geology of the Otavi Mountainland, South West Africa. Rep. Tsumeb Corporation Limited, Tsumeb (unpublished) Thomas, D.S.G., 1988. The nature and depositional setting of arid and semi-arid Kalahari sediments, Southern Africa. J. Arid Environ., 14: 17-26. Tsumeb Corporation., 1986. A field introduction to the geology of the Otavi Mountainland, northern SWA/Namibia. Rep. Workshop on Precambrian Carbonate Sedimentology, Tsumeb Corporation, Tsumeb, 29 pp. (unpublished). Welke, H.J., Allsopp, H.J. and Hughes, M.J., 1983. Lead isotopic studies relating to the genesis of the base metal deposits in the Owambo Basin, Namibia. In: R. McG. Miller (Editor), Evolution of the Damara Orogen of South West Africa/Namibia. Spec. Publ. Geol. Soc. S. Afr., 11: 321. Wellington, J.H., 1938. The Kunene River and the Etosha Plain. S. Afr. Geogr. J., 20:21-33.

Chapter 12

The Foreland Karoo Basin, South Africa

M.R. JOHNSON, C.J. VAN VUUREN, J.N.J. VISSER, D.I. COLE, H. DE V. WICKENS, A.D.M. CHRISTIE and D.L. ROBERTS

INTRODUCTION The Karoo succession of South Africa attained scientific prominence during the 19th Century on account of its rich terrestrial vertebrate fauna, distinctive Glossopteris flora, unrivalled glacial deposits (in places overlying classic glacial pavements) and the spectacular Drakensberg flood basalts and associated network of dolerite dykes and sills. The basin also assumed major economic importance with the discovery and exploitation of its extensive coal deposits. The main Karoo Basin contains strata which range in age from Late Carboniferous to Middle Jurassic and attain a maximum cumulative thickness of approximately 12 km (Johnson, 1976) in the southeastern portion of the basin towards the eastern end of the Karoo Trough (Fig. 1). The basin covers an area of approximately 700,000 km 2 but it was more extensive during the Late Carboniferous to Permian, when it formed one of the major depocentres in southwest Gondwana with an area of at least 1,500,000 km 2 (Visser, 1987). The existence of the Gondwana supercontinent was documented in detail by Du Toit (1937) who noted, for example, that the Permo-Carboniferous glacial deposits, Permian Glossopteris flora and certain Permo-Triassic tetrapods of the Karoo Basin are also present in South America, the Falkland Islands, Madagascar, India and Australia. The Karoo Basin is largely underlain by a stable floor comprising the Kaapvaal Craton and the Namaqua-Natal Metamorphic Belt (Fig. 1). It is bounded along the south by a fold-thrust belt (Cape Fold Belt) and along the east by a monoclinal downwarp. The Cape Fold Belt partly overlies less competent crust below the geophysically defined "Southern Cape Conductive Belt" (Fig. 1), interpreted by some as serpentinized basalt that was

obducted against the Namaqua-Natal Belt at approximately 800 Ma (De Beer et al., 1982). Following deformation and intrusion by granites (600-500 Ma) in the south, it was the site of subsidence and deposition of the Ordovician-Early Carboniferous Cape Supergroup, a precursor of Karoo Basin sedimentation. The Cape Supergroup consists of a passive margin clastic wedge up to 8 km thick that was derived from a northern cratonic provenance (Johnson, 1991). Similarly, the "Natal Trough" along the eastern basin margin is thought to be a pre-Karoo feature, originating as an aborted rift in which the early Palaeozoic Natal Group accumulated (Tankard et al., 1982). The Karoo Basin is filled with clastic and subordinate igneous rocks belonging to the Karoo Supergroup. The Dwyka Group forms the basal part of the succession and consists of diamictite and other glacial-related rock types deposited during the Late Carboniferous and Early Permian. The deposits are overlain by the mudrock-dominated Permian Ecca Group, generally representing suspension settling. Simultaneously, fluvio-deltaic sand prograded into the northeastern part of the basin while submarine fan and basin plain turbidite sand and silt as well as deltaic mud and sand were deposited in the Karoo Trough in the south. Extensive fluvial channel sand and flood basin/lacustrine mud were deposited during the Late Permian and Triassic (Beaufort Group, Molteno and Elliot Formations). By the early Jurassic the Karoo Basin was located between latitudes 30 ~ and 50 ~ (Visser, 1991), with desert conditions resulting in the deposition of fine-grained aeolian sand and silt and playa lake mud of the Clarens Formation. Infilling of the Karoo Basin was finally completed during the Early Jurassic by the outpouring of at least 1400 m of basaltic lava (Drakensberg Group), accompanied by the widespread intrusion of do-

African Basins. Sedimentary Basins of the World, 3 edited by R.C. Selley (Series Editor: K.J. Hsti), pp. 269-317. 9 1997 Elsevier Science B.V., Amsterdam. All rights reserved.

270

M.R. JOHNSON et al.

Fig. i. The Karoo Basin in its tectonic setting, showing the position of the cross-section depicted in Fig. 3. lerite dykes and sills into the underlying sedimentary pile.

STRATIGRAPHY AND PALAEOENVIRONMENTS

The stratigraphy and depositional history of the Karoo Supergroup have been reviewed by Tankard et al. (1982) and Smith (1990). Detailed studies of parts of the succession are referred to in the descriptions which follow. The major lithostratigraphic units of the Karoo Supergroup crop out concentrically around the basin (Fig. 2). This pattern reflects both the steep northward dip of strata along the southern margin of the basin (Fig. 3) and the gentler centripetal dips elsewhere. Sections portraying the stratigraphic succession in different parts of the basin are contained in Fig. 4. Lateral facies changes, particularly in the lower half of the succession, have given rise to intertonguing of formations in various parts of the basin (Figs. 2, 4). No significant unconformities are known to exist in the basin, with the possible exception of one at the base of the Molteno Formation.

Dwyka Group

The lower contact relationship of the glaciogene Permo-Carboniferous Dwyka Group is highly variable over the basin. Glaciated Precambrian bedrock surfaces are commonly found along the northern basin margin whereas a thin (1.5 m) sedimentary unit which has about the same lithological composition as the underlying early Palaeozoic bedrock and often contains soft-sediment glacial pavements occurs along the western and eastern basin margins. In the south, where the group mainly overlies mudrock, brecciation and deformation of the rocks on the contact took place. The Dwyka Group shows distinct lithological differences over the basin which led to the recognition of a northern and southern facies (Visser, 1986) (Fig. 5). The northern facies (Prieska, Virginia and Louwsburg stratigraphic sections in Fig. 6) mainly belongs to the Mbizane Formation and is characterised by rapid thickness changes (up to 200 m variation within short distances), a highly variable lithology, and a low massive diamictite (~20%) and high mudrock (~40%) content (Visser,

9 t" 7:

9 9

7: 9

Fig. 2. Areal distribution (schematic) of lithostratigraphic units in the Karoo Basin and location of sections A-E (Fig. 4) and 1-3 (Fig. 23).

t',9

272

M.R. JOHNSON et al.

Fig. 3. North-south section across the Karoo Basin. See Fig. 1 for location.

1986). Six lithofacies can be defined (Fig. 6). The massive diamictite facies consists mostly of highly compacted, clast-rich diamictite with rounded to angular, striated stones, up to 1.5 m across, which often reflect the composition of the underlying or surrounding bedrock. The facies is interpreted as lodgement or melt-out deposits which formed in subglacial positions. The stratified diamictite facies consists of diamictite with poorly to well-defined bedding planes or a rapid alternation of diamictite, mudrock, sandstone and conglomerate beds in which diamictite is the predominant constituent. The diamictite beds contain evidence of slumping, basal shear structures, vertically orientated clasts, and large clasts truncating several beds. The facies mostly formed by sediment gravity flow, although intermittent reworking of subglacial diamictons and rain-out of glacial debris also occurred during deposition. The conglomerate facies ranges from single-layer boulder beds to poorly sorted pebble and granule conglomerate. The upper contacts of the boulder beds (e.g. in the Prieska section) are often abraded to form striated boulder pavements. These are interpreted as lodgement deposits, whereas the poorly sorted conglomerates formed by water-reworking of diamicton and by high-density sediment gravity flows. The sandstone facies consists of either fine- to medium-grained, massive to ripple-laminated, or medium- to coarse-grained, trough cross-bedded, immature sandstones. The massive to ripple-laminated sandstones, often interbedded in mudrocks, are interpreted as turbidite deposits, whereas the coarsegrained cross-bedded types formed by tractional fallout from outwash streams (Von Brunn, 1981).

The mudrock with stones facies varies from a well-laminated mudrock or rhythmite with dispersed stones and granules deforming the laminations to a massive mudstone with numerous clasts (boulder mudstone). Viscous flow deformation and deformed outwash deposits are often associated with the latter type. The facies represents rain-out deposits with the variation in textural character dependant on the rate of rain-out and the location of the depository relative to the ice front. The mudrock facies consists of dark-coloured, often carbonaceous mudstone, shale or silty rhythmite. The facies is a product of suspension settling of mud as well as fall-out of silt from sedimentladen underflows. Anderson and McLachlan (1976) reported spores, pollen and plant remains from the mudrocks and the interbedded diamictite while Anderson (1981) noted the presence of arthropod trackways and fish trails on bedding planes. The southern facies (Elandsvlei and Klaarstroom stratigraphic sections in Fig. 6) constitutes the Elandsvlei Formation and is characterised by a regular increase in thickness towards the south (from about 100 m to 800 m), a fairly uniform lithology, and a high massive diamictite (~70%) and low mudrock (-~8%) content (Visser, 1986). Seven lithofacies are defined (Fig. 6). The massive diamictite facies consists of highly compacted, clast-poor diamictite often showing shear structures. The angular to subangular stones, up to 2 m across and often striated, are distantly derived except for those in the lowermost 1-2 m of the facies. The facies is interpreted as lodgement or melt-out deposits which formed in a subglacial position. The facies consisting of massive diamictite with deformed sandstone bodies differs from the massive

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t,9 Fig. 4. Generalised stratigraphy and lithology of the Karoo Supergroup in the Karoo Basin. For location of sections see Fig. 2.

274

M.R. JOHNSON et al.

STRATIGRAPHIC SECTIONS

!

9 Prince Albert and Whitehill Formations

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Fig. 5. Distribution of the Dwyka Group (southern and northern facies) and the Prince Albert and Whitehill Formations and location of the columnar sections in Figs. 6 and 7. diamictite facies only by the presence of lens-like to highly irregular and deformed sandstone bodies. The fine-grained wackes and medium- to coarse-grained quartz arenites exhibit horizontal bedding as well as cross-bedding. The sandstones formed by tractional fall-out from streams and represent esker, esker-delta and subaqueous outwash-fan deposits (Visser et al., 1987). The massive carbonate-rich diamictite facies, which is clast-poor, contains only small angular stones, concretions and irregular bodies of carbonate rock, and often lenses of mudrock with dispersed stones. The facies formed by the rain-out of debris, and the carbonate probably originated by crystallisation from interstitial waters. The carbonate concentrated in the cold waters by brine formation below an ice shelf where freeze-on of sea water took place (Visser, 1983). The stratified diamictite facies consists mainly of diamictite with poorly to well defined bedding planes; laminations can sometimes be seen in the argillaceous matrix. The facies mainly represents rain-out deposits with minor resedimentation in the proximal iceberg zone. The conglomerate facies represents single-layer and massive boulder beds similar in character to those of the northern facies. The mudrock with stones facies consists of a well-laminated shale or rhythmite with dispersed stones and granules deforming the laminations. The

facies represents rain-out deposits which formed in the distal iceberg zone. The mudrock facies is similar in character to that of the northern facies. Spores and acritarchs have been reported from the interglacial mudrocks (Anderson, 1977). The northern facies of the Dwyka Group represents predominantly valley fill deposits left by retreating glaciers (Visser, 1983). The southern facies represents thick platform deposits of diamicton (Visser, 1983) in which up to seven genetic increments left by ice lobes entering the basin from the north, south and east can be recognised. In the beginning debris eroded from the highlands was deposited by a grounded ice sheet, but near its unstable margin in the west (Theron and Blignault, 1975; e.g. Elandsvlei section in Fig. 6) fluctuations of the ice front deposited bedded diamictons and subglacial and subaqueous outwash sediments (Visser et al., 1987). A change to a subpolar climate caused the formation of predominantly floating ice where rain-out debris accumulated near the grounding line of large ice shelves and in the proximal iceberg zone (Gravenor et al., 1984; Visser, 1989). Disintegration of the ice sheet started about 280 Ma ago with glacial deposition then confined to valleys and low-lying areas along the cratonic highlands. Along the highland slopes and in the drowned valleys tidewater glaciers and debris-loaded icebergs left a thick sequence of massive diamicton, berg-

THE FORELAND KAROO BASIN, SOUTH AFRICA

275

LOUWSBURG

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stone mud, subaqueous outwash sediment and mud (Visser, 1982). Ecca Group The Permian Ecca Group comprises a total of 16 formations, reflecting the lateral facies changes that characterise this succession (Figs. 2, 4). Except for the fairly extensive Prince Albert and Whitehill Formations, the individual formations can be grouped into three geographical areas for descriptive purposes (southern, western + northwestern and northeastern areas). In addition, there is a relatively small area along the eastern flank of the basin, between the southern and northeastern outcrop areas, where the Ecca Group consists of 600-1000 m of undifferentiated mudrock which has not yet been studied in detail.

Basal formations in the south, west and northwest Prince Albert Formation. The Prince Albert Formation is confined to the southwestern half of the Karoo Basin (Fig. 5). Towards the northeast it thins and locally pinches out against the basement or merges into the Vryheid and/or Pietermaritzburg Formations (Cole and McLachlan, 1991; Cole, 1994) Along the western and southern outcrop belt the thickness is highly variable (40-150 m) and, with the help of borehole data, two areas of maximum thickness (up to 300 m) can be identified (Fig. 5). Comparison of the thickness variation with an isopach map of the Dwyka Group and the lower contact relationships (Visser, 1982, 1987) suggest a correspondence with the morphology of the upper Dwyka surface as well as a lateral facies change from shale to sandstonebeating diamictite in an eastern direction.

276 It is possible to recognise a northern and a southern facies in the formation (Fig. 5). The northern facies is characterised by the predominance of greyish to olive-green micaceous shale and grey silty shale, as well as a pronounced transition between the underlying glacial deposits and the mudrocks. In addition to the two above-mentioned rock types, dark-grey to black carbonaceous shale and fine- to medium-grained feldspathic arenite and wacke are also present. Cross-bedding and slump structures occur in the sandstones and parallel and ripple lamination in the silty shale. Brownish calcareous concretions and irregular carbonate bodies are present in both the sandstones and mudrocks. The lower contact in the east consists of a transition zone, up to 15 m thick, of mudrock with dispersed stones and granules and rhythmites. In the west the transition attains a maximum thickness of 80 m and consists of an upward-fining sequence of sandstones, siltstones, silty shales and rhythmites (Fig. 7), often containing ice-rafted debris. The upper contact with the Whitehill Formation commonly overlies a thin upward-coarsening sequence grading from carbonaceous mudrock into silty shale (Fig. 7). Marine fossils (cephalopods, lamellibranchs and brachiopods), as well as plant and palaeoniscoid fish remains and coprolites, have been recorded from near the base of the formation at Douglas (McLachlan and Anderson, 1973). Sedimentation of the mudrocks took place by suspension settling of mud from under- and overflows. The silt and sand were deposited by tractional fallout from turbidity currents entering the basin mainly from the north. In the Boshof area, where the silt and sand content is high, deltaic deposits have been recognised (Cole and McLachlan, 1991). The southern facies is characterised by the predominance of a dark-grey, carbonaceous, pyritebearing, splintery shale and the presence of darkcoloured chert and phosphatic nodules and lenses. Greenish and dark-brown mudrocks are also present where the formation is thick. The lower contact with the Dwyka Group is sharp, with a thin (1 m) black carbonaceous shale at the base. From east of Klaarstroom massive greenish-grey wacke and thin arenite beds are present near the base of the formation. Dark bluish-grey chert nodules and lenses are distributed throughout the formation, although a chertrich zone near the base of the sequence can be traced from the west to a position north of Port Elizabeth. Phosphatic nodules, up to 0.3 m across and consisting of amorphous collophanite and carbonaceous material, and lenticular phosphatic masses up to 10 m long and 40 cm thick, are distributed throughout the formation. However, the large lenses, which have casings of Fe-Mn rock, are largely concentrated in

M.R. JOHNSON et al. Black (white-weathering) highly carbonaceous shale

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THE FORELAND KAROO BASIN, SOUTH AFRICA the dark-brown shale (Strydom, 1950). Carbonate concretions and irregular calcareous bodies, commonly with a brown encrustation, are dispersed throughout the formation in the west. The remains of a fossil shark have been found near the base of the formation at Klaarstroom (Oelofsen, 1986) while at Laingsburg sponge spicules, foraminifera, radiolaria and acritarchs have been reported. Strydom (1950) also noted the presence of radiolaria remains in phosphatic nodules. The mudrocks reflect suspension settling of mud from under- and overflows in a marine environment. The sand was deposited by tractional fall-out from turbidity currents entering the basin, probably from the southeast. The phosphatic and siliceous rocks probably formed by chemical or biological deposition under reducing conditions in areas where upwelling of cold water occurred (Parrish and Curtis, 1982). This upwelling probably enhanced a rich marine life which supplied the organic materials trapped in the bottom muds. Early diagenetic processes enriched the deposits, forming nodules and lenses. WhitehUi Formation. Since it defines the top of the Prince Albert Formation, the distribution of the Whitehill Formation coincides with that of the Prince Albert. The mudrocks of this formation weather white on surface, making it a very useful marker on aerial photographs as well as in the field. In the subsurface, the predominant facies is black, carbonaceous, pyrite-bearing shale. The shale is very thinly laminated and contains up to 14% carbonaceous material (Du Toit, 1954). The lower contact of the formation is always sharp and well defined. The upper contact is often transitional along the northern basin margin (e.g. Loeriesfontein and Douglas Fig. 5) where silty beds are included within the formation because of the interbedded black carbonaceous shales and biostratigraphic considerations. The thickness varies from 10 to 80 m (Cole and Basson, 1991). The Whitehill Formation loses its distinctive lithological character towards the northeast with its lower part containing siltstone and very finegrained sandstone in the Boshof area (Fig. 5; Cole, 1994; Cole and McLachlan, 1991) These rock types become thicker and more abundant northward and there is a merging of the unit with the underlying Prince Albert Formation (Cole, 1994). Cole and McLachlan (1991) correlated the Whitehill Formation with one of the coal seams (Top Seam) and overlying thin glauconite-rich sandstone in the uppermost Vryheid Formation in the northern Orange Free State. Van Eeden (1972) showed a similar correlation. Stavrakis (1986), on the other hand, correlated a black shale at the base of the Pietermar-

277 itzburg Formation in the Orange Free State Coalfield with the Whitehill Formation. However, sequence stratigraphy using genetic sequences of strata (Cole, 1994, figs. $2 and $3) indicates a correlation between the Prince Albert Formation, including the merged Whitehill Formation in its upper part, and the Vryheid Formation in the northern Orange Free State. An interbedded, subordinate, grey silty to cherty shale facies, which shows vague ripple lamination, is present along the western and southern basin margins. Thin yellowish- or reddish-weathering tuffaceous beds are also present along the southern and the western outcrop belt (McLachlan and Anderson, 1977a). Between Douglas and Boshof, thin pale grey tuffaceous beds occur sporadically in the formation (McLachlan and Jonker, 1990). Ferruginous carbonate concretions with a dolomitic composition are dispersed throughout the formation. At Loeriesfontein and Laingsburg, however, they are mostly confined to a specific stratigraphic level. Halite imprints were found on bedding planes between Kenhardt and Douglas (Van der Westhuizen et al., 1981). The biostratigraphy of the Whitehill Formation suggests synchronous deposition in the Karoo Basin, as the fossil ranges are confined everywhere to the upper part of the sequence (Fig. 7). The two species of the swimming reptile Mesosaurus show specific lateral distributions as Mesosaurus stereosternum is confined to the shallow-water margins of the basin and Mesosaurus tenuidens to the central deeper water areas (Oelofsen and Araujo, 1987). Remains of plants, palaeoniscoid fish and arthropods (Notocaris tapscotti) are common while rare fossil insect wings have been reported (McLachlan and Anderson, 1977b). The black, laminated, carbonaceous shales were deposited largely from suspension settling in a young, underfilled foreland basin under anoxic bottom conditions. A sea-level highstand, basin tectonics and climate were the controlling factors, interplay of which resulted in bounding conditions for organic-rich mud deposition. The organic material consists mostly of amorphous kerogen, except near the northern basin margin (Douglas-Boshoff) where land-derived woody and herbaceous debris predominates, particularly in the lower sandier half of the formation (Cole and McLachlan, 1991). These authors attributed the anoxic conditions in the deeper part of the basin to the presence of benthic cyanobacterial or algal mats which covered the basin floor. No evidence for such mats has yet been found, which led Visser (1992) to suggest that the coalforming environments along the steep palaeo-eastern basin margin were the source of mud and organic matter. These were transported in fresh-water plumes in an offshore direction during episodic

278

M.R. JOHNSON et al.

flooding and erosion of the organic-rich deposits. The high concentration of organic matter in the water body and the restricted oceanic circulation in the morphologically complex basin created anoxia in the water column. Preservation of organic matter in the general absence of benthonic fauna was thus high. Less anoxic conditions prevailed in the shallow marginal regions where deposition of very finegrained sandstone, siltstone and carbonate rocks interbedded with the black shales took place. The sand and silt probably represent distal turbidites and storm deposits, i.e. the microhummocky lenses of Dott and Burgeois (1982). Air-borne volcanic ash deposited together with the muds as well as in discrete layers was derived from a tectonic arc in the palaeo-west. Continuous inflow of fresh-water plumes in the restricted basin progressively caused brackish conditions suitable for the proliferation of aquatic fauna.

Southern formations (above the Whitehill Formation) The stratigraphy of the Ecca Group along the southern basin margin between 21 ~ E and 25 ~ E was first described in detail by Rossouw (1953). Johnson (1966) undertook a stratigraphic revision of the group east of 24 ~ E and subsequently presented data on the sedimentology, palaeoenvironments and sandstone petrography (Johnson, 1976). Kingsley (1977, 1981) provided a comprehensive description of the Ecca east of 26 ~ E, based on a series of detailed measured sections. Viljoen (1992a, b) has described the Collingham and Laingsburg Formations while Viljoen and Wickens (1992) have provided basic information on the Vischkuil Formation. The descriptions which follow are based largely on the above sources. Collingham Formation. Outcrops of the Collingham Formation are confined to the southern and western margins of the Karoo Basin (Fig. 3). The formation is generally between 30 and 70 m thick and comprises a rhythmic alternation of thin ultra-tabular beds (average 5 cm) of hard, dark grey, siliceous mudrock and very thin beds (average 2 cm) of softer yellowish tuff (Fig. 8). In the western part of the area minor sandstone and siltstone units occur in the upper half of the formation, while a prominent chert bed 0.2-0.6 m thick is present in the lower half. A variety of trace fossils, including grazing trails which probably belong to the Nereites community, are present at various levels (Kingsley, 1977; Viljoen, 1992a). The paucity of traction current indicators, the trace-fossil assemblage and the great lateral extent and fine-grained character of individual beds point to deposition from suspension

Fig. 8. Collingham Formation southwest of Beaufort West. The thin light-coloured layers are tuffs. Photo supplied by J.H.A. Viljoen.

in relatively deep water. Low-density distal turbidity currents were probably responsible for deposition of the thin sandstone and siltstone layers (Kingsley, 1977; Viljoen, 1992a). The tuff layers are considered to be silicic air-fall tufts which underwent secondary alteration subsequent to some reworking (see Viljoen, 1992a).

Vischkuil Formation. The predominantly argillaceous Vischkuil Formation overlies the Coilingham Formation in the southwestern part of the basin (Fig. 3). The Vischkuil becomes more arenaceous towards the east and grades into the Ripon Formation. A western cut-off is located where the overlying Laingsburg Formation pinches out. The formation varies in thickness between 200 and 400 m. The Vischkuil Formation consists essentially of dark shales, alternating with subordinate sandstones, siltstones and minor yellowish tuff layers. The shale units are thinly laminated or structureless, and display sharp upper contacts and gradational to sharp lower contacts. The thickest beds of this facies (maximum 15 m) occur in the lower half of the formation where they contain phosphatic and calcareous lenses and ferruginous layers with liesegang structures.

THE FORELAND KAROO BASIN, SOUTH AFRICA Sandstone (greywacke) beds vary from 0.3 to 1.5 m in thickness, are fine-grained, parallel-sided and characterised by sharp lower contacts, gradational to sharp upper contacts, lateral persistence and a large variety of primary sedimentary structures. The thicker sandstones are mostly massive and/or planar laminated with erosional sole marks, load casts, horizontal streaks of rip-up clay clasts and dewatering structures. The thinner, very fine-grained sandstone beds in the lower part of the formation normally comprise a basal cross-laminated and/or planar laminated zone (5-18 cm), followed by a structureless zone containing pseudonodules and an upper structureless zone which grades into the overlying shales. The cross-laminated sets are 3-5 cm thick, display very low foreset dip angles (5-10 ~ and are often capped by wavy laminae (Wickens, 1985). Stow and Shanmugam (1980) describe similar sequences of structures in fine-grained turbidites. Yellowish-green to khaki-coloured tuff layers occur sporadically throughout the formation. They vary from 1 to 20 cm in thickness, are laterally persistent, lack any traction structures and sometimes show normal grading with slight mottling near the base. The presence of palygorskite, tridymite and cristobalite in these beds, as well as their similarity to the volcanic material in the Collingham Formation, confirms their volcanic origin. Kuenen (1963) described the lower part of the Ecca Group in the southwestern part of the basin as a flysch-like succession. His conclusions were later confirmed by Theron (1967) and Truswell and Ryan (1969). The former interpreted the Vischkuil as a distal turbidite succession. The facies assemblage, abundance of sole marks, nature of bedding contacts and the type and vertical arrangement of sedimentary structures reflect sediment gravity flow deposition in a basin plain to outer basin floor fan environment. Normal suspension deposition and deposition from thick muddy turbidites (lighter coloured silt and shale beds) were regularly interrupted by deposition of fine-grained turbidites displaying traction structures. The occurrence of volcanic ash layers and large-scale slumped beds reflect volcanic and seismic activity during deposition of the Vischkuil (Wickens, 1985). Transport of the Vischkuil and overlying Laingsburg Formations was from the south, with a subsidiary west-to-east axial flow trend (Truswell and Ryan, 1969).

Laingsburg Formation. The contacts of the Laingsburg Formation with the overlying Fort Brown and underlying Vischkuil Formations are generally gradational. The formation comprises four arenaceous units separated by shale units and is approximately 400 m thick in its type area. It wedges out towards the west and north. A vertical cut-off is

279 located in the east where the Vischkuil Formation merges with the Ripon Formation. The arenaceous units of the Laingsburg Formation comprise sandstone, siltstone and dark shales, with the former two rock types predominating and building prominent east-west trending ridges. These lithofacies are arranged in upward-thickening and upward-thinning sequences (in the lower and upper parts of the formation respectively) with respect to bed thickness (Wickens, 1985). The thick massive sandstones (up to 30 m thick, with individual beds up to 4 m thick) usually occur at the bases of upward-thinning cycles and are fine- to medium-grained and remarkably parallel-sided with abrupt upper and lower contacts. Their undersurfaces vary from fiat to highly irregular with a wide variety of sole structures such as flame structures and load, flute and groove casts. They grade upward sharply into planar-laminated siltstone and shale which commonly contain coalified plant fragments. Amalgamation of beds, clay pebble clasts, dewatering structures and calcareous concretions are also common features in the thicker sandstones. The majority of thinner sandstone beds, where they alternate with siltstone and shale, are planar-laminated near their bases or sometimes only in the upper few centimetres and also display a variety of sole structures. Graded bedding is present in places. Ripple cross-lamination normally succeeds planar lamination and also characterises the thinnest sandstone and siltstone beds where they occur rhythmically interbedded. The dark planar-laminated shales often contain lenticular calcareous concretions. Clastic dykes are present and are thickest in the upper shale successions. Sinuous paired "fish trails" (Undichna bina) and arthropod trackways (Umfolozia) (Anderson, 1975) are commonly found on bedding surfaces. Land-derived plant material is abundant, especially in the upper parts of the formation. The absence of traction structures and visible grading in the thick-bedded sandstones probably indicate deposition from high-density turbidity currents with the involvement of processes such as fluidised/liquified flows (Wickens, 1985). The structures observed by Kuenen (1963) and Theron (1967) and regarded by the latter as belonging to proximal turbidites were interpreted as indicative of submarine fan deposition by Visser et al. (1980). This was taken a step further by Wickens (1985) who interpreted the upward-thickening (10-20 m thick) cycles as progradation of submarine fan lobes and the upward-thinning cycles as abandonment of depositional lobes and channel-fill sequences in a midto outer-fan depositional environment. The turbidity flows were generated on a delta-front slope and were the precursors of deltas prograding from the south (Visser et al., 1980; Wickens, 1985).

280

Ripon Formation. The Ripon Formation is generally 600-700 m thick, but is over 1000 m in the eastern third of the outcrop area (along the southern basin margin). It consists of medium grey (frequently mottled), poorly sorted, fine- to very finegrained lithofeldspathic sandstones alternating with dark grey, fine-grained clastic rhythmite and mudrock units (Fig. 9). The Ripon in the eastern area can be subdivided into a lower Pluto's Vale Member consisting predominantly of sandstone, a middle Wonderfontein Member (mudrock/rhythmite) and an upper Trumpeters Member comprising alternating sandstone and mudrock (Fig. 9). Sandstones of the upper member pinch out towards the west, where the Ripon is correspondingly thinner (Fig. 4). Drilling has shown that the sandstones wedge out rapidly to the north of the outcrop area. In the lowermost 50 m of the succession sandstone units average about 0.3 m in thickness. For the rest of the formation the average is about 12 m, with a maximum thickness of 44 m. Most of the sandstones are composite, consisting of a number of individual sedimentation units. The sandstones are normally parallel-sided and tabular in individual outcrops. Lower boundaries are sharp, while the upper boundaries are normally gradational into mudrock or rhythmite over an interval ranging from a few millimetres to a few metres. Graded bedding is prominent in the lowermost 50 m of the formation, but poorly developed elsewhere, with most individual sedimentation units being superficially massive. Other internal current structures such as horizontal lamination, ripple lamination and wavy bedding are very subordinate. The very thin sandstone/siltstone beds present in the rhythmite intervals are usually graded, ripple-laminated or horizontally laminated. Groove casts and associated bounce, skip, and prod marks are common on the lower surfaces of sandstones in the lowermost 50 m, and also occur sporadically throughout the rest of the formation. Flute casts are rare. Poorly developed low-amplitude tipples are occasionally present. Deformation structures present include load casts, flame structures, convolute bedding, sandstone dykes and sills and slump structures. Roughly spherical calcareous concretions 10-15 cm in diameter, usually with shale-fragment nuclei, are fairly common in the sandstones. Larger brown-weathering calcareous sandstone bodies 30-150 cm in diameter are also present in places. Trace fossils of various kinds, including tracks, trails, tubes and burrows, occur sporadically throughout the Ripon Formation (Anderson, 1974; Kingsley, 1977). These generally reflect deep-water conditions. Unidentifiable carbonised plant remains and oval petrified logs displaying well-developed annual rings are present near the base of the formation.

M.R. JOHNSON et al.

Fig. 9. Type section of the Ripon Formation northeast of Port Elizabeth. Simplified after Johnson and Kingsley (1993).

THE FORELAND KAROO BASIN, SOUTH AFRICA Palaeocurrent data presented by Kingsley (1977) for the eastern area point to a general northnorthwesterly transport direction, with high vector strengths (average 0.85) being obtained at individual localities. Since the basin configuration suggests transport down the flank rather than along the axis of the depositional trough, the source area must have been located to the south and southeast of the southern basin margin. There is little doubt that the sandstones of the Ripon Formation were produced by sediment gravity flows. Apart from the distinctive suite of structures described above there is a complete absence of structures indicative of either shallow water or normal traction currents. The sandstones of the basal 50 m as well as those in the middle shale member presumably represent distal turbidites and those in the remainder of the formation proximal turbidites. Walker (1979, fig. 13) classified thick graded beds as proximal and massive sandstones as braided suprafan deposits. Lowe (1982) ascribed thick massive fine-grained sandstones to suspension sedimentation from high-density turbidity currents or liquefied flows. Deposition probably took place on turbidite fan complexes at the foot of an advancing delta slope. Interbedded mudrock units represent deposition of fines from suspension on the basin floor. Modal analyses of 17 Ripon Formation sandstones (Johnson, 1976) gave the following average mineral composition: quartz: 11%; feldspar: 21%; lithic fragments: 34%; matrix: 29%; cement (calcite): 3.5%; accessories: 1.5%. On a triangular QmFLt diagram these samples plot in the magmatic arc provenance field of Dickinson et al. (1983) (Johnson, 1991; Fig. 10). The abundance of volcanic, mostly felsitic, rock fragments (16% of the total sandstone composition) confirms that volcanic rocks were a prominent constituent of the source area. Since 90-95% of the feldspar in the Ripon Formation consists of plagioclase (Martini, 1974; Kingsley, 1977) it is probable that the plutonic component of the source area was granodioritic, tonalitic and dioritic in composition. The present albitic composition of the plagioclase appears to reflect in situ albitisation of an originally more calcic plagioclase (Martini, 1974). The micaceous and schistose fragments which on average constitute 12% of the total sandstone composition were probably derived from fine-grained low-grade metamorphic rocks, as no doubt was much of the matrix material. Fort Brown Formation. The Fort Brown Formation consists of rhythmite (Fig. 11) and mudrock with minor sandstone intercalations in places and displays an overall coarsening-upward tendency. One or more fairly prominent sandstones may, how-

281 Qm

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ever, be present some distance below the upper contact. Average formation thickness is about 1000 m, with values ranging from about 500 m to 1500 m. Outcrops are confined to the southern margin of the basin. Individual sand/silt and silt/clay layers comprising the rhythmite units are equal or subequal in thickness, ranging from a few mm to a few cm, and are laterally very persistent. Individual sand/silt layers display a general upward increase in thickness within the formation. Current structures are absent or poorly developed. Towards the top of the formation straight-crested oscillation ripples and associated wavy bedding are present on the upper surfaces of many of the more prominent sand/silt layers. In the central part of the outcrop area a vector mean of 105~ ~ (vector strength 0.70) was obtained on 80 measurements of ripple crest orientations (data from Ryan, 1967). In general both upper and lower contacts of the thinner sand/silt layers are gradational while the thicker layers often display sharp boundaries. The interbedded sandstones that occur sporadically at various levels within the formation are either massive or display various internal structures such as horizontal lamination, wavy lamination, tipple lamination and ripple-drift cross-lamination (Kingsley, 1977). Ball-and-pillow and related deformation structures are present in places towards the top of the formation. Calcareous concretions 30-100 cm in diameter and 10-30 cm thick occur sporadically. Trace fossils are represented by worm burrows and pairs of sinuous trails spaced 1-2.5 cm apart

282

M.R. JOHNSON et al.

Fig. I I. Rhythmite in the Fort Brown Formation northeast of Port Elizabeth.

( U n d i c h n a b i n a ~ Anderson, 1974). The latter are thought to represent markings produced by the ventral spines of a fish (Haughton, 1928). The presence of C r u z i a n a and S k o l i t h o s community traces towards the top of the formation (Kingsley, 1977) confirms the progressive decrease in water depth suggested by the increase in grain size and appearance of wave ripples. The Fort Brown rhythmites can be compared to the marine shale with silt layers (shelf and pro-delta deposits) and the alternating sand, silt and clay layers (distal bar deposits) described by Wright (1978) from modern deltas. The banded or striped silty mudrocks illustrated by Elliot (1978, figs. 6.40, 6.42) and assigned by him to the basal part of a prograding delta front also resemble these rhythmites. According to Elliot the diffuse banding defined by slight variations in grain size in these mudrocks reflects fluctuations in the supply of suspended sediment. Higher up in the succession the coarser siltstone and sandstone beds which are repeatedly intercalated in mudstone-siltstone background sediment represent distal flood-generated sediment incursions from distributary mouths. The presence of a crystal tuff comprising 90% plagioclase and 10% quartz in the Fort Brown Formation north of Port Elizabeth (Lock and Johnson, 1974) provides evidence for contemporaneous dacitic-andesitic volcanism.

Waterford Formation. The arenaceous Waterford Formation overlies the Fort Brown Formation along most of the southern outcrop belt (Fig. 2). Its thickness fluctuates between 200 and 800 m, except where it tapers out at either end of its

distribution range. The formation comprises alternating medium grey (commonly speckled), very finegrained, lithofeldspathic sandstones (~50% of total) and medium dark grey mudrock or clastic rhythmite units (Fig. 12). The Britskraal Shale Member, which occurs in the upper part of the Waterford Formation in the eastern outcrop area, averages 100 m in thickness and consists essentially of dark grey mudrock and clastic rhythmite. Individual sandstone units are moderately tabular with average and maximum thicknesses of 6 m and 18 m respectively. Most beds are superficially structureless; in order of abundance observed internal structures include horizontal lamination, low-angle cross-bedding (hummocky cross-stratification?)and (rarely) cross-bedding and ripple lamination. Beds (sedimentation units) range from a few cm to over a metre in thickness, and tend to be laterally persistent. Well-developed straight-crested oscillation ripples (wave length 5-8 cm, amplitude 7-10 mm, average ripple index about 7) are common in the more arenaceous rhythmites and the thin-bedded sandstones. Ball-and-pillow and related deformation structures are relatively common throughout the Waterford Formation, but are especially abundant in the upper part of the unit. Isolated, contorted sandstone blocks, balls, lenses and stringers often occur within massive, very fine-grained sandstone or siltstone units; slumping as well as foundering appear to have been involved in their formation. Individual "balls" or "pillows" vary from about 10 cm to over a metre in diameter. Thin mud-flake conglomerate layers

THE FORELAND KAROO BASIN, SOUTH AFRICA

283 are occasionally present. Brown-weathering, generally oval calcareous concretionary bodies 30-150 cm in diameter occur sporadically in the sandstones, while flattened bun-shaped concretions 30-100 cm in diameter and 10-30 cm thick are present in the argillaceous rocks. Trace fossils (trails, tubes, burrows) are fairly common throughout the formation. The abundance of wave tipples in the Waterford Formation indicates that it accumulated in relatively shallow water. The overall lithology and internal sandstone structures coupled with the fact that the formation overlies a thick succession of typical prodelta muds suggests a delta front environment. The virtual absence of true cross-bedding indicates that relatively sluggish coastal currents reworked material deposited by distributaries, thereby creating a series of delta front sheet sands. A relatively steep depositional slope prevailed, as evidenced by the abundance of ball-and-pillow and other soft sediment deformation structures. The marked thickness fluctuations point to the presence of a number of distinct delta lobes. The following average mineral composition (11 samples) was obtained (Johnson, 1976): quartz: 19%; feldspar: 23%; lithic fragments: 34%; matrix: 20%; accessories: 2%; cement: 2%. As in the case of the Ripon Formation the mean quartz:feldspar:rock fragments ratio falls within the magmatic arc provenance field of Dickinson et al. (1983) (Johnson, 1991; Fig. 10). Volcanic (felsitic) rock fragments on average constitute about 12% of the total composition, while micaceous/schistose fragments form 16%. The source was probably essentially the same as that proposed for the Ripon Formation.

Fig. 12. Lower two-thirds of the Waterford Formation type section northwest of Port Elizabeth.

Western and northwestern formations (above the Whitehill Formation) Tierberg Formation. The Tierberg Formation is a predominantly argillaceous succession considered to be a basin floor to distal delta front deposit (Wickens, 1984). It reaches a thickness of approximately 700 m along the western margin of the basin, thinning to about 350 m towards the northeast. It rests with a sharp contact on the Collingham or Whitehill Formations and grades upward through a 30-70 m thick transition zone into the arenaceous Waterford Formation or, in the east, into the Adelaide Subgroup (Beaufort Group). The upper boundary has generally been placed at the base of the first prominent sandstone in the succession or where a marked change in lithology from predominantly argillaceous to relatively arenaceous occurs. Where it is overlain by turbidites of the Skoorsteenberg Formation the contact is fairly sharp and the formation attains a thickness of about 460 m (Wickens, 1984). The bulk of the Tierberg Formation comprises well-laminated bluish-grey to almost black shale.

284 The shale is carbonaceous and contains abundant pyrite. According to Potgieter (1974) it is characterised by a remarkably high proportion of feldspar. Some yellowish tuffaceous beds up to 10 cm thick occur in the lower part of the succession along the western and northern margins of the basin (Visser et al., 1980; Jordaan, 1981; Wickens, 1984; Viljoen, 1992a) whereas calcareous concretions are common towards the top. Potgieter (1974) found fish scales and sponge spicules in some of the concretions. Chert beds up to one metre thick were reported by Visser et al. (1980) in the southwestern part of the area near the base of the formation. Trails of the Nereites ichnofacies are fairly common on bedding planes. Rhythmically interbedded grey sandstone or siltstone and dark shale occur at various levels in the sequence (Terblanche, 1979; Visser et al., 1980). This facies contains carbonate concretions and trace fossils of the Planolites ichnofacies in places. The planar lamination of the dark grey shale facies suggests settling from suspension in a lowenergy environment. Water depth in the Karoo Basin was probably at its maximum during deposition of this facies as indicated by the Nereites ichnofacies. Visser and Loock (1978), however, estimated that water depth did not exceed 500 m in the south. Towards the north, the presence of silt layers within this facies indicates a closer proximity to a provenance. The transition zone at the top of the formation consists of a number of upward-coarsening sequences 2-10 m thick comprising mudstone, siltstone and very fine-grained sandstone (Lemmer, 1977; Wickens, 1984; Zawada, 1988; Fig. 13). This zone is about 70 m thick along the western margin of the basin (Wickens, 1984) thinning to 30 m towards the northeast (Zawada, 1988). The sandstone beds increase in thickness upward in the transition zone but are normally less than 2 m thick. The thinner sandstone and siltstone beds are planar laminated and ripple cross-laminated while the thicker sandstones are massive, wave ripple cross-laminated, planar laminated, climbing ripple laminated and less frequently low-angle cross-bedded (Lemmer, 1977; Terblanche, 1979; Visser et al., 1980; Wickens, 1984; Zawada 1988). Thin clay-pellet conglomerates are present along the western margin of the basin (Wickens, 1984) and in places along the west-east outcrop (Terblanche, 1979; Siebrits, 1987; Zawada, 1988) while calcareous concretions occur fairly frequently. Slump structures, ball-and-pillow structures and other soft sediment deformation features are fairly common throughout the upper part of this succession. Structureless mudstones with curled and deformed fragments of sandstone are common along the western margin of the basin (Wickens, 1984).

M.R. JOHNSON et al. Wickens considered these to be the result of gravitational mass flow of water-saturated sediments in a coherent state. Nel (1977) reported an isolated occurrence of turbidites 6-9 m thick, consisting of 19 incomplete Bouma sequences, at the top of the transition zone east of 24~ The transition zone is bioturbated at intervals with Zoophycos ichnofacies at the base and Cruziana and occasionally Skolithos towards the top. A proximal pro-delta to distal delta front depositional environment is widely accepted for the transition beds. The abundance of wave ripples, distribution and types of trace fossils, and grain size trends indicate regression in a fairly shallow water environment. Soft-sediment deformation features are more common in delta front deposits where the sedimentation rates are high and where oversteepening leads to unstable conditions conducive to slumping and downslope gravitational flow of sediments. The absence of features indicative of thick beach and barrier deposits favours a fluvialdominated deltaic model. However, Siebrits (1987) interpreted the upper part of the Tierberg Formation and overlying Waterford Formation in the area between 22 ~ and 22~ as a storm shelf sequence, disagreeing with the deltaic model of Terblanche (1979). Siebrits based his model mainly on the abundance of low-angle cross-bedding (also reported by other authors in other areas) which he interpreted as hummocky cross-stratification. Palaeocurrent directions are consistently northeast to east-northeast for the entire area.

Skoorsteenberg Formation. The Skoorsteenberg Formation outcrops as an almost fiat-lying, lens-shaped, arenaceous unit in the southwestern part of the basin (Wickens, 1984). Although the Skoorsteenberg and Laingsburg Formations occupy similar stratigraphic positions, they do not link up with each other and are considered to be two separate basin floor fan complexes. The Skoorsteenberg Formation attains a maximum thickness of approximately 200 m at its type locality where it comprises five sandstone-rich units up to 65 m thick with shale units separating them. A sixth fan unit crops out in the extreme south. These sandstone-rich units extend over various distances laterally, and in places the formation consists of only one such unit (Fig. 14). The base and top of the formation have sharp contacts with the Tierberg and Kookfontein Formations respectively. The individual sandstones vary in thickness from a few centimetres to almost 6 m, are parallel-sided with sharp upper and lower boundaries and are fairly persistent along strike. Sedimentary features include massive bedding, rip-up clasts, dewatering structures, Bouma sequences (Bouma, 1962) and sole marks ranging

THE FORELAND KAROO BASIN, SOUTH AFRICA

285

Fig. 13. Representative sections of the upper part ("transition zone") of the Tierberg Formation. Section A after Wickens (1984), section B after Lemmer (1972) and section C after Zawada (1987). For location of sections see Fig. 14. from biogenic markings (tracks, trails and burrows) to large groove and fluid scour marks (Fig. 14). Coalified plant fragments mixed with silt and clay clasts commonly occur at the top of the sandstone beds. The depositional setting and larger-scale sedimentary features indicate a basin floor fan origin for the Skoorsteenberg sediments. The sedimentary structures observed in the Skoorsteenberg sandstones range from those constituting the classical Bouma sequences (Bouma, 1962) to those characteristic of grain and fluidized/liquified flow deposits as defined in the genetic classification of Middleton and Hampton (1976). Deposition is believed to have resuited from highly concentrated density flows where sediments were supported by a range of different mechanisms, including turbulence, as well as from relatively low concentration flows where sediment was mainly supported by turbulence. Traction structures and the general thinning and pinch-out directions of the formation indicate a northeasterly to easterly flow for the density currents. The cyclic nature of the Skoorsteenberg Formation sediments suggests successive, tectonically controlled progradation of several basin floor fans associated with a fluvially dominated delta system. The density flows most likely originated on the unstable delta front slope or at the river mouths during

extreme flood events and accumulated as coalescing fans and aprons in the basin floor to pro-delta environment. Palaeotransport was from the southwest, west and northwest with a NNW-SSE palaeoshoreline. Trace fossils are mainly horizontal feeding trails such as Lophoctenium and Plagiogmus, suggesting a water depth of approximately 500 m (Wickens, 1984). Plant fragments, such as Glossopteris, are also abundant.

Kookfontein Formation. The Kookfontein Formation overlies the Skoorsteenberg Formation with a sharp contact and grades upwards into the Waterford Formation, becoming more sandy towards the top. It correlates laterally with the upper part of the Tierberg Formation and is approximately 350 m thick at its type locality (Wickens, 1984). The Kookfontein mainly consists of siltstone, shale and fine-grained sandstones. The lower part is characterised by horizontally laminated dark-grey shales, alternating with rhythmically interbedded shales and siltstones which form minor upwardthickening cycles. The cycles become more prominent towards the top of the formation where they consist of alternating siltstone and thin sandstone beds, and are often capped by a thick sandstone bed. Slump and load features are very common

286

M.R. JOHNSON et al.

Fig. 14. Columnar sections illustrating some sedimentary characteristics of the Skoorsteenberg (1) and Waterford (2) Formations in the western part of the basin (after Wickens, 1984). Inset map shows location of sections depicted in Figs. 13 and 14. and occur interbedded with undisturbed beds. Planar lamination and ripple cross-lamination are the most common primary structures in the sandstones. The Kookfontein Formation represents a continuation of pro-delta sedimentation (after cessation of the Skoorsteenberg gravity flow events), changing upward into delta front deposition. The repetition of upward-coarsening cycles (2-15 m thick) and their association with interbedded slump and slide features indicate rapid progradation and switching of delta lobes with subsequent development of unstable

conditions. The abundance of wave ripple marks and bioturbation (Cruziana facies of Seilacher, 1967) reflects the overall shallowing of the depositional environment during deltaic progradation.

Waterford Formation. The description of the Waterford Formation along the western flank of the basin is based mainly on the work of Wickens (1984). In this area the formation has a mean thickness of 130 m and overlies the Kookfontein and Tierberg Formations with a gradational contact

THE FORELAND KAROO BASIN, SOUTH AFRICA (Wickens, 1984). The contact with the overlying Abrahamskraal Formation is relatively sharp and the break in lithology represents a change from the lower delta plain to a mud-rich, subaerially exposed upper delta plain environment. The major rock types are sandstone, siltstone and shale (Fig. 14). The sandstones are fine- to medium-grained, with the coarsest grain sizes associated with clay-pebble conglomerates and basal channel-fill deposits. The lower part of the formation is characterised by upward-coarsening cycles (capped by extensive sheet-like sandstones) which alternate with chaotic slump and slide deposits, while the upper part consists of alternating thick horizontally laminated and low-angle trough (hummocky?) cross-stratified sandstones up to 8 m thick, siltstones, ball-and-pillow layers and channel-fill deposits. The thinner sandstone and siltstone beds are mostly horizontally laminated with wave-rippled upper surfaces. Tabular beds of massive silty sandstone with deformed fragments of coarser sandstone also occur. Slumped and load-deformed beds are common and alternate with rhythmically interbedded sandstone and siltstone beds. Fossil wood is abundant in the sandstones and a variety of trace fossils, mainly vertical to subvertical burrows of the Cruziana-Skolithos ichnofacies, are present. The Waterford Formation in the western outcrop area represents the terminal topset deposits of fluvially dominated deltas that prograded eastward (Visser et al., 1980; Jordaan, 1981; Wickens, 1984). The delta topset succession can be subdivided into delta front deposits (mainly upward-coarsening distributary mouth bars) and lower delta plain deposits (interdistributary bay-fill sequences, crevasse splays and channel fills) which are overlain by subaerially deposited upper delta plain muds. This lithological break between lower and upper delta plain represents the transition (diachronous boundary) between the Ecca and Beaufort Groups. The Waterford Formation, together with the underlying Tierberg and Kookfontein Formations, thus forms an upwardcoarsening deltaic megacycle (shelf, pro-delta and delta front) which culminates in argillaceous fluviatile upper delta plain deposits (Visser and Loock, 1978; Wickens, 1984). Palaeotransport was from the southwest, west and northwest (Ryan, 1969; Wickens, 1984). The easterly extension of the Waterford Formation along the northwestern margin of the Karoo basin has been described by Terblanche (1979), Siebrits (1987) and Rust et al. (1991). It attains a thickness of 250 m north of Beaufort West (Siebrits, 1987) but thins and eventually pinches out towards the east. The formation constitutes an overall regressive sequence, grading from the argillaceous Tierberg Formation upward through a sequence of

287

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iERBERG FORMATION Fig. 15. Representative composite sections of the Waterford Formation north of Beaufort West (after Siebrits, 1987). Note that the uppermost part of the formation has been removed by erosion. alternating sandstone, siltstone and mudstone into a sandstone-dominated interval at the top (Fig. 15). The sandstones are greyish in colour (yellowish when weathered), fine- to very fine-grained and attain a thickness of 8 m. Hummocky cross-stratifica-

288 tion and ripple cross-lamination (mostly with external symmetrical tipple forms) are the most common internal primary sedimentary structures in the sandstone and coarser-grained siltstones. The former occurs more abundantly towards the top of the Waterford Formation (Siebrits, 1987). Less common structures include interference ripples, climbing tipple lamination and flaser lamination (Terblanche, 1979; Siebrits, 1987). The mudrocks and fine-grained siltstones are grey to dark grey in unweathered samples and are horizontally laminated. Clay-pellet conglomerates, some with erosional bases, with a maximum thickness of 50 cm and maximum lateral extent of 3 m occur sporadically throughout the formation. Clay clasts are up to 5 cm in diameter. Sandstones overlying these conglomerates are hummocky cross-stratified. Brownweathering calcareous concretions with a maximum diameter of 2 m occur in all the rock types. Slump structures, ball-and-pillow structures and load casts are common throughout the formation. A large proportion of the formation has been bioturbated to a greater or lesser degree; in places the bedding has been completely destroyed. According to Siebrits (1987) Planolites (Cruziana ichnofacies) and Scoyenia ichnofacies are the most common trace fossils while Skolithos occurs sporadically. The different facies commonly form upwardcoarsening sequences ranging in thickness from 4 to 38 m. Terblanche (1979) also described upward-fining units 1-4 m thick in the topmost part of the formation. He considered the sequence to have resulted from fluvially dominated deltaic sedimentation. He recognised two first-order upward-coarsening regressive cycles, each grading from pro-delta shale upward into distributary mouth bar sands, the upper sequence being overlain by upward-fining distributary channel fill and point bar sediments. Smaller upward-coarsening cycles occurring within the first order cycles he considered to be the result of avulsion and progradation of small lobate deltas into shallow water. Siebrits (1987) concluded that the Waterford Formation in the northern outcrop area had been deposited in a wave-dominated shallow shelf environment (~50 m water depth) with frequent reworking of sediments during storms. He based his model mainly on the abundance of hummocky cross-stratification and wave ripples and the absence of any real evidence for subaerial deposition. He showed that the low-angle cross-bedding which Terblanche (1979) took to represent deposition on distributary mouth bars was in fact hummocky cross-stratification. Whereas Terblanche found upward-fining units only at the top of the formation, Siebrits (1987) recognised such units in one or two places lower

M.R. JOHNSON et al. down in the succession. The sandstones in these fining-upward sequences are hummocky cross-bedded which rules out the possibility of deposition in channels. Siebrits ascribed the mud-pebble conglomerates to erosion of compacted mud beds by bottom currents generated during storms. Siebrits considered the presence of Scoyenia burrows, which normally characterises fresh-water sedimentation, to be the result of fresh-water flood incursions into the basin. Skolithos developed during storms and the Cruziana ichnofacies during the intervening quiet periods. Siebrits deduced from the orientation of wave tipple crests and current lineation that the shoreline had a north-northwest orientation and that the sediments were derived from the west.

Northeastern formations Pietermaritzburg Formation. The Pietermaritzburg Formation comprises a monotonous succession of dark bluish-grey silty mudrock which crops out only along the eastern margin of the basin. Extensive drilling has shown that it underlies the Vryheid Formation over almost the entire northeastem part of the Karoo Basin. It attains its maximum thickness of over 400 m in the southeast (Du Toit, 1954), thinning towards the north and pinching out against the Dwyka Group at about 26~ and against a palaeoscarp near the northwestern margin of the Karoo Basin. Towards the south, where the Vryheid Formation pinches out, it merges with the Volksrust Formation, and thus passes laterally into undifferentiated Ecca Group strata. It can be broadly correlated with the Prince Albert Formation (see p. 275). The contact between the Pietermaritzburg Formation and the underlying Dwyka Group is mostly sharp. However, Mathew (1974) and Van Vuuren (1983) reported local occurrences of siltstone with scattered pebbles (glacial dropstones) in the lowermost part of the formation. The contact between the Pietermaritzburg and Vryheid Formations is strongly diachronous. Sandstones successively higher up in the succession shale out towards the southeast. Widespread carbonate concretions, lenses and beds record changes in Eh-pH levels consistent with relatively shallow water (Visser and Loock, 1978). Invertebrate trace fossils can be seen on bedding planes of carbonate-cemented mudrock. Microbioturbation may account for the general lack of structures in Ecca shelf mudrocks (Hobday, 1973). Sharp-based burrowed sandstones showing graded bedding near Vryheid bear witness to resedimentation events on an unstable shelf (Roberts, 1986). The formation coarsens upward with heavily bioturbated and penecontemporaneously deformed sandy and silty beds appearing near the top; these secondary

THE FORELAND KAROO BASIN, SOUTH AFRICA

289 particularly in the thin northwestern part where they constitute the entire Vryheid in places. A relatively thin fluvial interval which grades distally into deltaic deposits towards the southwest and south occurs approximately in the middle of the formation in the east and northeast (Figs. 16, 17). The base of an idealised upward-coarsening deltaic cycle in the eastern part of the formation (Fig. 18) consists of dark grey, muddy siltstone resuiting from shelf suspension deposition in anoxic water of moderate depth. Pro-delta sediments are represented by alternations of bioturbated, immature sandstones, dark siltstones and mudstones on a centimetre to decimetre scale. The rhythmic nature of this facies reflects seasonal variations in fluvial input or storm/fair weather deposition. Distal distributary mouth bar sediments comprise repetitive units of horizontally laminated, medium-grained, well-sorted sandstone grading upwards into ripple cross-laminated, fine-grained sandstone and siltstone, bioturbated in places. Hummocky cross-bedding may also be present (Roberts, 1986). Slump structures and high-density turbidites are fairly common and testify to slope instability at the delta front (Van Vuuren, 1983). Proximal mouth bar deposits take the form of trough cross-stratified medium- to coarse-grained sandstone with polymodal palaeocurrent patterns; the variable trough orientation records the influence of both distributary channel and basinal currents. This facies is erosively overlain by coarse to pebbly, planar cross-stratified, feldspathic sandstone with unimodal palaeocurrent patterns interpreted as distributary channel fills. The delta front deposits have a sheet-like geometry due to sediment redistribution by basinal currents

structures record colonisation of a more competent substrate by filter-feeding organisms and foundering into undercompacted muddy silts. The coarser sediments reflect shoreline progradation as the climate warmed and rivers, fed by glacial meltwater, entered the basin in the northwest and northeast. This heraided a new phase in the depositional history of the northern Karoo Basin. The salinity of the primeval Ecca Sea has long been a contentious issue but in recent years evidence has been mounting in favour of marine conditions. Glauconite and trace fossils indicative of saline water have been recorded in subaqueous Vryheid Formation sediments which overlie the Pietermaritzburg Formation strata (Cairncross, 1979; Le Blanc Smith, 1980; Mason et al., 1983). Vryheid Formation. The Vryheid Formation comprises a siliclastic wedge (Fig. 16) thinning towards the north, west and south from a maximum of approximately 500 m along section E (Fig. 2) The uneven pre-Karoo topography in the vicinity of the northern and northwestern margins of the basin where the Vryheid rests directly on pre-Karoo rocks or Dwyka Group tillite gives rise to marked variations in thickness. In these areas the Vryheid pinches out against numerous local basement highs. Thinning and pinch-out towards the southwest and south is due to a facies gradation of its lower and upper parts into shales of the Pietermaritzburg and Volksrust Formations respectively (Figs. 16, 17). The different lithofacies of the Vryheid Formation are mainly arranged in upward-coarsening cycles which are to a great extent deltaic in origin. Linear coastline cycles are, however, fairly common

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290

M.R. JOHNSON et al.

Fig. 17. Schematic west-east section through the Ecca Group in the northeastern part of the Karoo Basin (after Van Vuuren, 1983). (Le Blanc Smith, 1980) and were probably lobate in form. Delta plain sediments are frequently heavily bioturbated as a result of inundation following abandonment. Many variations in the pattern of deltaic deposition have been noted due to factors such as delta switching, degree of wave influence, fluvial erosion and water salinity. Fining-upward fluvial cycles are typically sheetlike in geometry, although some form valley fill deposits. They comprise coarse-grained to pebbly, immature sandstones which are erosively based and planar and trough cross-stratified, with an abrupt upward transition into fine-grained sediments and coal (Fig. 18). Reactivation surfaces and internal scours are common, and the sequences are usually vertically repetitive. Such sequences have been attributed to either meandering rivers (Hobday, 1973; Van Vuuren, 1981) or braided streams (Le Blanc Smith, 1980, and many others). Probably both depositional modes were active, depending on local conditions such as bank stabilisation by vegetation or proximity to source areas, etc. Peat swamps developed on broad abandoned alluvial plains and, less commonly, in interfluves.

Nondeltaic progradation cycles have been reported in various localities in the Vryheid Formation (Vos and Hobday, 1977; Van Vuuren, 1981, 1983; Tavener-Smith, 1982; Roberts, 1986). The lower portion of most units resembles the distal deltaic facies described above except that turbidites and slump structures are usually absent. Instead of the distributary mouth bar/distributary channel association, however, fine- to medium-grained, relatively mature sandstones are developed. Internal structures include bioturbation, wave ripple lamination, fiattopped and interference ripples and low-angle swash lamination, together with plant rootlets in some instances. These facies associations point to a foreshore setting. A heavy mineral bearing succession characterised by poorly sorted, medium- to largescale, erosively based sandstone in its lower part was interpreted by Behr (1965) and Vos and Hobday (1977) as a storm-beach succession. Van Vuuren and Cole (1979), Van Vuuren (1981, 1983), Le Blanc Smith (1980), Winter (1985) and Caimcross and Cadle (1987) found that the different rock types can be grouped into a limited number of

THE FORELAND KAROO BASIN, SOUTH AFRICA

291

Fig. 18. Idealised upward-coarsening deltaic and upward- fining fluvial cycles in the eastern part of the Vryheid Formation (after Roberts, 1986). genetic sequences as defined by Busch (1971), which can be correlated regionally. Van Vuuren and Cole (1979) and Van Vuuren (1983) furthermore correlated these sequences over the entire Vryheid Formation although the correlation is doubtful in places. Each of the sequences comprises one or more progradational phases, usually deltaic in origin, and/or an aggradational fluvial phase with a transgressive marine sandstone unit terminating the sequences in places.

Eastern facies. For convenience of description, a simple three-fold subdivision of the Vryheid Formation into a lower deltaic interval, a middle fluvial interval and an upper deltaic interval (Figs. 16, 17) is used in the east. These correspond approximately to the "lower sandstones", "coal zone" and "upper sandstones" of Blignaut and Furter (1940). Lower deltaic interval. Van Vuuren (1983), Winter (1985) and Cairncross and Cadle (1987) recognised two genetic sequences with up to five upward-

coarsening cycles in total below the top of the fluvial succession along the eastern part of the area (Figs. 17, 19). The fluvial succession forms the top part of the upper sequence. Le Blanc Smith and Eriksson (1979) interpreted the lowermost cycle (below the No. 2 Coal Seam) near the northemmost margin of the basin as the product of Gilbertian deltas resuiting from glacial outwash streams entering fresh water lakes. Bottomsets consist of fine-grained sandstone which grades downward into lacustrine pebbly mudstone with sandstone and siltstone lenses. The foresets are up to 5 m thick and comprise conglomerates and sandstones topped by distributary channel lag conglomerates and sandstones. In the extreme north upward-coarsening cycles are replaced by very coarse-grained sandstone to conglomerate fluvial sequences, a consequence of proximity to the source area (Cairncross, 1979; Winter, 1985; Fig. 19). In the southeast individual cycles attain a thickness of up to 80 m (Fig. 20). Regional marker

292

M.R. JOHNSON et al.

Fig. 19. Generalised (composite) section of the northernmost part of the Vryheid Formation (east of Johannesburg) illustrating lateral lithological variations (after Winter, 1985).

horizons in the easternmost part are scarce and correlation is extremely difficult. The shoreline appears to have advanced and retreated according to proximity of loci of sediment input, which varied due to delta switching, giving rise to a complex sequence of interfingering and overlapping delta lobes. Correlation of genetic sequences is however

Fig. 20. Lithology and depositional environments of the Vryheid Formation at Tugela Ferry 140 km northeast of Durban (after Hobday, 1973).

THE FORELAND KAROO BASIN, SOUTH AFRICA easier further west (Van Vuuren, 1983). Transgressions are typically recorded by intensely burrowed delta plain deposits. The lowermost cycle was commonly formed by outbuilding of Gilbertian deltas from the northeast. Successive phases of progradation produced foresets up to 35 m thick comprising medium to coarse feldspathic sandstone separated by dark silty drapes. Bottomsets are tangential, finergrained and burrowed. The topsets are truncated by distributary channel scours infilled by planar to trough cross-stratified pebbly sandstone. Roberts (1986) proposed that the homopycnal flows necessary for the formation of these structures developed via dilution of seawater around distributary mouths during floods. Giant foresets near Vryheid have also been interpreted as offshore sand ridges (Smith and Tavener-Smith, 1987). Delta cycles above the lowermost cycle conform to the model outlined above. However, successive units become more regressive upwards, with distributary mouth bar and channel sediments becoming more prominent. Delta front sandstones in the uppermost cycle are also cleaner, better sorted and in places represent the nondeltaic progradational sequences described above (Van Vuuren, 1981; Roberts, 1986). This evidence of wave reworking signalled a phase of major regression as also recorded by the fluvial succession. Fluvial interval. The fluvial succession comprises one or more superimposed upward-fining cycles of the type illustrated in Fig. 18. It attains its maximum development in the east where up to six cycles are present (Fig. 21) with a total thickness of about 60 m. Palaeocurrent patterns indicate a northeasterly source, hence the southwestward attenuation of the succession as a result of gradation into deltaic sediments. Most of the economically significant coals are associated with the fluvial succession. Most workers (e.g. Le Blanc Smith, 1980; Cadle et al., 1993) attributed the coals to colonisation of broad abandoned alluvial and upper delta plains by plants, although backswamp coals may also be present (Van Vuuren, 1981). The cyclicity of the fluvial sediments has been attributed by the majority of researchers (e.g. Cairncross, 1979; Le Blanc Smith, 1980; Roberts 1986) to vertical aggradation in multichannel systems followed by avulsion. When slope advantage was regained due to subsidence, the aggradational phase was repeated. Finer-grained burrowed sediments replace the fluvial sequence in some areas towards the southeast, suggesting the presence of marine embayments in these localities (Van Vuuren, 1983; Roberts, 1986). The fluvial interval is present as far south as Tugela Ferry (Hobday, 1973), but is barren of coals here. This is attributed to basin instability and frequent marine inundation by Van Vuuren and Cole (1979).

293

Fig. 21. Lithology of the Vryheid Formation in the type area west of Vryheid, 250 km NNE of Durban (after Roberts, 1986). The lower zone, coal zone and upper zone are equivalent to the lower deltaic, fluvial and upper deltaic intervals described in the text. Note that the upper zone is incomplete at this locality.

294

M.R. JOHNSON et al.

Upper deltaic interval. A major transgression terminated deposition of the fluvial interval, marked by glauconitic, medium-grained, burrowed sandstones above the No. 4 coal seam along the eastern part of the basin (Fig. 19). The sandstones in the upper deltaic interval are somewhat more feldspathic, micaceous and finer-grained than those of the lower deltaic succession, suggestive of reduced wave reworking. The No. 5 coal seam (Fig. 19) is succeeded by a glauconitic sandstone interpreted by Cadle and Hobday (1977), Le Blanc Smith (1980) and other workers as a transgressive sheet deposit. The No. 5 seam apparently formed in back barrier settings in places. Extensive torbanite lenses are associated with this coal seam, recording algal blooms in backbarrier lakes (Christie, 1988). The sequence above the No. 5 seam has been denuded over most of the northern area of the basin, but where preserved it comprises several upward-coarsening cycles, similar to those below although somewhat finer-grained. In the southeast the transgression which terminated deposition of the fluvial succession is marked by pro-delta silts and sands. Up to six upward-coarsening deltaic cycles which are similar to those in the north although thicker and finer-grained have been recorded in the upper deltaic interval (Van Vuuren, 1983; Roberts, 1986). As in the lower deltaic succession, marker horizons are scarce and correlation between cycles is difficult. Coals are rare and discontinuous. In the far south and southwest where the fluvial sequence is absent, the upper and lower deltaic successions merge and become indistinguishable. Van Vuuren (1983) and Roberts (1986) noted that the ichnofauna of the lower deltaic succession was marked by great diversity and dense population in contrast to the upper deltaic interval where well-defined structures are rare. Possibly a change in basin salinity may explain this observation; alternatively, higher sedimentation rates in the upper deltaic succession may have inhibited the invertebrate fauna. Western facies. The number of genetic sequences decrease concomitantly with the decrease in the thickness of the Vryheid Formation southwestward and towards the northwestern margin of the basin where the formation usually comprises two genetic sequences, each consisting of one or more upwardcoarsening cycles (Van Vuuren and Cole, 1979; Van Vuuren, 1983; Fig. 22). The sequences in this area are much better defined and correlation is easier than in the east. The sandstones are furthermore better sorted and finer-grained and the sediments in general more extensively bioturbated. Van Vuuren (1983) found deposits resembling the storm-beach sequence of Behr (1965) and Vos and

Fig. 22. Typical borehole section of the Vryheid Formation in the west (southwest of Johannesburg) (after Van Vuuren, 1983). Hobday (1977) to be fairly widespread in this area, constituting the entire Vryheid Formation in places. Fair weather beach sequences are however also fairly common. The bulk of the Vryheid Formation here

THE FORELAND KAROO BASIN, SOUTH AFRICA is believed by Van Vuuren (1983) to be of waveinfluenced deltaic origin. However, Stavrakis (1986) reported fluvial sediments along the northwesternmost margin of the basin. The Vryheid Formation rests on a very uneven floor of pre-Karoo rocks and Dwyka Group sediments along the northwestern margin of the basin. The rugged topography provided sheltered environments for the development of coal swamps. These include glacially scoured valleys blocked by moraines and lagoons enclosed by islands of pre-Karoo rocks in conjunction with beach ridges or barrier bars. Volksrust Formation. The Volksrust Formation is a predominantly argillaceous succession confined to the northeastern part of the basin (Fig. 3). It interfingers with the overlying Beaufort Group and underlying Vryheid Formation. Where the latter pinches out towards the southwest the Volksrust becomes the Tierberg Formation (in the northern outcrop area) or merges with the Pietermaritzburg Formation in the undifferentiated Ecca Group in the southeast (Fig. 3). Drilling has shown that it reaches a thickness of 380 m about 120 km northeast of B loemfontein, thinning to 250 m towards the east (SACS, 1980; Taverner-Smith et al., 1988) and to 100 m towards the northern margin of the basin (Van Vuuren, 1983). The formation consists of grey to black silty shale with thin siltstone or sandstone lenses, laminae and beds particularly towards its upper and lower boundaries. In outcrop it is mainly finely laminated although structureless beds, believed to be the result of intense bioturbation, are also present and may predominate in places. Ripple lamination and convolute bedding occur sporadically. The silty/sandy interbeds are usually bioturbated. Thin phosphate and carbonate beds and concretions, some of which have a lateral extent of up to 30 m (Linstrrm, 1981) are fairly common in the formation. Tavemer-Smith et al. (1988) reported siderite concretions up to 1.5 m in diameter and beds up to 0.75 m thick. Although the Volksrust Formation is widely believed to be an open "shelf" sequence, TavernerSmith et al. (1988) concluded from a detailed study of an isolated outcrop northeast of Durban that its upper and lower parts were deposited in lacustrine to possibly lagoonal and shallow coastal embayment environments. The presence of palaeosols and plant remains (Glossopteris, Phyllotheca) and trace fossil types were taken as evidence of shallow water environments. It should however be pointed out that the section measured is relatively proximal to the provenance, which is believed to have been located towards the northeast. Deeper water conditions probably prevailed throughout deposition of the formation towards the southwest.

295

Beaufort Group Adelaide Subgroup In the southeastern part of the basin the Late Permian Adelaide Subgroup comprises the Koonap, Middleton and Balfour Formations, in ascending order. In the western part of the basin the Abrahamskraal and Teekloof Formations are the approximate equivalents of the Koonap and Middleton Formations respectively (Figs. 2, 4). While the Middleton and Teekloof Formations are characterised by a greater relative abundance of red mudstone compared to the underlying and (in the case of the former) overlying units, in practice the boundaries are linked to specific sandstone-rich marker units (members). In the northeastern region the Estcourt and Normandien Formations are laterally equivalent units within the subgroup (Figs. 2, 4), although the Estcourt Formation will probably in future be incorporated in the Normandien Formation. From a maximum of 5000 m in the southeast (Johnson, 1976), the thickness of the Adelaide subgroup decreases rapidly to about 800 m in the centre of the basin and thereafter more gradually to around 100-200 m in the extreme north (Groenewald, 1989) (Fig. 23). The Koonap Formation attains a maximum thickness of about 1300 m, the Middleton 1600 m (although it may be as much as 2500 m north of Port Elizabeth) and the Balfour 2000 m (Johnson, 1976). In the west the Abrahamskraal and Teekloof Formations are up to 2500 m and 1000 m thick respectively. The Normandien Formation attains a maximum of 320 m (Groenewald, 1984) while the Estcourt Formation is up to 500 m thick (Linstr~m, 1981). In the southern and central parts of the basin the Adelaide Subgroup consists of alternating bluishgrey, greenish-grey or greyish-red mudrocks and grey, moderately to well-sorted, very fine to medium-grained lithofeldspathic sandstones. In the northem part of the basin coarse and very coarse sandstones, or even granulestones, are also common in the Normandien Formation (Groenewald, 1984). Sandstone generally constitutes 20-30% of the total thickness, but in certain areas may be as little as 10% while some sandstone-rich intervals may in places contain up to 60% sandstone. Except in the lower part of the Estcourt Formation, where coarsening-upward cycles are common, the sandstone and mudrock units normally form fining-upward cycles. These cycles rest on erosion surfaces displaying a few centimetres to a few metres of relief. In many cases a thin intraformational mud-pellet conglomerate is present at the base of the sandstone unit. The cycles vary from a few metres to a few tens of metres in thickness. Individual sandstone units are thickest (average 6 m; maximum 60 m) in the extreme southeast and

296

M.R. JOHNSON et al.

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west and become thinner northwards, except for the extreme northeast where thick, extensive units are also present in the Normandien Formation. They are subtabular to moderately lenticular (occasionally highly lenticular) and generally extend laterally for a few hundred metres to a few kilometres. Thicker sandstones tend to be multistorey, with cut-andfill features (including mud drapes) being common. Internally the sandstones are characterised by horizontal lamination accompanied by parting lineation and less abundant trough cross-bedding and ripple lamination. Massive beds may occur towards the base of sandstone units. Ripple lamination tends to be confined to thin sandstones or the finer material towards the top of thicker units. Small oscillation ripples with an average wave length of 2.5 cm occur and appear to be oriented roughly parallel to the general palaeocurrent direction (Johnson, 1976). Calcareous concretions 20-100 cm in diameter are present in some sandstones. Current ripples are common in the Abrahamskraal Formation (Theron, 1983). Stear (1978, 1980) describes and illustrates a variety of ripple marks and other structures from this formation in the Beaufort West area. In the Daggaboersnek Member towards the middle of the Balfour Formation the sandstones tend to be thin and tabular, with cross-bedding rare or absent. The wave ripples which are fairly common in this unit are appreciably larger than those occurring elsewhere in the subgroup, with an average wave length of about 7 cm (Johnson, 1976). The mudrocks in the Adelaide Subgroup are generally massive and blocky-weathering except in the Estcourt Formation and Daggaboersnek Member where horizontal lamination is common. Desiccation cracks and raindrop impressions are occasionally present. Calcareous nodules and concretions, varying considerably in size and shape from subrounded lumps a few cm in diameter to flattened, irregular discontinuous bodies a few metres long occur in mudstones throughout the Beaufort Group. In the Abrahamskraal Formation brown-weathering limestone (palaeocaliche) layers are up to 1.5 m thick and extend up to 2 km. A number of greenish grey, cherty layers a few cm to two metres thick and extending a few tens of kilometres in some cases are also present in this formation. Most are massive, but tipple lamination, bioturbation and ripple marks are not uncommon (Theron, 1983). It appears that at least some of these layers represent reworked, silicified volcanic ash (Martini, 1974). Vertebrate fossils comprising numerous reptilian genera as well as amphibian and fish remains are common in the Adelaide Subgroup. Nonmarine molluscs, invertebrate burrows and trails, silicified wood and stem impressions occur sporadically throughout the subgroup. Well-preserved leaf impressions

THE FORELAND KAROO BASIN, SOUTH AFRICA

297

are present in places, but are common in the Daggaboersnek Member (Johnson, 1976) and Escourt Formation, the latter unit also having yielded a varied insect assemblage (Riek, 1973, 1976). Palaeocurrent data for the Adelaide Subgroup are summarised on Fig. 24. From this data it would appear that the bulk of the sediment was derived from a source area situated to the south and southeast of the basin, with subordinate influx from the southwest, west-northwest and northeast. In the area east of B loemfontein Theron (1970) postulated an intrabasinal source, the Clocolan Dome, which shed detritus towards the southwest. The ubiquitous presence of fining-upward cycles, the terrestrial biota, the abundance of red mudrocks and a characteristic suite of sedimentary structures point unequivocally to deposition under fluvial conditions. The high mud-sand ratios and fine-grained character of the sandstones indicate meandering rather than braided rivers, with the sandstones having formed as bed load channel (point bar) deposits and the mudstones representing suspended material dumped on the adjacent floodplains (flood basins) during overbank flooding. In the southwestern part of the basin Turner (1978) distinguished between a low sinuosity channel facies association (high sandstone : mudstone ratio), a high sinuosity channel facies association (intermediate sandstone:mudstone ratio) and a floodbasin facies association (low sand-

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Typical fluvial characters are, however, uncommon in the Estcourt Formation and the Daggaboersnek Member. In both cases the preponderance of rhythmites and shale (with plant remains) and thin tabular sandstones rather than massive mudstones and thick, lenticular sandstones as well as the absence of classic fining-upward cycles point to largely subaqueous deposition under "lacustrine" conditions. Small deltas probably prograded into this "lake" (Hobday, 1978). In the Daggaboersnek Member the presence

of numerous moderately sized wave ripples displaying a mean E N E - W S W orientation - - approximately at right angles to the regional palaeoslope as deduced from current directions in the overlying and underlying s t r a t a - also suggest deposition in a fairly extensive inland sea or lake (Johnson, 1976). In both cases the dark, carbonaceous shales with plant remains presumably accumulated in coastal or deltaic marshes, swamps and interdistributary bays.

THE FORELAND KAROO BASIN, SOUTH AFRICA

299

Modal analyses of 52 sandstone samples collected from the Adelaide Subgroup in the southeastern part of the basin (Johnson, 1976) gave the following average composition: quartz: 19%; feldspar: 28%; lithic fragments: 36%; matrix: 12%; accessories: 1%; cement: 4%. No significant differences were apparent between the Koonap, Middleton and Balfour samples. Volcanic rock fragments constitute over 20% of the total and micaceous-schistose fragments 9%. As in the case of the sandstone-rich units of the Ecca Group in this area, the samples fall in the magmatic arc provenance field of Dickinson et al. (1983) (Johnson, 1991; Fig. 10). It is concluded that the Ecca Group and Adelaide Subgroup in the southern and southeastern part of the basin were derived from the same source. The sandstones in the southwestern part of the basin are also characterised by high feldspar and rock (including volcanic) fragment content (Theron, 1983) and were probably derived from a westward extension of the above provenance. Those sandstones which show an east-southeast palaeocurrent direction in the extreme west-northwest of the basin (Fig. 24) may have been derived from granite-gneiss of the Namaqua-Natal Belt in Namaqualand (Toens and Le Roux, 1978). In the northeastern part of the basin Groenewald (1984) postulated a granitic source area situated to the east of the basin for the Normandien Formation.

Tarkastad Subgroup The Early Triassic Tarkastad Subgroup is characterised by a greater abundance of both sandstone and red mudstone than the Adelaide Subgroup. The boundary between these subgroups is the only one in the Beaufort Group that can be traced throughout the Karoo Basin. The subgroup attains a maximum thickness of close on 2000 m in the south (Johnson, 1976). This reduces to around 800 m in the middle of its outcrop area and it ultimately thins to 150 m or less in the far north (Groenewald, 1984) (Fig. 23). In the south the Tarkastad Subgroup comprises a lower Katberg Formation (sandstone-rich) and an upper Burgersdorp Formation (mudstone-rich). Sandstone constitutes over 90% of the Katberg Formation in the small coastal exposures (where it is over 900 m thick) and the southernmost part of the main outcrop area (Fig. 26). However, the sandstone:mudstone ratio decreases rapidly northwards until the formation becomes difficult to distinguish from the Burgersdorp Formation. The latter is around 1000 m thick in the southern outcrop area, with the overall sandstone content diminishing from possibly 50% in the coastal exposures to around 20-30% or less further north in the main outcrop area (Fig. 27). In the northeastern part of the basin west of Durban (Fig. 2), Botha and Linstrrm (1978) defined a

Fig. 26. Representative sections of the lower (A) and upper (B) parts of the Katberg Formation in the type area northeast of Port Elizabeth (after Stavrakis, 1979). lower Belmont Formation, up to 165 m thick, and an upper Otterbum Formation, between 130 and 180 m thick. The Belmont Formation consists of sand-

300

Fig. 27. Type section of the Burgersdorp Formation near Queenstown, northeast of Port Elizabeth (after Johnson and Hiller, 1990). Note that only the middle part of the formation is represented.

stone (70%) interbedded with predominantly red mudstone, whereas the Otterburn Formation consists of red mudstone and subordinate green mudstone interbedded with minor sandstone (25%). In the extreme northeast Groenewald (1984) recognised a lower Verkykerskop Formation comprising up to 80 m of fine- to very coarse-grained sandstone and an upper Driekoppen Formation which is up to 70 m thick and consists almost entirely of mudstone. Compared with those in the underlying Adelaide Subgroup, sandstones in the Katberg Formation are coarser-grained (ranging up to medium-grained, with scattered pebbles up to 15 cm in diameter present in the coastal outcrops) and lighter in colour (light brownish grey or greenish grey). Horizontal lamination plus parting lineation, trough cross-bedding and

M.R. JOHNSON et al. planar cross-bedding are the main internal sedimentary structures (Fig. 26). Oval to spherical calcareous concretions 3-10 cm in diameter are common, with the former displaying a preferred orientation parallel to the regional palaeoslope (Johnson, 1989). Deformed cross-bedding is present in places (Hobday, 1978). Burgersdorp Formation sandstones are finegrained, greenish grey or light brownish grey and display horizontal lamination, cross-bedding and ripple lamination (Fig. 27). In the middle part of the formation in the main outcrop area they average 2 m in thickness. Intraformational mud-pellet conglomerates are common in both the Burgersdorp and Katberg Formations. Red colours predominate in the mudstones of both formations. Reptile, amphibian and (to a lesser extent) fish remains are fairly common in the Tarkastad Subgroup, while plant stems are sporadically present. Except for the Katberg Formation in the south the sandstone and mudstone units in the Tarkastad Subgroup tend to form fining-upward cycles comparable to those in the Adelaide Subgroup. A meandering river palaeoenvironment may therefore have predominated. The southern Katberg Formation, however, displays features indicative of a braided stream environment, namely relatively coarse grain size, virtual absence of interbedded mudstone layers and hence lack of distinct fining-upward cycles composed of alternating channel and floodbasin deposits, rarity of ripple lamination and low scatter of palaeocurrent data (Johnson, 1976). The fan-shaped palaeocurrent pattern displayed by the coastal occurrences led Hiller and Stavrakis (1980) to suggest that these represent deposits formed by braided streams on the distal reaches of an alluvial fan. The preponderance of red colours in the mudrocks indicates that more arid conditions prevailed than was the case during deposition of the Adelaide Subgroup. In the northeastern part of the basin Botha and Linstr6m (1978) concluded that the sediments of the Belmont Formation were deposited in a braided river environment and those of the Otterburn Formation in a meandering river environment. In the area further north, Groenewald (1984) deduced a braided river environment for the coarse-grained Verkykerskop Formation. The overlying fine-grained Driekoppen Formation (Fig. 23) was interpreted as a distal meandering river facies of the underlying unit. Palaeocurrent data for the Tarkastad Subgroup suggest a northerly to northwesterly regional palaeoslope similar to that which prevailed during deposition of the underlying Adelaide Subgroup (Fig. 24). Once again the major source area must have been located to the southeast of the basin. The composition of the Tarkastad Subgroup sandstones in the southeastern outcrop area (Johnson,

THE FORELAND KAROO BASIN, SOUTH AFRICA

301

1976) show significant differences compared with the Adelaide Subgroup (Fig. 10). For the Katberg 14 samples gave an average quartz content of 35% (nearly double that of the Adelaide Subgroup) and feldspar content of 14% (half that of the Adelaide). Rock fragments constitute 38%, matrix 8%, accessories 1.5% and cement 1.5%. For the overlying Burgersdorp Formation the following average composition, based on 10 samples, was obtained: quartz: 42%; feldspar: 9%; rock fragments: 34%; matrix: 14%, cement: 1%; accessories: 0.5%. Volcanic fragments continue to be an important component of both the Katberg and Burgersdorp Formations, constituting 17% of the total rock composition in the case of the former and 13% in the case of the latter. Quartzite, quartzitic sandstone and "arkose" together constitute nearly 50% of the pebbles in the coastal occurrences of the Katberg Formation. In order of abundance the remainder comprise devitrifled lava, quartz-feldspar porphyry, granite/gneiss and chert/silicified wood (Johnson, 1976). The rapid increase in quartz content at the expense of feldspar in the Katberg Formation would appear to reflect strong uplift and denudation of a fold-thrust belt, comprising mainly Cape Supergroup strata, located between the magmatic arc and the basin (Johnson, 1976, 1991). As a result the role of the igneous component steadily diminished with time. The preponderance of red and greenish rather than white quartzites suggest that at this stage pre-Cape rather than Cape Supergroup strata were being subjected to erosion. For his Verkykerskop Formation in the northeastern part of the basin Groenewald (1984) postulated the continued existence of a granitic source lying to the east. Molteno Formation

The Molteno Formation comprises a Late Triassic (Turner, 1983) northward-thinning, foreland wedge of dominantly fluvial sediments. The formation attains a maximum thickness of about 600 m in its southern outcrop area, where it can be subdivided into five members (Turner, 1975; Christie, 1981) (Fig. 28). In the extreme north it is less than 10 m thick. Deposition was predominantly by bedload-dominated rivers flowing from a tectonically active source area situated to the south and southeast (Turner, 1975, 1983; Christie, 1981), believed to represent either accelerated uplift of the Cape Fold Belt (Tankard et al., 1982) or basin margin faulting associated with initial rifting prior to the separation of the Falkland Plateau and southeast Africa (Turner, 1983). In the south the Bamboesberg Member forms the base of the Molteno Formation (Fig. 28) which here conformably overlies the Beaufort Group. Elsewhere, the Indwe Sandstone Member, which is the

Fig. 28. Typical section of the Molteno Formation in the southern outcrop area (Fig. 2, Section D) (after Christie, 1981).

only representative of the Molteno Formation in the north, constitutes the base and appears to rest unconformably on the Beaufort Group (Turner, 1975). The Bamboesberg Member is a composite succession of upward-fining sequences comprising erosively-based, fine- and medium-grained sandstone beds and thin, lenticular mudrock intercalations. Two coal seams are present. The base of the Indwe Sandstone Member is defined by a distinctive, erosively-based pebble and cobble horizon. The bulk of the Indwe Sandstone comprises 2-8-m-thick sequences fining upward from very coarse-grained, pebbly to coarse- or medium-grained sandstones

302 (Christie, 1981). Less resistant, medium-grained sandstone composing the upper part of the member fines upward into the overlying Mayaputi Member. The Mayaputi Member is a dominantly argillaceous unit, 15-50 m thick, containing thin, lenticular, finegrained sandstone beds. The coal seams present are discontinuous and generally shaly, occurring near the top of the member which is terminated by the erosively based Qiba Member. The Qiba Member comprises a monotonous succession of upward-fining, fine- to medium-grained sandstone beds with thin mudrock partings. It varies in thickness from 20 to 58 m and contains a shaly, erratically developed coal seam 20-40 m above its base. The Tsomo Member displays a repetitive pattern of erosively-based, coarse-grained, pebbly sandstones up to 25 m thick alternating with mudrock units which may be as much as 60 m thick. Thin, lenticular coal seams are sporadically developed. The Bamboesberg Member was deposited by sandy, ephemeral streams which were broad and shallow (Christie, 1981, 1986). In contrast to the warm, arid climate prevailing during Beaufort Group sedimentation, the climate was wet and cool. Peat accumulated in alluvial-plain swamps removed from the locus of fluvial activity. While the Bamboesberg Member is believed to have been the product of an epeirogenic phase of moderate uplift in the source area, Indwe Sandstone deposition was initiated by rapid and significant uplift which resulted in an increase in the gradient and competency of streams (Turner, 1975). Large amounts of coarse detritus were released and deposition was by high-energy, coalescing, braided streams characterised by vertical channel aggradation and rapid channel shifting. The rivers drained an extensive alluvial plain which may have constituted the distal parts of an alluvial fan complex (Turner, 1975, 1983; Christie, 1981). Decrease in sediment supply and river size allowed the development of extensive floodplains responsible for the Mayaputi Member. The coals present at the top of the member represent localised peat swamps. The Qiba Member was deposited by shallow, high-energy ephemeral streams similar to those active during Bamboesberg Member time. The Tsomo Member reflects periods of intermittent tectonic activity which resulted in the deposition of coarse-grained, braided-river sheet sandstones alternating with periods of little fluvial activity during which extensive lacustrine and floodplain silts and clays were deposited. Evidence of an increasingly warm, arid environment is provided by the appearance of maroon and green mudstones at the top of the unit. Conditions were, however, still favourable for the formation of localised peat swamps. It is suggested that the upper part of the Molteno Formation passes northwards into sediments of the Elliot For-

M.R. JOHNSON et al. mation deposited within an interior drainage basin (Turner, 1975; Christie, 1981). The relatively quartz-rich nature of the Molteno Formation sandstones compared with the underlying units (Fig. 10), together with the absence of clasts other than quartzite suggests that uplifted Cape Supergroup rocks in the Cape Fold Belt had by now replaced the magmatic arc as the main provenance for the Karoo Basin (Johnson, 1991). Elliot Formation

The Late Triassic Elliot Formation comprises an alternating sequence of fine- to medium-grained sandstone and mudrock (Fig. 29). It attains a maximum thickness of about 500 m in the south (Visser and Botha, 1980) and overlies the Molteno Formation conformably and gradationally. The maroon and green mudrock units typically range in thickness between 25 and 100 m and contain subordinate beds and lenses of fine-grained sandstone. Desiccation cracks and occasional vertebrate footprints are present. The sandstone layers are yellowish grey to pale red, up to 22 m thick (average 6-15 m) and grade laterally into mudrock. Their bases are characteristically erosional and contain numerous mudrock intraclasts. Flat bedding and trough cross-bedding predominate. Botha (1968) and Visser and Botha (1980) interpreted the Elliot Formation in terms of major sand-laden rivers separated by vast floodplains. Initially the rivers had meandering channels, but progressive warming and aridity resulted in the rivers becoming broader, shallower and more ephemeral. Palaeocurrent patterns suggest northerly to westerly transport. The uppermost part of the Elliot Formation contains evidence of aeolian processes (Visser and Botha, 1980). Red colouration of the mudrocks is attributed to the presence of finely disseminated iron oxide formed by the early diagenetic oxidation of ferrous iron, whereas the green sediments may have resulted from the reduction of ferric iron by migrating groundwater (cf. McBride, 1974). Clarens Formation

The final phase of Karoo sedimentation is represented by the Late Triassic/Early Jurassic Clarens Formation. Progressive warming and desiccation, which reached a climax during the Late Triassic, is reflected by fine-grained aeolian sand and associated playa lake, sheetflood and ephemeral stream deposits (Beukes, 1970; Eriksson, 1981). Beukes (1969, 1970) divided the northern Clarens Formation, which is generally in the order of 100 m thick, into three "zones" (Fig. 30). The central zone represents true desert conditions dominated by an

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303 to mid-latitude westerly winds (Bigarella and Van Eeden, 1971). The upper and lower zones reflect wetter environments characterised by shallow playa lakes (which were host to fish, crustaceans and small dinosaurs) and ephemeral rivers. In the southem outcrop area the formation comprises up to 300 m of homogenous siltstone and silty fine-grained sandstone, interpreted as loess deposits (Johnson, 1976). Minor lava flows interlayered with sediments in the uppermost part of the Clarens signal the termination of Karoo sedimentation.

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A prolonged outpouring of basalts during the Early Jurassic (c. 180 Ma - - Hooper et al., 1993) produced vast lava fields, with the Drakensberg plateau (140,000 km 2) being an erosive remnant of a once more extensive and thicker suite of volcanic rocks. This episode of volcanism lasted for a relatively short period and preceded the breakup of Gondwanaland. The remnants of these flood basalts attain a maximum thickness of 1400 m (Eales et al., 1984). Considering that the Karoo sedimentary strata accumulated essentially without interruption for some 150 Ma, the transition into the volcanic sequence is remarkably rapid. However, the history of early volcanism was complex, with minor pyroclastic rocks occurring in the upper part of the Elliot Formation (Botha and Theron, 1967). Within the upper Clarens Formation volcanism became more widespread and there is considerable interfingering of volcanic and clastic deposits (Botha and Theron, 1967; Lock et al., 1974). Some of the lava flows were covered by sand on which dinosaur tracks are preserved (Rust, 1975). By all accounts, the pre-volcanic surface displayed considerable relief (in the order of 100 m); volcanic flows are seen to wedge out against large dune features and fill basins and palaeovalleys (Lock et al., 1974). Eales et al. (1984) described early volcanic activity as being confined to discrete centres (vents and diatremes) around which lava shields were built up. During periods of explosive activity, thick sequences of pyroclastic rocks were deposited over wide areas. Evidence that aqueous environments existed during the early stages of lava extrusions is provided by deposits of playa lakes, pillow lavas and lava-pod complexes. Shortly after the start of volcanism the main phase of fissure eruptions was initiated, rapidly building into a thick succession of horizontal flows which extended over large parts of southern Africa (Marsh and Eales, 1984). The bulk of the lavas were basaltic (tholeiitic) in composition, of the pahoehoe type and amygdaloidal (Eales et al., 1984). Individual flows varied in thickness from less than half a metre to more than 50 m.

304

M.R. JOHNSON et al.

Fig. 30. Representative sections of the Clarens Formation (after Beukes, 1970). The Karoo Basin sedimentary sequence is extensively intruded and locally metamorphosed by dolerite sills, dykes and irregularly shaped bodies, except for a portion along the southern margin compressed by the Cape Fold Belt (Tankard et al., 1982). Dolerite intrusions are rare in the pre-Karoo basement and appear to terminate at a critical distance of a " . . . few thousand feet below the basalt" (Winter

and Venter, 1970). The inference is that the basic magma feeding the lavas moved into the upper crust along narrow, localised zones. Many thick sills show subtle internal differentiation produced by gravity settling of phenocrysts (Eales et al., 1984). Less commonly, pronounced differentiation produces igneous lamination and cryptic layering. As an example, layering in the Insizwa

THE FORELAND KAROO BASIN, SOUTH AFRICA sheet intrusion in Transkei is well defined and comprises an olivine-rich, ore-bearing basal zone, a gabbro and norite central zone, and an acid roof zone (Scholtz, 1936).

BASIN HISTORY Tectonic setting

Infilling of the Karoo Basin was preceded in the south by the deposition of up to 8 km of Ordovician to Early Carboniferous strata belonging to the Cape Supergroup in the Cape Basin. This succession comprises mudrocks and quartzitic sandstones which form a southward-thickening passive margin wedge (Fig. 31a). Along the eastern flank of the basin accumulation of the Natal Group (quartzitic sandstones up to 1 km thick) took place in an elongate downwarp (the Natal Embayment) which was probably the failed arm of a triple junction (Tankard and Hobday, 1979). The Karoo Basin can readily be classified as a foreland basin, since it contains a thick flyschmolasse wedge which flanks the front of a mountain chain and wedges out northward over the adjacent craton (Figs. 1, 3, 31). The basin constitutes a retro-arc foreland basin as defined by Dickinson (1974), situated behind a magmatic arc and associated fold-thrust belt (Cape Fold Belt) produced by northward subduction of oceanic lithosphere located south of the arc (Johnson, 1991; Fig. 31b). Maximum subsidence, as reflected by thickness of basin fill, took place in a linear belt (the Karoo Trough) situated along the southern edge of the basin. This trough is underlain by a zone of high electrical conductivity termed the Southern Cape Conductive Belt (De Beer et al., 1982; Fig. 1). The high conductivity is thought by these authors to reflect the presence of denser oceanic basalt which had been obducted against the continental margin during the Late Proterozoic Pan-African orogeny. A subordinate trough along the eastern edge of the basin (the Natal Trough) probably represented reactivation of the early Palaeozoic zone of rifting (Tankard et al., 1982; Cole, 1992). A 30 Ma erosional period (Namurian-Westphalian) followed deposition of the Cape Supergroup (Visser, 1987) before renewed subsidence accompanied by sedimentation in the form of glacial deposits (Dwyka Group) took place in the Karoo Trough. A longer hiatus (Silurian-Late Carboniferous?) is probably present in the Natal Trough due to abortion of the rift before reaching the spreading stage (Hobday and Von Brunn, 1979). A northward shift of the Karoo Trough axis (Dingle et al., 1983) and a westward shift of the Natal Trough axis (Tankard

305 et al., 1982) occurred between early Palaeozoic and Permian times. The Karoo Basin developed from the Late Carboniferous onward by gradual subsidence of essentially cratonic crust. Visser (1987) estimated that an ice sheet up to 4200 m thick could have covered the basin floor during the Late Carboniferous to Early Permian. This ice load may have been a contributory factor in the initial subsidence of the Karoo Basin (Visser, 1987). Loading of the continental margin through folding and thrusting in the Cape Fold Belt as well as sedimentary loading may later have become important mechanisms for maintaining basin subsidence, particularly in the Karoo Trough where up to 10 km of Karoo sediments were deposited. The possible presence of denser material at the base of the crust, under the Southern Cape Conductive Belt (De Beer, 1983), could also have aided subsidence (Cole, 1992). Along the southern edge of the basin the Cape Fold Belt developed while sedimentation of at least the upper Karoo units was still in progress. This orogeny resulted in intense deformation of the Cape Supergroup and underlying basement as well as the lower units of the Karoo Supergroup. The Cape Fold Belt partly coincides with the Southern Cape Conductive Belt (Fig. 1), a zone of weak crust sandwiched between rigid crust to the north (Namaqua-Natal metamorphic belt) and south (Saldanian Province granite-intruded sedimentary rocks). De Beer et al. (1982) suggested that the Cape Fold Belt was caused by compression from the south which was initially transmitted by the rigid block (Saldanian Province). 4~ step heating analysis of cleaved pelites from the Cape Fold Belt indicated four main episodes of intense folding and thrusting (H~ilbich, 1983; H~ilbich et al., 1983). Dates of 277 -+-5 Ma, 259 -+-5 Ma, 246 -t- 2 Ma and 229 + 5 Ma were obtained for these paroxysms. Du Toit (1937) proposed that the Cape Fold Belt formed a segment of the Gondwanide Orogen, which included the Sierra de la Ventana of Argentina and the Trans-Antarctic Mountains (Fig. 32). De Wit (1977) suggested that Andean-type subduction, whereby the palaeo-Pacific oceanic plate descended below the Gondwana continental plate (Fig. 32), was the cause of the Gondwanide Orogen. Lock (1980) attempted to explain the excessive distance of the Cape Fold Belt from the subduction zone by means of a fiat-plate subduction model. This model involves the attachment of the subducted plate to the Gondwana plate and lateral compression and folding along the line of separation in the north. A magmatic arc associated with the subduction zone has been identified in southern South America (Forsythe, 1982) and the Antarctic Peninsula (Smellie, 1981) (Fig. 32). The silicic volcanic ash layers in the Collingham Formation and ubiquitous volcanic rock fragments

306

M.R. J O H N S O N et al.

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EMFB EIIsworth Mountains Fold Belt PMFB Pensacola Mountains Fold Belt Fig. 32. Gondwana reconstruction (after De Wit et al., 1988), showing the probable location of the magmatic arc, Gondwanide Orogen and other tectonic elements. Modified after Johnson (1991).

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308 in sandstones of the Ecca and Beaufort Groups in the southern Karoo Basin (Elliot and Watts, 1974; Martini, 1974; Johnson, 1976, 1991) may have been derived from this arc. However, the distance separating the postulated arc from the Karoo Trough (~1700 km) represents a problem, and it may be necessary to consider the possibility of an originally closer arc source located along the northern edge of the Falkland/Malvinas Plateau (Ramos, 1986; Johnson, 1991; Collinson et al., 1992; Fig. 32). The progressive northeastward decrease across the Karoo Basin in thickness and frequency of tufts in the Ecca Group led Viljoen (1987) to suggest a source in Patagonia, where Permian volcanic centres presently crop out. Mitchell and Reading (1986) point out that as the basin deepens many successions pass from a shallow water (pre-orogenic) phase through a starved basin (pre-flysch) phase to a deep marine clastic (flysch) phase and finally to a continental clastic (molasse) phase. The Cape-Karoo succession fits this classical geosynclinal cycle rather well, although a major glacial episode is here interposed between the preorogenic stage (Cape Supergroup) and pre-flysch stage (Prince Albert-Collingham Formations). Dwyka glaciers and ice sheets The Dwyka basin was located within the polar circle since the end of the Devonian (Smith et al., 1981). Decreases in temperature during the mid-Carboniferous caused the global Namurian regression (Veevers and Powell, 1987) as well as the build-up of an extensive ice cover over the southern mountain chain (fringing the palaeo-Pacific margin) and the cratonic highlands to the north and east. The onset of the major glacial depositional phase at about 300 Ma reflects the spreading of a large marine ice sheet across the basin. The areal extent of the basin was at least double its present size, and probably included the present Falkland Islands, situated at the southeastern comer of the Karoo Basin (Mitchell et al., 1986; Fig. 32). At first deposition took place largely from a grounded ice sheet, but the onset of slightly warmer conditions resulted in rain-out debris accumulating from a predominantly floating ice shelf. Further warming led to dissolution of the ice sheet, with glacial deposition confined largely to valleys bordering the cratonic highlands, and inundation of the basin by a shallow body of water into which argillaceous material was transported. Ecca seas and deltas The change-over from primarily glacial diamicton to offshore mud sedimentation (Prince Albert

M.R. JOHNSON et al. Formation) in the Permian is attributed to a climatic amelioration in the region and a major regional transgression. These resulted in a marine basin with a drowned northern margin. The basin, which was located between palaeolatitudes 50 ~ and 70 ~ (Smith et al., 1981), was greatly influenced by the influx of sediment-laden cold meltwaters from the glaciers on the highlands. Outwash fans formed where streams debouched into the basin, underflows carried mud and silt basinwards to settle as a blanket deposit on the bottom while possible mud fans built up along bottom depressions in the western and eastern sectors of the basin. Upwelling of cold water in certain areas enhanced a rich marine life. The abrupt change in depositional conditions between the Prince Albert and Whitehill Formations can be attributed to shallowing of the basin and termination of oceanic circulation which led to stratification of the water and reducing conditions during Whitehill time. This organic-rich mud facies together with fossils of the reptile Mesosaurus tenuidens are also present in the Warmbad, western Kalahari and Parana Basins which suggests that the areal extent of the interconnected basins may have been as large as 4.5 x 106 km 2. By the end of the Artinskian (~270 Ma) all the ice on the highlands had melted and a lush vegetation developed in the region which would have lowered the effluent as well as sediment supply to the basin. In the south the accumulation of organic-rich muds under low-energy conditions (Whitehill Formation) was followed by deposition of numerous thin air-fall tufts which periodically interrupted the suspension settling of mud (Collingham Formation). The sudden influx of coarser detritus from a provenance situated south of the present basin which followed this distal volcanic episode marked the transition from "pre-flysch" to flysch stages. Rapid downwarping of the Karoo Trough accompanied the build-up of sandy and silty submarine fans and basin plain turbidites of the Ripon, Laingsburg, Vischkuil and Skoorsteenberg Formations, which were deposited in a water body which according to Visser and Loock (1978) was up to 500 m deep. The turbidity processes in turn gave way to suspension settling of rhythmically bedded pro-delta mud (Fort Brown Formation) as the delta slope prograded across the turbidite fans at its foot. Turbidite deposition did not, however, extend very far into the basin and to the north of the Karoo Trough suspension settling of mud (Tierberg Formation) took place in relatively shallow water. In the northeastern part of the basin the Ecca sea was initially starved of sediment due to glacial cover in the source areas. As the climate warmed, detritus was carried mainly from a granitic highland situated towards the northeast and to a lesser extent

THE FORELAND KAROO BASIN, SOUTH AFRICA from a relatively low-lying quartzitic provenance to the north (Witwatersrand a r c h Fig. 1) and deposited as a prominent fluvio-deltaic wedge (Vryheid Formation). The presence of prominent coal seams in the Vryheid Formation points to luxuriant plant growth and also indicates that the organic-rich muds of the Whitehill Formation are probably a distal equivalent of this formation. Mud carried in suspension from the deltas was distributed further basinward and deposited as a thick blanket during the advances of the coastline (Pietermaritzburg Formation) and also during its retreat (Volksrust Formation). During this period several major transgressions caused by epeirogenic subsidence and/or sea level rises interrupted delta progradation, while shoreline sands formed in places. The climate was cold throughout deposition of the sequence. In the south the transition from flysch to molasse is denoted by the progradation of delta front sands (Waterford Formation) across pro-delta Fort Brown Formation muds. In the west and northwest deltaic sediments advanced more or less simultaneously into the Tierberg sea. Parts of these units may, however, represent coastline rather than deltaic deposits.

Beaufort alluvial plains Progradation of sandstone-rich delta front and lower delta plain sediments into the Ecca sea was followed by the subaerial deposition of upper delta plain and fluvial mud and sand of the Adelaide Subgroup during the Late Permian. Gradual denudation of the provenance caused a sourceward shift of the mixed-load fluvial deposits that were replaced by flood basin and lacustrine muds in the western part of the basin (Turner, 1985). In the northeast the Estcourt Formation provides evidence that deltaic conditions persisted locally for some time after they had ceased elsewhere. Palaeocurrent studies indicate that while the main source areas were located south and southeast of the basin, provenances to the west, north and east (Theron, 1975; cf. Fig. 24) also supplied sediments during the Late Permian. An intrabasinal provenance, the Clocolan Dome, may have been the source of localised coarse fluvial sands and granulestones prior to its burial by younger sediments (Theron, 1970). Strong uplift associated with the major Cape Fold Belt orogeny along the southern margin of the basin at the beginning of the Triassic led to the influx of medium-grained, pebbly, bed-load fluvial sandstones of the Katberg Formation. These sediments extended right across the basin, but with denudation of the provenance a sourceward shift of facies occurred, resulting in the overstep of bed-load fluvial deposits by mixed-load and flood basin deposits of

309 the Burgersdorp Formation (Hiller and Stavrakis, 1984). The outline of the basin may have almost coincided with its present limits apart from an extension towards the northwest as shown by Beaufort Group xenoliths in a kimberlite pipe from the Finsch Mine, 140 km west-northwest of Kimberley (Visser, 1972).

Post-Beaufort floodplains and deserts A Middle Triassic hiatus which increases northwards (Turner, 1983) was followed by a Late Triassic cycle consisting of a northward-thinning bedload-dominated fluvial wedge (Molteno Formation). This cycle is linked by H~ilbich (1983) with intensification of Cape Fold Belt tectonism dated at 229 4- 5 Ma, but a closer provenance than for the previous cycle is implied by the presence of Cape Supergroup quartzite pebbles and boulders. Turner (1983) attributed northward-tapering wedges in the Molteno Formation to phases of fault-controlled uplift of provenances along the southeastern basin margin. Denudation of the provenance again led to a sourceward shift of the facies with overstepping of the bedload fluvial deposits by mixed-load fluvial and flood basin/lacustrine deposits of the Elliot Formation. These deposits are overlain by the Late Triassic-Early Jurassic Clarens Formation consisting largely of aeolian fine-grained sands derived from a western source (Beukes, 1970). A progressive increase in aridity is evident in the MoltenoElliot-Clarens depositional sequence (Visser, 1991 ).

Igneous events and Gondwana break-up The main Karoo Basin was completely filled during the Jurassic with the outpouring of at least 1400 m of basaltic lavas (Drakensberg Group). The initial break-up of the southern African component of Gondwana commenced during the Middle Jurassic with the formation of rift-associated sedimentary basins around the continental margins of southern Africa and within the Cape Fold Belt (Dingle et al., 1983). The Falkland Islands also appear to have rotated 120 degrees as they moved from their position off the southeast coast of South Africa (Fig. 32) to a site approximately 500 km southeast of Cape Town (Mitchell et al., 1986). The main breakup probably commenced during the Early Cretaceous (Larson and Ladd, 1973) with the opening up of the Atlantic Ocean, lateral movement along the Agulhas/ Falkland Fracture Zone including a further rotation of 60 degrees of the Falkland Islands microplate (Mitchell et al., 1986) and separation of East Antarctica from southern Africa (Dingle et al., 1983).

310 ECONOMIC RESOURCES

Coal The main Karoo Basin is host to the major coal resources of South Africa, though significant deposits are also present in the minor, contemporaneous basins to the north. Coal is developed in both the Permian Vryheid Formation and the mid-Triassic Molteno Formation. The extent of the economic coal deposits is illustrated in Fig. 33 together with the arbitrarily defined boundaries of the various coalfields. Coal seams throughout the basin are virtually horizontal. The only significant disturbances are those associated with dolerite sills and dykes which not only displace and replace the strata but also devolatilise the coal. Distribution and thickness of coal seams in those coalfields peripheral to the basin (for example, the Springs-Witbank coalfield) are controlled by pre-Karoo and Dwyka glaciation topographic features and, to a lesser extent, by sedimentological factors. Coal seams more distally situated were influenced to some degree by basin-floor topography and tectonic events but sedimentological criteria, such as the type of peat environment, local rates of subsidence, and timing of marine transgressions and fluvial clastic influxes exerted much stronger controls (Fig. 34). The wide range of depositional settings within which peats accumulated, combined with variations in climate and plant communities as well as Jurassic dolerite intrusions, impart to the coals significant differences in grade, type and rank. These differences have important practical implications with respect to beneficiation methods and utilisation in the metallurgical, synthetic fuels and power (steam) generation processes. In general, coals of the Karoo Basin are more variable in type and contain a much higher inertic and transitional-reactive inertic component than those of Europe and the USA (Falcon, 1986). With total recoverable Karoo Supergroup coal reserves of 55 333 Mt (in situ resources of 121 218 Mt), of which 37 625 Mt are present in the main Karoo Basin (Bredell, 1987), South Africa ranks fifth in the world. Although total saleable (beneficiated) reserves within the basin are in the order of 29,000 Mt, a large component (77%) of this is lowgrade (< 25.5 Mj/kg) bituminous coal. High-grade (noncoking) bituminous reserves (12%), which contributed the bulk of the 50 Mt of coal exported in 1992 (Tinney, 1993) are mainly the product of coal seams from which prime and middlings products are prepared. Of the approximately 174 Mt of coal produced during 1992 (Tinney, 1993), noncoking bituminous coal constituted about 94% of the total. Coking coal and anthracite together account for approximately 4% and 1.5% of saleable reserves re-

M.R. JOHNSON et al. spectively but, with the introduction of direct reduction processes and declining demand for anthracite (especially in the export market), these reserves appear to be adequate in the short to medium term. The Molteno Formation, although providing most of South Africa's coal between 1900 and 1904, is no longer productive. An extensive investigation into the feasibility of exploiting a torbanite deposit (developed within the No. 5 seam in the Highveld coalfield) with a view to establishing a retorting plant to produce shale oil was undertaken recently. Oil and gas

Numerous oil shows are known in the northern part of the Karoo Basin but only two small uneconomic accumulations have been found. It was established early in the exploration for oil during the 1960s and early 1970s that only the rocks north of latitude 28~ were still in a diagenetic stage consistent with the generation and preservation of oil, except where metamorphosed by dolerite intrusions (Rowsell and De Swardt, 1976). The source rock potential of the Pietermaritzburg Formation was rated as fairly good near its top, but otherwise poor, that of the Vryheid Formation as fair to good and that of the Volksrust as mostly poor and lignitic. Because the best oil shows were encountered in the upper part of the Vryheid, this formation was probably the source. The volume of shale in the Vryheid is, however, too insignificant to be an important source. The primary porosity and permeability of the Vryheid are in general poor, although leaching has improved the quality considerably in places towards the north. Widespread intrusion by dolerites probably led to large-scale conversion of oil into gas and some escape along fractures. The Ecca shales qualify as fairly good gas source rocks but only two small uneconomic gas fields have been discovered during exploration for other mineral deposits in the northern part of the Karoo Basin. Uranium and molybdenum

Uranium occurrences, which are presently subeconomic, are located in the western and central parts of the Karoo Basin within the Adelaide Subgroup and Molteno and Elliot Formations (Cole and Labuschagne, 1985; Le Roux and Toens, 1986). The occurrences are epigenetic, tabular and sandstone-hosted, forming discrete pods and lenses less than 10,000 m 3 in volume. The sandstones represent fluvial channel deposits and are interbedded with mudrock of flood basin and/or lacustrine origin. The thickest sandstone bodies (up to 60 m thick) contain the highest proportion of mineralisation. In the Adelaide Subgroup the sandstone bodies cluster

THE F O R E L A N D KAROO BASIN, SOUTH AFRICA I

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into packages (members), with the Moordenaars and Poortjie Members, in the middle of the subgroup, containing approximately 50% of the uranium occurrences.

The dominant uranium is commonly associated senopyrite, chalcopyrite, uraninite may also occur.

mineral is coffinite, which with calcite, pyrite, arbornite and chalcocite; In the Adelaide Subgroup

312

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in the southwestern part of the basin molybdenum, in the form of molybdenite and jordisite, is present (Cole and Wipplinger, 1991). Here, a few occurrences exceed one million tons of ore but the possibility of exploiting the uranium and recovering molybdenum as a by-product is mitigated against by the marginal grades, which average less than

1500 ppm U and 800 ppm Mo. Further negative factors are the presence of calcite and clay minerals, which would cause the reagent consumption to be high in an acid leach, and the presence of sulphides, which would cause a high reagent consumption in an alkaline leach (Le Roux and Toens, 1986).

T H E F O R E L A N D K A R O O BASIN, S O U T H A F R I C A REFERENCES

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M.R. J O H N S O N et al. Van Vuuren, C.J., 1972. Geological well completion report of the Swartberg (SW1/67) borehole. Rep. Southern Oil Explor. Corp. (unpublished). Van Vuuren, C.J., 1981. Depositional models for the Vryheid Formation in the northeastern part of the Karoo Basin - - a review. Ann. Geol. Surv. S. Afr., 15: 1-11. Van Vuuren, C.J., 1983. A basin analysis of the northern facies of the Ecca Group. Ph.D. Thesis, Univ. Orange Free State, Bloemfontein, 249 pp. (unpublished). Van Vuuren, C.J. and Cole, D.I., 1979. The stratigraphy and depositional environments of the Ecca Group in the northern part of the Karoo basin. In: A.M. Anderson and W.J. van Biljon (Editors), Some Sedimentary Basins and Associated Ore Deposits of South Africa. Spec. Publ. Geol. Soc. S. Afr., 6:103-111. Veevers, J.J. and Powell, C.M., 1987. Late Paleozoic glacial episodes in Gondwanaland reflected in transgressive-regressive depositional sequences in Euramerica. Bull. Geol. Soc. Am., 98: 475-487. Viljoen, J.H.A., 1987. Subaqueous fallout tuffs of the Ecca Group in the southern Cape Province. In: G. Brown and V.A. Preston (Compilers), Workshop on Pyroclastic Volcanism and Associated Deposits. Dept. Geol. Miner., Univ. Natal, Pietermaritzburg, pp. 45-48. Viljoen, J.H.A., 1992a. Lithostratigraphy of the Collingham Formation (Ecca Group), including the Zoute Kloof, Buffels River and Wilgehout River Members and the Matjiesfontein Chert Bed. Lithostrat. Ser. S. Afr. Comm. Strat., 22. Viljoen, J.H.A., 1992b. Lithostratigraphy of the Laingsburg Formation (Ecca Group). Lithostrat. Ser. S. Afr. Comm. Strat., 20. Viljoen, J.H.A. and Wickens, H. de V., 1992. Lithostratigraphy of the Vischkuil Formation (Ecca Group). Lithostrat. Ser. S. Afr. Comm. Strat., 19. Visser, J.N.J., 1972. Sediment6re insluitsels van Karoo-ouderdom in kimberliet van die Finsch-diamantmyn. Tydskr. Natuurwet., 12: 32-36. Visser, J.N.J., 1982. Upper Carboniferous glacial sedimentation in the Karoo Basin near Prieska, South Africa. Palaeogeogr., Palaeoclim., Palaeoecol., 38: 63-92. Visser, J.N.J., 1983. Glacial-marine sedimentation in the Late Paleozoic Karoo Basin, Southern Africa. In: B.F. Molnia (Editor), Glacial-Marine Sedimentation. Plenum Publishing Corporation, New York, pp. 667-701. Visser, J.N.J., 1986. Lateral lithofacies relationships in the glacigene Dwyka Formation in the western and central parts of the Karoo Basin. Trans. Geol. Soc. S. Afr., 89: 373-383. Visser, J.N.J., 1987. The palaeogeography of part of southwestern Gondwana during the Permo-Carboniferous glaciation: Palaeogeogr., Palaeoclimatol., Palaeoecol., 61:205-219. Visser, J.N.J., 1989. The Permo-Carboniferous Dwyka Formation of Southern Africa: deposition by a predominantly subpolar marine ice sheet. Palaeogeogr., Palaeoclimatol., Palaeoecol., 70: 377-391. Visser, J.N.J., 1991. Geography and climatology of the Late Carboniferous to Jurassic Karoo Basin in south-western Gondwana. Ann. S. Afr. Mus., 99(12): 415-431. Visser, J.N.J., 1992. Deposition of the Early to Late Permian Whitehill Formation during a sea-level highstand in a juvenile foreland basin. S. Afr. J. Geol., 95:181-193. Visser, J.N.J. and Botha, B.J.V., 1980. Meander channel, pointbar, crevasse splay and aeolian deposits from the Elliot Formation in Barkly Pass, northeastern Cape. Trans. Geol. Soc. S. Afr., 83: 55-62. Visser, J.N.J. and Loock, J.C., 1978. Water depth in the main Karoo basin, South Africa, during Ecca (Permian) sedimentation. Trans. Geol. Soc. S. Afr., 81: 185-191. Visser, J.N.J., Loock, J.C. and Jordaan, M.J., 1980. Permian

T H E F O R E L A N D K A R O O BASIN, S O U T H A F R I C A deltaic sedimentation in the western half of the Karoo Basin. Trans. Geol. Soc. S. Afr., 83: 415-424. Visser, J.N.J., Loock, J.C. and Colliston, W.P., 1987. Subaqueous outwash fan and esker sandstones in the Permo-Carboniferous Dwyka Formation of South Africa. J. Sed. Petrol., 57: 467478. Von Brunn, V., 1981. Sedimentary facies related to Late Paleozoic (Dwyka) deglaciation in the eastern Karroo Basin, South Africa. In: M.M. Creswell and P. Vella (Editors), Gondwana Five. A.A. Balkema, Rotterdam, pp. 117-123. Vos, R.E. and Hobday, D.K., 1977. Storm beach deposits in the late Palaeozoic Ecca Group of South Africa: Sed. Geol., 19: 217-232. Walker, R.G., 1979. Turbidites and associated coarse clastic deposits. In: R.G. Walker (Editor), Facies Models. Geol. Ass. Canada, Waterloo, pp. 91-107. Wickens, H. de V., 1984. Die stratigrafie en sedimentologie van die Groep Ecca wes van Sutherland. M.Sc. Thesis, Univ. Port Elizabeth, 86 pp. (unpublished). Wickens, H. de V., 1985. Sedimentologiese ondersoek van die

317 Formasies Vischkuil en Laingsburg, Groep Ecca, Laingsburgomgewing. Rep. Geol. Surv. S. Afr. (unpublished). Winter, H. de la R. and Venter, J.J., 1970. Lithostratigraphic correlation of recent deep boreholesin the Karoo-Cape sequence. In: Second Gondwana Symposium: Proceedings and Papers. Counc. Sci. Ind. Res., Pretoria, pp. 395-408. Winter, M.E, 1985. Lower Permian palaeoenvironments of the northern Highveld Coalfield and their relationship to the characteristics of coal seams. Ph.D. Thesis, Univ. Witwatersrand, Johannesburg, 261 pp. (unpublished). Wright, L.D., 1978. River deltas. In: R.A. Davis (Editor), Coastal Sedimentary Environments. Springer-Verlag, New York, pp. 5-68. Zawada, P.K., 1987. The stratigraphy and sedimentology of the Ecca and Beaufort Groups in the Fauresmith area. M.Sc. Thesis, Univ. Witwatersrand, Johannesburg, 192 pp. (unpublished). Zawada, P.K., 1988. The stratigraphy and sedimentology of the Ecca and Beaufort Groups in the Fauresmith area, south-western Orange Free State. Bull. Geol. Soc. S. Afr., 90:48 pp.

C h a p t e r 13

Late Mesozoic Sedimentary Basins Off the South Coast of South Africa

I.K. McMILLAN, G.J. BRINK, D.S. BROAD and J.J. M A I E R

Late Mesozoic sedimentation off the south coast of South Africa records a history of initial continental rifting (?Middle-Late Jurassic to latest Valanginian), followed by a transitional episode (latest Valanginian to Early Aptian) and a drifting episode (Early Aptian to present day), as Africa separated from South America. Rifting appears to have been initiated by separation of East and West Gondwana during the Middle to Late Jurassic. Sediments associated with rifting are now confined to four major basins Bredasdorp, Pletmos, Gamtoos and Algoa, which are underlain and bounded by rocks of the Ordovician-Devonian Cape Supergroup that form prominent arches between the basins. During the rifting phase when half-graben basin styles were typical, sediments accumulated in a wide range of environments (non-marine to slope). In the Bredasdorp Basin, where major bounding faults are less well developed, sediments were laid down in non-marine and marginal marine environments resulting in widespread development of red and green claystones overlain by clean, porous glauconitic littoral sandstones. In contrast, transitional early drift sedimentation, which began after a major regional unconformity (seismic horizon 1Atl) in the latest Valanginian, is characterised by deep-marine, poorly oxygenated conditions. The pre-lA, 13A and 14A sequences are of considerable economic significance for hydrocarbons, particularly in the Bredasdorp Basin where commercial gas and condensate production began in 1992. Late drift sedimentation since Late Aptian has occurred in generally well-oxygenated environments, and has led to the steady southwards development of the continental shelf and the formation of an elongate basin parallel to the relict shelf break. This basin, the Outeniqua Basin, is composed of essentially mid-Aptian to Maastrichtian deposits, and overlies the pre-existing rift basins with a transverse structural grain.

INTRODUCTION

Regional setting The locations of the offshore basins, and boreholes drilled to date (January 1992), on the southern South African continental margin are shown in Fig. 1. The basins lie at the southernmost end of the African continent, where the plate margin was sheared by fight-lateral movement along the Agulhas-Falkland Fracture Zone: in contrast elsewhere in Southern Africa the margins are extensional pull-apart in style. Four major depocentres, the Bredasdorp, Pletmos, Gamtoos and Algoa basins, with the smaller Infanta Embayment, formed at the c o m m e n c e m e n t of rifting along the southern margin of the African plate. Cross-sections from west to east across the southern South African margin illustrate the variations in structural style (Fig. 2). All five basins have now been extensively explored via acquisition and interpretation of multi-channel seismic,

deep borehole drilling, and a wide variety of service geology disciplines. The latter include petrography, core analysis, palaeontology (foraminifera, ostracods, dinoflagellates, acritarchs, pollen and spores, as well as some work on nannofossils), geochemistry and vitrinite reflectance and seismic attribute studies. A summary of the number of boreholes drilled, kilometres of seismic line shot and the area of each basin is shown in Table 1. As can be seen, greatest interest has focused on the Bredasdorp Basin. Drilling for hydrocarbons was initiated in 1967 in the late Mesozoic rocks of the onshore part of the Algoa Basin, though one hole in this area dates back to 1908. Offshore drilling c o m m e n c e d in 1968 with the exploration of the Superior High and surrounding region in central Pletmos Basin. In the early 1980s interest concentrated on the gas fairway along the northem flank of the Bredasdorp Basin, where significant gas discoveries led to the Mossgas development project which began gas and

African Basins. Sedimentary Basins of the World, 3 edited by R.C. Selley (Series Editor: K.J. Hsti), pp. 319-376. 9 1997 Elsevier Science B.V., Amsterdam. All rights reserved.

tao to

Fig. 1. Location map of South African southern offshore sedimentary basins. Numbered boreholes are referred to in the text. In part after Broad (1990) and reproduced with permission of the Geological Society of South Africa. 1,,=.I

tt" >. Z

t"

9 N 9 t") t'rl

7~ -] >.

>. o~

9 ,-4 9

9 9

>.

>.

Fig. 2. Schematic profiles across the Bredasdorp, Pletmos, Gamtoos and Algoa basins, South African offshore. to i--,t

322

I.K. McMILLAN et al.

Table 1 Areas, kilometres of seismic coverage and number of wells drilled in the southern offshore basins and their onshore extensions (as of January 1992) Basin

Area, onshore and offshore (km 2)

Seismic line, offshore (kin)

Boreholes

Bredasdorp Pletmos-lnfanta Gamtoos

18,150 21,350 5,038

46,617 23,700 4,272

8,193

5,000

132 37 2 onshore 10 offshore 22 onshore 9 offshore

Algoa

condensate production from the F-A platform in early 1992. Little exploration interest has been centred to date on the distal parts of the Outeniqua Basin (Fig. 1), primarily because of the generally excessive water depths and strong Agulhas Current. Consequently, comparatively little will be found herein on the Southern Outeniqua Basin sensu stricto: a more comprehensive review of its salient features can be found in Dingle et al. (1983). It has long been recognised that the southern offshore basins exhibit features characteristic of rift basins, and more specifically of divergent margin basins (Atlantic-type passive margin basins of Bally and Snelson, 1980) which have been modified by transform movements. Their geological history is discussed in this paper in terms of the following stages of development: rift, transitional-early drift and late drift, which follows the nomenclature used by Edwards and Santogrossi (1990). A similar terminology had been used previously for the adjacent west coast of southern Africa (Gerrard and Smith, 1982). Previous work The vast majority of work undertaken by Soekor on the southern offshore basins remains in internal reports, and what has been published is essentially a synthesis. Published compilations dealing with aspects of the southern offshore geology, stratigraphy, palaeontology and geochemistry, together with relevant onshore data from the southern Cape coast include those of Du Toit (1976, 1979), Winter (1973, 1979), Leith and Rowsell (1979), McLachlan et al. (1976), McLachlan (1977), McLachlan and McMillan (1979), De Swardt and McLachlan (1982), Light et al. (1982) and Marot and McLachlan (1982). A summary of much of this work has been made by Dingle et al. (1983), and additional comments can be found in Tankard et al. (1982). More recently, intensive seismic-stratigraphic analyses of the transitional to late drift interval

(1Atl to horizon K: latest Valanginian to about Santonian) have been undertaken for the Bredasdorp and Pletmos basins (Beamish et al., 1988 and Brink et al., 1994, respectively), and resulted in detailed subdivision of the sedimentary sequences of these two basins. Between the previously identified major unconformities 1Atl, 5Atl/6Atl, 13Atl and 15Atl (see definition of a type 1 unconformity in Van Wagoner et al., 1987), there were found to occur other smaller (higher order) seismic events, each associated with a lowstand sedimentary episode on the upper slope, and separated from each other by a highstand episode. These detailed sequence-stratigraphic studies have led to the compilation of an atlas of seismic stratigraphy which draws examples from the Bredasdorp and Pletmos basins (Brown et al., 1995). Attempts have been made therein to correlate sequences and sequence boundaries with eustatic relative sea-level curves. Seismic stratigraphic terminology used in the present article follows the definitions given by Van Wagoner et al. (1987). Although the original seismic-stratigraphic studies recognised the synchroneity of the major unconformities 1Atl, 5Atl/6Atl, 13Atl and 15Atl between the Pletmos and Bredasdorp basins based on the palaeontological age-datings above and below the breaks, the smaller unconformities were identified independently in each basin, and no correlation is implied between the two basins. Consequently, although 8At l, for example, has chronostratigraphic value in the Bredasdorp Basin, its time relationship with 8Atl of the Pletmos Basin remains to be established. Although seismic and sequence-stratigraphic identification of highstand and lowstand sedimentary packages has resulted in ever-finer stratigraphic subdivisions of the sedimentary infill in the South African offshore basins, confirmatory evidence of sea-level change has often proved ambiguous or inconclusive. Repeated changes in sea-level would be expected to result in sedimentary packages containing a wide variety of deep, medial or shallow water benthonic microfaunas at any one locality. However, from foraminiferal evidence, borehole sections tend to show either little change in depositional environment or a gently shallowing upward trend. Though the relationships of lowstand and highstand sedimentary packages and associated deposits are not in dispute, it is considered here that they reflect changes in rates of continental margin subsidence and rates of sediment input amended by seismic, tectonic and sea-floor erosion processes. The reader is referred to Brown et al. (1995) and Brink et al. (1994) where the eustatic influence on sedimentary packages is emphasised. Details of holes drilled in the southern offshore basins can be found in the American Association

LATE MESOZOIC SEDIMENTARY BASINS OFF THE SOUTH COAST OF SOUTH AFRICA of Petroleum Geologists' annual reviews for central and southern Africa. General structure and history of the southern offshore basins The Bredasdorp, Pletmos, Gamtoos and Algoa Basins have comparable histories, although responses to specific events affecting the four basins are often distinctly different in each case. The oldest datable sediments drilled to date are of Kimmeridgian age (Late Jurassic) in the Gamtoos and Algoa basins, and it is considered that rocks of equivalent age also occur in the depocentres of the other basins. Even older Mesozoic sediments may occur, but lie too deeply buffed to be of economic significance in oil exploration. These earliest deposits commenced accumulation at the initiation of rifting. In southern Africa rifting is regarded as having commenced with the preliminary fracturing of Gondwana into east (Antarctica-Australia-India) and west (South America-Africa) portions. Kimmeridgian or Portlandian sedimentary rocks are known from northern Mozambique (Silva, 1966), Bajocian-Bathonian rocks in the Majunga Basin of northwest Madagascar (Espitali6 and Sigal, 1963) and coastal Tanzania (Quennell et al., 1956), and Toarcian deposits in Somalia (Kent, 1974). These ages give an impression of a seaway opening down the length of the east coast of Africa from Early Jurassic in the north to Late Jurassic in the south (Norton and Sclater, 1979). Generation of the earliest oceanic crust between East and West Gondwana is regarded as Kimmeridgian for the Mozambique Basin (in the present day southern Mozambique Channel), as described by Srgoufin (1978), Norton and Sclater (1979) and Powell et al. (1980). More recently Martin and Hartnady (1986) dated initial separation of Antarctica from the eastern margin of the Falkland Plateau as 145-122 my (M21-M 10). A summary of the plate tectonic setting of the Mesozoic southern offshore basins of South Africa is presented by Fouch6 et al. (1992). Deep Sea Drilling Project results from holes drilled on the eastern flank of the Falkland Plateau (sites 330 and 511) revealed the oldest sediments, unconformably overlying Precambrian granitic and gneissic rocks, to be of Kimmeridgian-Portlandian age, perhaps extending as far back as Oxfordian (Jones and Plafker, 1977; Jeletzky, 1983). Considerable similarities exist between the stratigraphic sequences and lithologies intersected at sites 330 and 511 and boreholes in the deeper parts of the southern Gamtoos Basin. Martin et al. (1981) have also previously commented on the similarity in structure and stratigraphy of the Falkland Plateau and Outeniqua (sensu lato) basins, and they considered

323

that the two constituted a single feature during the Late Jurassic to Early Cretaceous rift-onset to driftonset period. Thus, for the rift period, the Bredasdorp, Pletmos-Infanta, Gamtoos and Algoa basins can be regarded as proximal tongues of the large Falkland Plateau Basin that commenced subsiding during the Kimmeridgian or slightly earlier. Dingle et al. (1983) have regarded the Middle Jurassic date of 162 -+- 7 my derived from a whole-rock K/AR determination of basalt from the Suurberg Group in the northernmost onshore Algoa Basin as indicative of the commencement of basin formation. However, although pre-Kimmeridgian rocks are to be expected at several localities in the offshore Gamtoos and Algoa basins, nowhere onshore in the southern Cape basins can marine rocks be dated older than Portlandian (Colchester Member and its equivalents), and closely associated nonmarine rocks (Enon Conglomerate) are likely to be only slightly older. The relationship of the volcanic rocks of the Suurberg Group with basin formation remains uncertain. Some difficulty exists in interpretation of the drift-onset or break-up unconformity across the southern continental margin of South Africa. The term was first used by Falvey (1974) and has recently been reviewed by Braun and Beaumont (1989) in terms of its relationship to flank uplift. Two major unconformities have been proposed in the past, though they have been somewhat confused in interpretation. McLachlan and McMillan (1979), Du Toit (1979) and De Swardt and McLachlan (1982) all comment on the substantial change in sedimentary and tectonic style seen at the unconformity associated with seismic horizon C (later renamed 1At 1), which was previously dated microfaunally as near the Hauterivian-Barremian boundary. Subsequently it has become clear that the I Atl unconformity is of latest Valanginian age and although representing only a short time break, profound changes in depositional environment occurred. In contrast, in the onshore Algoa Basin although the stratigraphic position of the 1Atl unconformity can be identified microfaunally within the Sundays River Formation it cannot be recognised seismically or with electric logs. In the Pletmos, Gamtoos and Algoa basins, and to a lesser extent in the Bredasdorp Basin, a major unconformity, associated with seismic horizon 6Atl (5Atl in the Bredasdorp), and regarded as latest Hauterivian to earliest Barremian in age, marks the end of half-graben infilling and the commencement of prograding shelf sedimentation. The northern, onshore portions of the Algoa and Gamtoos basins ceased active subsidence and sedimentation at this time. The 6Atl hiatus reflects a phase of profound erosion of the underlying Hauterivian rocks, so much so that in the southern Pletmos, Gamtoos and Algoa,

324 horizons 1Atl and 6Atl combine as a compound unconformity (6Atl = 1At l) often as a result of canyon-cutting. Even in the Bredasdorp Basin where erosion was less severe, there is evidence of canyons at this level. Aspects of this are dealt with more fully later. For the present work, the Late Valanginian 1Atl unconformity is regarded as the drift-onset or break-up unconformity and marks the change from rift to transitional-early drift sedimentation, although this is probably an oversimplification. It may be that the intense erosion and canyon cutting noted at the level of 6Atl in the southern parts of the Pletmos, Gamtoos and Algoa basins reflects their proximity to the uplifting marginal fracture ridge; less severe erosion at this time in the Bredasdorp Basin perhaps accords with the greater distance of this basin from the marginal fracture ridge (Fig. 1). As noted above, latest Valanginian and later sediments accumulated as drifting between South America and Africa was activated. In terms of the southern continental margin of South Africa, this involved the shearing of the greater Falkland PlateauOuteniqua basins along the Agulhas Fracture Zone over a distance of approximately 1300 kilometres. The subsequent history of transverse movement of the Falkland Plateau past the Agulhas Bank has been described in detail by Martin et al. (1981, 1982) and Martin and Hartnady (1986). The latter authors, relying on a revised sequence of plate tectonic reconstructions, suggest that the eastern end of the Falkland Plateau cleared the tip of the South African continental margin in the late Albian. It is not unexpected that these major transform movements had a profound effect on the structural development of southern Africa. The proximity of the Agulhas Fracture Zone to the basins discussed in this paper implies that many structural aspects of these basins may be explained in terms of strike-slip or wrench faulting, particularly in those basins closest to the fracture zone. This subject was first addressed by Du Toit (1976) who observed the progressive clockwise rotation of faults from a roughly easterly orientation in the Bredasdorp Basin to a southerly orientation in the Gamtoos and eastern Pletmos basins near the fracture zone, thus implying that movement along the shear zone was responsible for the southerly curvature of the faults. However, a recent structural study of the Gamtoos and Algoa basins by Cartwright (1989) proposes that the clockwise bending of these faults, basins and arches is more likely to be due to either "tension gash" sinusoidal pullapart movements dating from the early propagating transform, or inheritance of a trend in the underlying Cape Fold Belt, rather than late bending during continental separation (Malan et al., 1990). The Agulhas marginal fracture ridge (Scrutton and Du Plessis, 1973; Ben-Avraham et al, 1993)

I.K. McMILLAN et al. and the contiguous Agulhas Arch (De Swardt and McLachlan, 1982) are regarded as having been uplifted during the drifting episode, due to thermal expansion and phase boundary migration within the lithosphere. In turn this led to crustal thinning by erosion (Falvey, 1974), although this may have been a more complex procedure for the sheared AgulhasFalkland margin than for the pull-apart Atlantic type margins described by Falvey (1974) and Braun and Beaumont (1989). Further discussion of the marginal fracture ridge and its history is provided by Ben-Avraham et al. (1993). However, neither the Agulhas marginal fracture ridge, nor the departing Falkland Plateau appear to have substantially hindered marine incursions into the basins of the South African southern offshore for any length of time, either during the rift, transitional-early drift or late drift phases. The rifted basins were smothered by deep-marine prograding and aggrading sediments during the latest Valanginian to mid-Aptian. Although subsidence remained substantial in the Bredasdorp Basin until Cenomanian times, deposition thereafter resulted in the development and subsequent outbuilding of progressively more linear continental shelf and slope systems. Figure 3 represents a generalised foraminiferal correlation of the southern offshore basins from latest rift times (immediately pre-lAtl) to late drift times (15Atl), in order to provide a biostratigraphic framework of episodes of sedimentation within this critical period (considered both in a tectonic and economic sense) for the three following sections of this article: Bredasdorp, Infanta and Pletmos, and Gamtoos and Algoa basins.

BREDASDORP BASIN

Introduction

The Bredasdorp Basin is defined by basement arches aligned parallel to the structural grain of the orogenic Cape Fold Belt. The bounding Infanta Arch to the northeast and the Agulhas Arch to the southwest define a southeasterly elongate basin approximately 200 kilometres long by 80 kilometres wide and about 18 000 square kilometres in area (Table 1, Fig. 1). Minor onshore extensions occur to the west of Cape Infanta (Dingle et al., 1983, Malan and Viljoen, 1990). Economic basement (horizon D, Fig. 4) attains a maximum depth of about 7 kilometres (4 seconds of two-way time), and has been intersected only along the northern margins and on basement highs flanking both the northeastern and southwestern margins of the basin. With the exception of three wells on the north eastern margin of the basin, intersected

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LATE MESOZOIC SEDIMENTARY BASINS OFF THE SOUTH COAST OF SOUTH AFRICA basement consists of black slates of the Devonian Bokkeveld Group: only close to the Infanta Embayment have quartzites of the Ordovician-Silurian Table Mountain Group been encountered. Initial rifting was characterised by the development of horst and graben tectonics in an extensional stress regime. Active boundary faults largely controlled rift clastic deposition. Thick successions laid down in continental, paralic and shallow-marine environments are interpreted to have developed in response to major marine transgressive and regressive cycles principally induced by syndepositional normal faulting events. Termination of active rift sedimentation is marked by the 1Atl unconformity, which records significant uplift and truncation of underlying deposits along the basin margins during the late Valanginian (Figs. 3 and 5). Features interpreted as caused by inversion tectonics occur at several localities in the Bredasdorp Basin, and at several horizons, particularly in the D to 1Atl interval, the 1Atl to 6Atl interval, and at 14Atl (van der Merwe and Fouch6, 1992). Post-lAtl to 13Atl onlap-fill sequences are associated with both rapid thermal subsidence along the basin axis and episodes of re-activated faulting. They represent transitional tectonic processes subsequent to the onset of drifting. Permanent marine conditions in the central Bredasdorp Basin were established from 1Atl times onward, with the early sediments accumulating under deep marine, poorly oxygenated conditions. The regional 13Atl unconformity in the Early Aptian (Fig. 6) marks the onset of renewed and gradually accelerating subsidence, especially in the central and southern Bredasdorp Basin. Relatively thick progradational sequences are widespread, and reflect the epeirogenic nature of Late Cretaceous sedimentation from 15Atl to horizon L (end of the Cretaceous) in the more distal parts of the basin. Subsidence of the Bredasdorp Basin ended at horizon L, at the end of the Cretaceous, with sediments accumulating very slowly during the Tertiary (McLachlan and McMillan, 1979) under stable shelf conditions. Sporadic seaward tilting caused thickness differences in an otherwise very uniform Tertiary sedimentary veneer. From the Tertiary the basin was a component of the Agulhas Bank continental margin and greater Outeniqua Basin. The Agulhas Arch was no longer a positive feature from about Late Oligocene times. Consequently, Tertiary deposits accumulating over the Bredasdorp Basin region generally reflect well-oxygenated, open ocean regimes, and are composed mainly of biogenic and limey muds, often rich in glauconite. Igneous activity is evident at several stratigraphic levels as indicated by changes in seismic character, and confirmed by drilling and petrographic evidence. Minor tuff layers within the rift succession have

329

been recognised, particularly along the northeastern flank of the Bredasdorp Basin, and Early Tertiary lamprophyres have been encountered in boreholes in the southwestern and southeastern parts of the basin. Seafloor outcropping intrusions of a similar age in the northwestern Bredasdorp Basin are trachytic and have been described by Dingle and Gentle (1972). Basin evolution Rift tectonics and sedimentation (D to 1Atl) Widespread normal faults and complex horsts and grabens are seen on seismic sections across the basin (Figs. 7, 8 and 9), indicating that the basin was subjected to an extensional stress regime during the rifting phase. Maximum known throw on the Arniston Fault which defines a major half-graben in the extreme northwest of the basin, is approximately 3850 metres, which is comparable to the throw on the Superior and Pletmos faults in the Pletmos Basin. Other large half grabens are present elsewhere within the Bredasdorp Basin but are less clearly delineated except locally in the Mossel Bay gasfields area. Detailed lithostratigraphic study of rift sedimentation along the northeastern flank of the basin indicates that differential subsidence of the basement floor strongly influenced sedimentation rates from horizon D to 1Atl (?Kimmeridgian to Late Valanginian): all lithogenetic units thicken within graben and condense over horsts. For economic reasons, rift sedimentation history has been studied in greatest detail for the gas fields area (Light et al., 1982; Strauss et al., 1990; Fatti et al., 1995) but similar lithostratigraphic successions have been described elsewhere on the flanks of the Bredasdorp Basin (Broad and Turner, 1982). Throughout the rift episode, clastic supply into the basin was mainly from the north and northeast and was derived from erosion of orthoquartzites and slates of the Cape Supergroup and sandstones and shales of the Karoo Supergroup. During Late Jurassic and Early Cretaceous times the Cape mountains are presumed to have been much higher than their present maximum of about 2000 metres due to the generally high elevation of the southern African part of Gondwana (De Swardt and Bennet, 1974; Partridge and Maud, 1987). De Swardt and Rowsell (1974) considered the Ordovician-Devonian rocks of the Cape Supergroup to have been metamorphosed by burial to a depth greater than 7000 metres prior to being subjected to folding. This implies that the Dwyka, Ecca and Beaufort Groups of the Carboniferous to Triassic Karoo sequence formerly overlay the Cape Supergroup in the Cape Mountains, and that they have subsequently been removed by erosion.

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LATE MESOZOIC SEDIMENTARY BASINS OFF THE SOUTH COAST OF SOUTH AFRICA Lithogenetic units of the rift sequence which have been identified in the gas fields area are: (1) a lower fluvial interval, (2) a lower shallow marine interval, (3) an upper fluvial interval and (4) an upper shallow marine interval. The generalised distribution of facies is illustrated in Figs. 10 and 11. The generalised distribution of facies around the basin margin during deposition of the upper shallow marine unit is illustrated in Fig. 10A (Broad and Turner, 1982). The rift succession has not been intersected in the central basin area and the basinward extent of sandstone is therefore unknown, although a facies change to shale is envisaged. In more proximal localities, especially in the northern part of the basin, rift sediments consist entirely of red and green claystone lithologies with non marine sands, or, more proximally still, greenish quartz pebble beds (as in borehole 1, line A-A', Fig. 7). The favourable reservoir quality of the upper shallow marine sandstones in trapping situations accounts for the presence of the Mossel Bay gas fields along the northern flank of the Bredasdorp Basin within the gas fairway (Figs. 1, 10B). Figs. 10B and 11 (after Holmes, in Strauss et al., 1990) illustrate the areal extent and thickness of the upper shallow marine sandstone unit as discussed later. (1) The lower fluvial interval. This fluvial interval, of which only the upper 500 metres has been intersected, consists of red and minor green argillites with subordinate reddish sandstones and rare conglomerates. Inferred locally steepened palaeogradients and onlapping infill relationships in the region of elevated basement highs suggest deposition concurrent with progressive tilting and differential subsidence. Initial graben fill is envisaged to consist of local scarp-developed alluvial fans, and wide flood plains supplied by transverse fluvial systems. Extensive oxidation and abundant caliches suggest an arid climate. Borehole 3 in the southwestern part of the basin (Figs. 7 and 8) intersected 180 m of interbedded red claystones and evaporites that are interpreted as a coastal sabkha deposit. Although these features suggest a hot and dry climate, palaeolatitude evidence indicates that the southern offshore of South Africa was sited about 50~ in the Late Jurassic to Early Cretaceous (Smith and Briden, 1977). At this latitude a more temperate, cool and wet climate would be expected and hence it is likely that the evaporites accumulated in a small subbasin of restricted access which was severely influenced by coastal winds. In the northeastern gas fields area, tufts were encountered in the lower red beds sequence and indicate extra-basinal volcanic activity. To date, no clear microfaunal age dating has proved possible for the lower fluvial interval. (2) Lower shallow marine interval. On the northern flank of the basin the first marine incursion

341

occurs at horizon V (Fig. 11) which is an erosional regional unconformity with glauconitic sands overlying red beds. These glauconitic sandstones are clean, fine-grained, well sorted and bioturbated, and contain ostracods, ooliths, comminuted crinoid and oyster shell debris, with rare whole oyster shells and very rare foraminifera. They are interpreted as tidally influenced shallow-marine deposits above active wave-base. Datable elements of the foraminifera suggest a Portlandian age. Thicknesses are usually of the order of 50 metres. Local upward-coarsening sandstone beds suggest progradational strand plain/ barrier bar construction. Local thickening of the horizon V glauconitic sandstone interval is attributed to syndepositional differential subsidence associated with fault reactivation (Strauss et al., 1990). (3) Upper fluvial interval. The overlying upper red bed interval is characterised by fining-upward features, and is composed of interbedded non-glauconitic sandstones, red and green claystones and siltstones. It is regarded as having accumulated in an alluvial flood-plain setting characterised by meandering fluvial channels. No marine microfossil or macrofossil remains have been encountered. (4) Upper shallow marine interval. The upper glauconitic sandstone (Figs. 10 and 11) developed above an unconformity following a second major marine transgression into the Bredasdorp Basin. This was followed by an overall regressive phase in which cyclical progradation was dominant. The western and eastern regions of the gas field area were subjected to significantly different subsidence rates and depositional styles at this time, as demonstrated by the facies changes illustrated in Figs. 10B and 11. The uniformly thick upper shallow marine unit in the east gives way to interbedded marine and non-marine intervals in the west. Intermittent reactivation of faulting accounts for exceptionally thick, stacked cycles in synsedimentary tectonic settings which were predominantly vertical in the east and mainly tilted in the west. The reactivation of basinmargin normal faulting at this time may have been responsible for this second marine transgression. In the eastern gas fields area the upper glauconitic sandstones achieve a maximum intersected thickness of 237 metres. The sequence is composed of blocky or cyclic upward-coarsening and upward-cleaning units that are commonly cross-bedded. These sandstones are generally rich in quartz grains, poor in lithics, and variably glauconitic. Shell debris, consisting of comminuted crinoid and oyster debris, and rare ooliths, is generally much rarer than in the horizon V sandstones and is usually restricted to an interval immediately overlying the basal unconformity. Identifiable macrofaunal and microfaunal remains have proved elusive, and the age of the interval may be anywhere between Portlandian and Late

342

I.K. M c M I L L A N et al.

Fig. 10. Areal distribution of rift lithofacies in the Bredasdorp Basin and adjacent Infanta Embayment. A. Idealised and generalised distribution of lithofacies along the flanks of the basin. Rift facies have not been intersected in the deeper central basin area. The basin shape is restored by removal of the Late Cretaceous intrusive high area in the southeast (compare with Figs. 4 and 5). B. As for A but based on more detailed lithofacies studies in the gas field areas and illustrating maximum extent of transgressions and regressions. A-A': line of section illustrated in Fig. 11. (After Holmes in Strauss et al., 1990).

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344 Valanginian. A youngest possible age would suggest that this upper glauconitic sandstone interval accumulated in the Late Valanginian, and thus equates to the basal Sundays River Formation of the onshore Algoa Basin and its equivalent in the northern Pletmos Basin (Fig. 3). The marine components of progradational cycles in the western gas fields are characterised by a greater abundance of coarse sandstones and conglomerates, and carbonaceous detritus, corresponding with a disappearance of the meagre shell debris, compared with the eastern sandstones. This is interpreted to indicate a greater proximity to source, less wave and current reworking, and probably slightly lowered salinity levels. Gradation from typical foreshore sandstones, exhibiting low-angle planar cross-lamination and heavy mineral laminar concentrates, to fluvio-deltaic distributary bar-sands and interdistributary bay muds within individual cycles (Wickens, 1989) reflect a progradation of the palaeoshoreline towards the southeast. These progradational cycles are from 100 to 200 metres thick and are considered to have been caused by fault activated increases in depositional slope. This is indicated by widespread and abundant penecontemporaneous faulting, and by soft sediment deformation observed in cores. The upper shallow marine sandstones are the best reservoirs of the gas fields area and possess significant porosities and permeabilities (Light et al., 1982; Marot and McLachlan, 1982, Questiaux et al., 1985, Strauss et al., 1990, Pferdek~imper et al., 1992). Termination of active rift tectonics is indicated by the basinwide angular unconformity 1Atl which is nearly everywhere a compound unconformity, and by the termination of faults at 1Atl in the Bredasdorp Basin (Figs. 7, 8 and 9). This substantial break in sedimentation appears to have been predominantly tectonically controlled. Rift sediments were variably eroded, especially over structural highs and along the margins of the Bredasdorp Basin.

Transitional-early drift tectonics and sedimentation (1Atl to 13Atl) General remarks. Tectonic development and depositional history of only the 6Atl to 13Atl part of this Late Valanginian to Early Aptian interval has been intensively studied using seismic sequence techniques (Beamish et al., 1988, Brown et al., 1995). Major subsidence of the basin occurred in latest Valanginian (1At l) times, following which relatively uniform, slow, thermally-driven regional subsidence prevailed. 1Atl to 5Atl sedimentation is confined to several concurrent depocentres, notably in the northwest against the bounding Arniston Fault, and in the central and southern areas along

I.K. McMILLAN et al. the basin axis. Deep-marine sedimentation predominated from 1At 1 to 13At 1; poor circulation in the overlying water column led to profoundly lowered oxygen levels at the sea-floor, and benthonic faunas (ostracods and foraminifera) are correspondingly relatively rare, and confined to proximal boreholes close to the Arniston Fault and on the flanks of the Agulhas and Infanta arches (Fig. 3). In the central basin area, benthonic faunas are lacking, and faunal composition frequently consists only of Radiolaria with shell debris (particularly Inoceramus prisms and echinoderm debris) transported from the shelf. Argillaceous marine sequences exhibiting onlapfill geometries were repeatedly eroded in proximal areas due to slower rates of subsidence, and the material carried by turbidity flows into deep water.

1Atl to 5Atl (Late Valanginian to Hauterivian). During this period the northern flank of the Bredasdorp Basin, including the gas fields area, either lacked active sedimentation or suffered erosion. Southerly trending submarine valleys and canyons were cut into pre-1Atl sediments and provided conduits for sediment passing into deeper parts of the basin. Distal deposits are mainly argillaceous, accumulated in poorly oxygenated conditions and locally exhibit source rock potential (Davies, 1990). Rare turbidite sandstones occur, as detailed by Hodges and Winters (1990). The channels and canyons effectively subdivide the pre-lAtl upper shallow marine sandstones into discrete areas and provide part of the trapping mechanism for the gas reservoirs. 5Atl to 13Atl (Barremian to Early Aptian). At 5Atl time several changes in the area of active sedimentation occurred. The bounding Arniston Fault in the north ceased effective movement, the depocentre to the south of the fault merged with the depocentre in the central and southern Bredasdorp Basin, and, most significantly, a major shoreward advance of sedimentation occurred along the entire northeastern flank of the basin, so that deposition consequently occurred over the entire region of the gas fields. Concurrently, marine sedimentation commenced in the Infanta Embayment (see later). Channelised and mounded structures are recognisable in the central Bredasdorp Basin at 5Atl times, and some are of potential economic significance. Sedimentation during 5Atl to 13Atl times was again dominated by turbidity flows into a poorly circulating and poorly oxygenated, deep marine basin. Progradation from the northern margin and from the Infanta Arch is reflected by distinctly shallowing-upward sequences. In proximal areas to the north and cast (including the gas field area) these sequences culminate in shallow marine and shelf

LATE MESOZOIC SEDIMENTARY BASINS OFF THE SOUTH COAST OF SOUTH AFRICA sands that are often clean and highly porous, and constitute the shallowest parts of the highstand systems tracts. Locally matrix-supported conglomerates occur (Hodges and Winters, 1990: Facies I). An episode of channelling occurred during the Early Barremian (6At l) and three major channel systems were cut, originating on the northwest, west and southwestern flanks of the basin, and extending towards the basin deep, which at that time was located about 25 kilometres due south of the gas fields area. The main channel, which is easterly-trending and originates from the Agulhas Arch near borehole 5, can be traced for 60 kilometres (Brink and Wickens, 1990). It lies slightly to the north of the later and much larger 13A canyon system illustrated in Fig. 12, thus demonstrating the southerly shift of the basin axis during the Early Cretaceous. Channels on the 6Atl surface are also recognised throughout the gas fields area. These southerly-trending submarine valleys and canyons were cut deeply into pre-lAtl rocks and provided conduits for sediment passing into deeper parts of the basin. These canyons cross-cut the pre-lAtl west-east structural grain and disrupt the continuity of the upper shallow marine reservoir sandstones. The canyons are clay-plugged and provide part of the trapping mechanism for the gas fields. Towards 13At 1 times (Early Aptian) subsidence rates and faulting show marked declines, presaging a more stable Bredasdorp Basin. A major change to a northerly sediment input direction can be seen at this time: sandstones became widespread over almost the entire basin in highstand tracts proximally, and in lowstand turbidites distally. Lithologies are described by Hodges and Winters (1990: their Facies III), and petrography details are provided by Hill (1990). Turbidite systems intersected in boreholes comprise slightly carbonaceous, fine-grained sandstones in lobe, depositional channel and abandonment environments, much as described by Mutti and Normark (1987). Organic enrichment of the 5Atl to 13Atl interval is intermittent. Over the northeastern flank of the Bredasdorp Basin, some enrichment is seen in the 5A sequence and locally in the 6A sequence whereas in the central Bredasdorp Basin the 9Atl to 12Atl interval also shows some organic enrichment (Davies, 1990). The pattern of distribution is not yet fully understood.

Late drift tectonics and sedimentation (13Atl to present day) The 13Atl unconformity in the Early Aptian ushered in a very different sedimentation regime. Although there is seismic evidence for major erosion around the basin margin, erosion deeper in the basin is confined primarily to cutting of the 13Atl canyon which is a submarine channel about 5 km wide and

345

approximately 50 km long, that trends in an easterly direction across the central part of the basin (Fig. 12). Updip tributaries of the system are clayplugged. The channel system served as a conduit supplying mass-flow deposits, predominantly thinly bedded turbidites, to the deeper parts of the basin. The 13A channel is the site of oil accumulations which are located in seismically-defined mounded sequences comprising submarine fan-channel complexes. In the overlying 14A sequence sandy basinfloor fans also occur in the central Bredasdorp basin, and again have oil-bearing reservoirs. Argillaceous slope-front fans have also been identified, located immediately beyond the relict shelf-edges of successive lowstand systems tracts. The interval overlying horizon 13Atl is distinguished by high-gamma organic rich claystones in the central and southern Bredasdorp Basin that only proximally contain benthonic fauna but almost everywhere are rich in planktonic foraminifera and Radiolaria (Fig. 3: 13Atl to Top Anoxic). Though distinctly thinner in the proximal, shelf areas, these claystones accumulated under severely lowered oxygen conditions and they extend northwards almost to the basin margin, and northeastwards across the gas fields area. Along the steeper southern flank against the Agulhas Arch, the 13A transgression did not advance as far as in the north and northeast. The anoxic interval, though tending to be sand-rich in the west, is elsewhere one of the most organically enriched and best source rock intervals yet intersected in southern offshore drilling (Davies, 1990). A maximum thickness of 200 metres is developed along the Bredasdorp Basin axis, which at that time was sited close to the southern flank of the basin. Source rocks are mature and viable and have supplied hydrocarbons (mainly oil) to adjacent 13A and 14A deep-water sandstone reservoirs. The Top Anoxic surface, as defined by foraminiferal data (Fig. 3), marks an abrupt return to well-oxygenated sea-floor conditions, reflecting a great improvement in the water mass circulation; benthonic faunas are thereafter mainly abundant and often diverse. Sedimentation rates and basin subsidence rates generally show signs of decline, especially in the gasfield areas of the northeastern shelf of the Bredasdorp Basin. To the northwest however, clastic input remained high, and high-stand shelf sandstones are common throughout the interval, attaining a maximum advance into the central basin area between horizons 14Ctl and 15Atl. These sands are thickest towards the basin axis and thin towards the northeastern and southwestern flanks. Distinctive lowstand-tract sandstones developed as deepwater fan-channel complexes at 14Atl times, and extend considerably into the distal, southern parts of the basin from the western and southern

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LATE MESOZOIC SEDIMENTARY BASINS OFF THE SOUTH COAST OF SOUTH AFRICA flanks. Their maximum thickness appears to be controlled by the position of the basin axis at 14Atl times. Some detail of their character is presented by Gilbert (1990), Hill (1990) and Hodges and Winters (1990, Facies II). These sandstones have substantial economic significance as oil reservoirs, and their full potential is still being evaluated. Maximum sedimentation within the 13A sequence is confined to a small region in the vicinity of borehole 3 (Figs. 7 and 8), that lay close to the shelf-break (outer shelf to upper slope) whereas maximum thickness of the 14A sequence lies about 37 kilometres east-southeast. The 15Atl unconformity shows clear microfaunal evidence of erosion, suggesting minor warping and uplift during the Late Cenomanian. Maximum erosion appears to have occurred distally, particularly in the easternmost part of the basin (McMillan, 1990a). An episode of sediment-starved, oxygen-starved, organically-enriched, plankton-enriched claystone is found immediately overlying 15Atl and shows some value as a source rock in the southern half of the Bredasdorp Basin, but is absent further north. However, the interval is generally too immature for it to be regarded as a viable source rock except in the distal, southernmost parts of the basin. This interval is dated as Early Turonian on the basis of planktonic and benthonic foraminifera. Overlying sediments show pronounced progradation, especially in the eastern parts of the Bredasdorp Basin. This prograding episode, of Turonian to mid-Coniacian age (15Atl to horizon K) can be recognised eastwards as far as the Pletmos Basin. Above horizon K sediments are predominantly biogenic, and clastic input to the basin was greatly reduced, though progradation and truncation can be recognised on several surfaces, especially at horizon X (late Santonian-early Campanian) and at horizon L (Maastrichtian-Palaeocene boundary). In the southeastern Bredasdorp Basin, a major domal structure (the F - F structure), as yet little explored and lying seaward of the end of seismic line B-B', appears to have formed during the latest Cretaceous, as indicated by the thin interval between horizon X and horizon L. It is one of the few major late structures in the basin, and it may have developed in association with, or immediately before, the many volcanic intrusions mapped in that part of the basin and which were intersected in the crestal well. Considerable thicknesses of horizon 15Atl to horizon L sediments occur in the vicinity of the F-X prospect described by Fouch6 (1990) in the southeastern part of the basin. Post-L sedimentation (Tertiary to present day) is exclusively composed of highstand shelf deposition of glauconitic clays, occasional sands and widespread biogenic clays. Relative uplift of the Agulhas Arch since shortly before horizon L led

347

to erosion of Late Cretaceous sediments from the flanks of the arch and redeposition of the fossiliferous debris into Palaeocene shallow-marine clayey and glauconitic sands. To the east, away from the Agulhas Arch, middle to outer shelf deposits accumulated. A major Late Oligocene unconformity can be recognised, in keeping with many other localities worldwide (Vail et al., 1977). Early Miocene biogenic clays are widely developed over the entire southern half of the basin, and their depositional environments imply that by this time the Agulhas Arch had essentially foundered. Later deposition has been minimal, and where sampled, thin veneers of one metre or less, of Holocene and latest Pleistocene sediments unconformably overly Early Miocene rocks.

PLETMOS BASIN AND INFANTA EMBAYMENT Introduction

The Pletmos Basin together with the Infanta Embayment, is the largest of the southern offshore basins, being some 270 km long and 110 km wide (Fig. 1). As with the Bredasdorp Basin, these areas are defined by arches of Palaeozoic rocks (the lnfanta and St. Francis arches), and are aligned parallel to the east-southeasterly-trending structural grain of the Cape Fold Belt. The pre-Mesozoic basement has been penetrated on basement highs only and consists mainly of Ordovician-Silurian Table Mountain Group quartzites belonging to the Palaeozoic Cape Supergroup. The Pletmos Basin is a structurally intricate basin, much more so than the Bredasdorp Basin. It is divided into two major areas by the complex Superior Fault (Fig. 1), and during the rift phase distinctly different sub-basins existed to the north and south of the fault. The major bounding faults, notably the Plettenberg Fault in the northeast, and the Pletmos Fault in the southwest, virtually define the limits of rift sedimentation (Figs. 1, 2). The northwestern parts of the basin have been little explored even though some major local depocentres have been recognised there. Drilling results indicate that depositional environments in that region are mainly proximal, with red and green non-marine beds and shallow marine claystones and sandstones predominating: source rocks seem to be lacking and gas levels while drilling have not been encouraging. A number of onshore extensions of the Pletmos Basin, notably in the vicinity of Mossel Bay, at Knysna, and at Plettenberg Bay, are present. Where sediments are exposed, lithologies comprise varicoloured claystones of non-marine origin or shallow marine claystones and sandstones, and have been

348 described by Rigassi (1970), Rigassi and Dixon (1972), McLachlan and McMillan (1976), McLachlan et al. (1976), McLachlan and McMillan (1979) and Rust (1983). Additional comments and a summary are given by Dingle et al. (1983) and Malan and Viljoen (1990). The tectonostratigraphic evolution of the Pletmos Basin has been discussed by Bate and Malan (1992). Because of its poor hydrocarbon potential the Infanta Embayment has been little drilled. The embayment, approximately 80 km by 40 km, is aligned parallel to the Cape Fold Belt structural grain and lies seawards of, and is partly enclosed by, basement highs of the Infanta Arch. The embayment extends southeastwards over a low basement ridge into the Bredasdorp Basin, and similarly northward into the Pletmos Basin. A major graben is developed eastwards down the axis of the distal Infanta Arch (Fig.

13). In both the Pletmos Basin and Infanta Embayment, rift, transitional-early drift and late drift phases of sedimentation are recognised and the basin-wide unconformities D, 1Atl and 13Atl delineate the onset of these episodes. From the top of basement (horizon D) to horizon 1Atl (Kimmeridgian to Late Valanginian), an extensional stress regime led to horst and graben tectonics and, locally, extremely thick accumulations of sediment, most notably in the graben just south of the Plettenberg Fault (Figs. 4, 5). Thick D to 1Atl intervals also occur just to the north of the Superior Fault, west of section line E-E' (Fig. 14), as well as in the southernmost Pletmos Basin, north of the Pletmos Fault. D to 1Atl sediments are composed of inner to outer shelf sandstones and claystones with localised non-marine red and green beds. Since oxygen levels at the time of deposition appear to have been near normal in both the Pletmos Basin and Infanta Embayment, organic enrichment in these rocks is very rare. Like the Bredasdorp Basin during this time period, the local relative sea level variations probably reflect an interplay of fluctuations in rates of subsidence caused by syn-depositional faulting and sediment supply. A marked change in sedimentation pattern occurs above 1Atl. Compared with the underlying sediments the Late Valanginian to Hauterivian (1Atl to 6At l) deposits accumulated in a substantially deeper marine environment, and in profoundly lowered oxygen conditions with local organic enrichment occurring over the 1At 1 surface. The 6Atl unconformity marks a phase of uplift and erosion prior to deposition of a second, mainly deep marine, poorly oxygenated sequence (6Atl to 13Atl). As with 5Atl marking the termination of the Arniston bounding fault in the Bredasdorp Basin, 6Atl marks the end of substantial normal movement of the Pletmos,

I.K. McMILLAN et al. Plettenberg and Superior faults, though subsequent strike-slip reactivation can be seen along the Superior Fault. Concurrent with the 6Atl phase of uplift, sedimentation rates in the Pletmos Basin declined substantially north of the Superior Fault. Only the southernmost Pletmos Basin shows thicknesses comparable to the Bredasdorp Basin from 6Atl through to horizon L at the top of the Cretaceous. The Tertiary sediments accumulated exclusively on the continental shelf where they form a relatively thin veneer thickening gradually in a southerly, distal direction. Basin evolution Rift tectonics and sedimentation (D to l A t l ) Pletmos Basin. Initial rifting was developed by complex horsts and grabens, with the rift sediments of widely variable lithologies being deposited mainly in the graben and reflecting rapid changes in depositional environment. The horizon D surface (economic basement) is shown in Fig. 4. The rapid changes in facies in the earliest sediments, together with a lack of complete sequences drilled in the basin deeps has hindered a full understanding of the early history of the Pletmos Basin and the Infanta Embayment. The interval is too deeply buried in much of the Pletmos Basin for it to be of economic significance. The D to I Atl interval thicknesses intersected in boreholes range from 2500 metres on the southern margin of the Plettenberg Graben to as little as 500 metres on bevelled highs. Sedimentation probably commenced in Kimmeridgian times, especially in the deep Plettenberg Graben in the northeast of the basin (Fig. 15). Here the basement surface probably had a gentle northeasterly dip into the Plettenberg Fault. However, over the Superior High (line E-E', Fig. 14a), on the eastern flanks of the southern parts of the Infanta Arch, and on the northern margins of the basin off Plettenberg Bay, the oldest sediments are Portlandian. On the basement highs such as the Superior High, the earliest sediments are often stained red and comprise a veneer of pebbly and sandy beds. These are overlain by shallow marine claystones and sandstones which contain benthonic microfaunas that suggest a slightly restricted environment. Foraminiferal correlation with other basins is not easy, but the assemblages found in the lower glauconitic sandstone at horizon V in the Bredasdorp Basin (Fig. 11) show some elements in common with those seen in the Portlandian "Colchester" equivalent on the northern margin of the Pletmos Basin (McLachlan et al., 1976). The deep marine, poorly oxygenated environments present in the Gamtoos and Algoa Basins during the Kimmeridgian and Portlandian have not been recognised in the Pletmos and Bredasdorp

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LATE MESOZOIC SEDIMENTARY BASINS OFF THE SOUTH COAST OF SOUTH AFRICA basins. The deepest marine environments during this time are restricted to that part of the basin north of the Superior Fault, where they contain diverse microfaunas that appear no deeper than outer shelf. To the south of the Superior Fault environments were shallower, ranging from inner shelf to transitional, and the resulting variably glauconitic and shelly sandstones are well-developed. However, microfaunas are rare. Slight deepening occurred in the later Portlandian when marine sedimentation became more argillaceous. Similar depositional conditions appear to have been maintained during the Berriasian, though sandstones are less well developed probably due to a shift in the point source. On the northern rim of the basin during this period, minor red beds with shallow-marine sandstones (often very shelly with crinoid skeletal debris) indicate a localised regression. The red beds occupy the same time-span as the Kirkwood Formation red beds of the northern Algoa Basin. However, although conditions are marginally marine throughout the interval south of the Superior Fault, no equivalent regression can be identified there. In the Early Valanginian, depositional environments almost everywhere show signs of shallowing upward, with shelly and glauconitic sandstones being widespread. Most of these sandstones are marginal marine and tidally-influenced in character. Only in the Plettenberg Graben to the northeast did middle to outer shelf claystone and sandstone sedimentation continue. In the Pletmos Basin sandstones were best developed during late Early to Late Valanginian. Permeabilities, porosities and thicknesses of sandstone units attain a maximum in this interval, and constitute economically significant reservoirs in the Superior High area (Fig. 16) (Maier, 1990). These sandstones are generally fine grained and glauconitic with varying amounts of argillaceous matrix, and are interbedded with glauconitic and locally calcareous claystones and siltstones. Sandstone distribution is indicated on the log correlation of the Pletmos boreholes (Fig. 16). Tectonically the basin is relatively complex. Large normal faults developed during the early rift period, and movement on the hanging wall contemporaneous with sedimentation is reflected in section E-E' and F-F' (Figs. 14 and 15). Maximum known throw of basement on the Pletmos, Superior and Plettenberg faults is approximately 2600 metres, 5000 metres and 5600 metres, respectively. The irregularity of the main fault lineaments reflects stresses originating from later strike slip movement along the Agulhas Fracture Zone, to be discussed later. The resulting fault patterns created potential structural traps for hydrocarbons in the D to 1Atl interval.

355

Infanta Embayment. Pre- 1At 1 sedimentation in the Infanta Embayment is comparable to that seen in the Pletmos Basin southeast of the Superior Fault, and in the Bredasdorp Basin. Initial sediments in the embayment are non-marine, with red beds dominant, and these have been interpreted to represent fluvial meandering-channel sandstones with overbank siltstones and claystones. Sedimentation probably began during the Kimmeridgian but no datable microfossils have yet been found. Deposition was initiated in two depocentres, one being the Infanta Embayment proper, the other being the graben along the length of the distal Infanta Arch. A summary of the evolution of the Infanta Embayment during the Late Jurassic to Early Cretaceous period (D to 1At 1) is presented by Turner (1990). Glauconitic sandstones and interbedded claystones were laid down during a Portlandian marine episode and overlie an unconformity mapped as seismic horizon IV (Fig. 13). The sandstones contain comminuted crinoid and oyster debris, ooliths and ostracods, and rare foraminifera, and are interpreted to have accumulated in a shoreline system. Though palaeontological data is not entirely confirmatory, it seems likely that horizon IV of the Infanta Embayment and the graben is equivalent to horizon V of the northeastern Bredasdorp Basin. In the vicinity of borehole 7 (section D-D', Fig. 13) in the graben, and west of the distal Infanta Arch, the basal deposits overlying the horizon IV unconformity consist of oolitic and shelly grey limestones and claystones, although the majority of the marine sequence is as described above. In general terms the thickness of marine beds is much greater here than in the Bredasdorp Basin, and is more in keeping with the Pletmos Basin. Following a regressive episode, during which thin but widespread non-marine red and minor green beds were deposited, a further marine transgression occurred, and variably shelly glauconitic sandstones and interbedded claystones were again deposited. Reliable micropalaeontological data from these upper marine beds is lacking: their age may be as young as Late Valanginian, and thus they could be equivalent to the lower beds of the Sundays River Formation in the Algoa Basin, and to the upper shallow marine gas-bearing sandstones of the northern flank of the Bredasdorp Basin. Transitional-early drift tectonics and sedimentation (1Atl to 13Atl) 1Atl to 6Atl interval. The rifling episode was terminated at 1Atl times by major tectonism in the Late Valanginian. In common with the Bredasdorp Basin, 1Atl to 6Atl sediments, of latest Valanginian to Hauterivian age, were deposited in markedly deeper water than those immediately underlying

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LATE MESOZOIC SEDIMENTARY BASINS OFF THE SOUTH COAST OF SOUTH AFRICA horizon 1At1, implying abrupt subsidence of the Pletmos Basin. In addition, oxygen levels at the sea-floor were much reduced, leading to widespread accumulation of organic-rich claystones above horizon IAtl. These claystones have some potential as source rocks locally in the Plettenberg Graben and at isolated sites southwest of the Superior Fault. They contain a distinctive foraminifera fauna composed mainly of a variety of benthonic agglutinated foraminifera, notably Trochamminoides sp. A, with floods of dictyomitroid and spherical Radiolaria. The overlying beds consisting of cyclic deep marine claystones with local channel sandstones, are poorly fossiliferous with much land-derived plant debris and few Radiolaria. The overall 1Atl to 6Atl succession is regressive and becomes sandier up section. The 1Atl to 6Atl sediments are bounded in the northeast by the Plettenberg Fault, and in the southwest by the Pletmos Fault. Along the northern margin of the Pletmos Basin are several boreholes which indicate that the 1A to 5A sequences were deposited in an environment shallower than much of the basin, probably on the shelf. Indications are that sea-floor oxygen levels were nearer to normal, with the benthonic foraminifera faunas correspondingly more diverse and with calcareous forms dominant. Here, sequences consist of grey claystones with thin argillaceous sandstones. Locally, thicker channel and bar facies occur, composed of fine- to medium-grained, glauconitic and porous sandstones, that were sourced from the north and northeast of the Pletmos Basin. The cyclicity of the 1A to 6A sequences is regarded as having developed in response to third order eustatic sea-level changes (Brink et al., 1994); however, movement on the major bounding fault planes may well have influenced cyclical sedimentation. Adjacent to the Plettenberg Fault sedimentation rates were high greater than 200 metres per million years ~ and subsidence rates were also correspondingly high. Borehole thicknesses of the 1Atl to 6Atl interval attain 2800 metres in the Plettenberg Graben, but only 50 to 100 metres distally in the southernmost Pletmos Basin. Substantial uplift at 6Atl times led to severe erosion of the 1A to 5A sequences as well as the upper pre-lAtl section in the central and southern Pletmos. The most complete pre-6Atl sequences are thus found along the northern margin of the Pletmos Basin, such as that intersected by borehole PB-A1 (McLachlan et al., 1976), and elsewhere close to the Plettenberg Fault. Southward, sequences are increasingly abbreviated, as in boreholes 12 and 13 (Fig. 16), and south of the Superior Fault only the basal portion (1Atl to 2Atl or 3Atl?) is preserved in basin deeps. The 1A to 5A sequences have not been recognised in the Infanta Embayment, and it is

357

unclear whether they were removed by 6At 1 erosion, or whether they were never deposited. 6Atl to 13Atl interval. Following major uplift and erosion at 6Atl time, sedimentation reflects marked changes in clastic supply and basin shape. Horizon 6Atl marks the end of substantial movement on the bounding faults (especially the Pletmos and Plettenberg faults). Again, sedimentation during this interval was cyclic, and shows substantial aggradation and progradation. High frequency sequences and sequence sets have been interpreted as responses to the interplay between lower order eustatic fall and rise cycles, high sediment input rates and a period of relatively stable subsidence rates (Brink et al., 1994). Sedimentation in the Pletmos Basin initially occurred simultaneously in two major sub-basins, one north of the Superior High, and one to the south. During 6A deposition, sediment sources migrated from the north to the northwest margin of the basin, and progradation advanced down the axes of the two sub-basins. In time the Superior High was inundated: in general the northern sub-basin was infilled by proximal, sandier sequences, while the southern sub-basin was infilled by distal clays. Gouws (1990) has described a lowstand prograding wedge in the 8E sequence which was explored for hydrocarbons. Distal sedimentation occurred in poorly oxygenated, deep marine conditions on the 6Atl surface and led to the accumulation of organic-rich claystones having source potential. The interval is characterised by distinctive agglutinated benthonic foraminifera (see Fig. 3) with occasional floods of planktonic foraminifera, and more widespread floods of Radiolaria. The overlying prograding beds in the 6Atl to 13Atl interval are poorly fossiliferous, often containing restricted foraminifera faunas of low diversity, with scattered Radiolaria. These sediments display progressive shallowing upwards features and show a tendency to become sandier, with lignite and localised coal beds indicative of marginal marine conditions near 13Atl. The Superior High tended to cause sediment starvation in the southern Pletmos Basin, so that the Barremian to mid-Aptian sequences become very thin distally, as in the vicinity of borehole 9 (Figs. 14 and 16). Although high porosities sometimes occur in the proximal sandstones, source rocks are too distant for these sandstones to be of economic significance. The 6A to 12A sequences in the Infanta Embayment show close similarities to those on the northeastern flank of the Bredasdorp Basin, but are distinctly more proximal in character. In a borehole located close to the depocentre, uppermost slope grey claystones with some organic enrichment, and variable Radiolaria, grade upward to thick, shallow marine, bar sandstones which are glauconitic, fine

358 to medium grained and well sorted. The maximum basinward advance of shelf sandstones in all three basins, Bredasdorp, Pletmos and Infanta, appears to have been roughly coeval during 10Atl to 13Atl times (Early Aptian) (Fig. 3). Thicknesses of the 6At 1 to 13At 1 interval penetrated in boreholes attain a maximum of 2000 metres, with the interval thinning northward to the basin margin and southward into deep water, where they are of the order of 100 metres.

Late drift tectonics and sedimentation (13Atl to present day) After the mid-Aptian, sedimentation rates show a marked decline over the entire Pletmos Basin and Infanta Embayment, and sandstones become comparatively rare. Thereafter, the basin subsided slowly within a widening continental shelf setting. The Superior High and the distal Infanta Arch have little or no influence on sedimentation patterns from 13Atl times onward. Faunal associations and sediment thicknesses become very uniform from the eastern Bredasdorp Basin across to the Pletmos Basin. Only in the southernmost Pletmos Basin have higher sedimentation and subsidence rates been encountered. The Early Aptian to mid-Cenomanian interval, 13Atl to 15Atl, is characterised by an organicenriched episode at its base (13Atl to Top Anoxic, see Fig. 3) that is distinguished by floods of planktonic foraminifera and Radiolaria and a lack of benthonic species. This claystone has some potential as a source rock. Though widespread on the upper slope in the southern Pletmos Basin, it is everywhere considerably thinner than its equivalent in the Bredasdorp Basin (where it is a major source rock), and it cannot be clearly recognised on the shelf in the northern Pletmos Basin. The Top Anoxic to 15Atl interval consists of well-oxygenated shelf and slope claystones with minor sandstones: only on the inner shelf overlying the old Plettenberg Graben are sandier sequences found, with shelly and glauconitic sandstones common. Diverse foraminiferal faunas are widespread, and are correlatable with Bredasdorp Basin sequences. These correlations suggest that over much of the Pletmos Basin the majority of the Cenomanian section has been eroded away following mild tectonic uplift at 15Atl times (Late Cenomanian), as described by McMillan (1990a). Thicknesses of the 13Atl to 15Atl interval intersected in boreholes are never greater than 300 metres on the shelf but thicken to 700 metres or so on the upper slope. Overlying the 15Atl surface is an Early Turonian interval of oxygen-starved, organically enriched sediments that may prove to be of some value as a source interval in the south, although it is generally regarded as thermally immature. Thereafter follows

I.K. McMILLAN et al. a phase of substantial progradation which culminated on the shelf with a latest Turonian fall of sea level and accumulation of fine to medium grained, clean, glauconitic, shelly sheet sands (horizon 11 sand). In deeper water, progradation continued until horizon K (top middle Coniacian). Sediments consist of clays with local channel-derived argillaceous, fine-grained sandstones. From mid-Coniacian to mid-Campanian (horizon K to horizon X) sedimentation rates in both the Pletmos Basin and the Infanta Embayment show a marked decline, corresponding to an increase in the biogenic component and a predominance of orientated lime rich claystones. Thereafter, later Cretaceous sediments are glauconitic claystones with occasional sandy stringers that indicate a reactivation of southerly-orientated progradation. A mid-Campanian transgression at horizon X is inferred along the northern margin of the Pletmos Basin. This phase of increased subsidence and faster sedimentation was brought to an abrupt termination by tectonic activity at horizon L, at the end of the Maastrichtian. Sedimentation thereafter was extremely condensed. Deposition commenced early in the Palaeocene, and consists of glauconitic claystones. Biogenic limey clays predominate in the Middle and Late Eocene, with greenish siltier glauconitic clays in the Early Oligocene. Local channel and minor bar sandstones can be recognised. Unconformably overlying the Early Oligocene are Early Miocene white biogenic limey clays topped by a substantial hardground, in turn overlain by an extremely thin veneer of latest Pleistocene and Holocene. All over the Pletmos Basin and Infanta Embayment explored to date, Cainozoic rocks accumulated in shelf environments and they show little variation in foraminiferal faunas or depositional environments. Maximum borehole thicknesses of the 15Atl to horizon L interval are of the order of 1500 metres, and of the horizon L to sea-floor interval about 700 metres in the southernmost Pletmos Basin.

GAMTOOS AND ALGOA BASINS

Introduction The Gamtoos and Algoa Basins (Fig. 1) are half grabens bounded by major faults to the northeast. Although each basin is substantially smaller than the Bredasdorp and Pletmos Basins (see Table 1) they nevertheless contain comparable thicknesses of sediment (Fig. 2). The Gamtoos Basin is a relatively simple half graben (Figs. 17 and 18), but the Algoa Basin is subdivided into three fault troughs, the Port Elizabeth Trough, the Uitenhage Trough, and the Sundays River Trough in a west to east direction

LATE MESOZOIC SEDIMENTARY BASINS OFF THE SOUTH COAST OF SOUTH AFRICA (Figs. 1, 19 and 20). The last-named trough occurs almost entirely onshore and contains over 4200 metres of Portlandian to Hauterivian sediments. The Gamtoos Basin and the three Algoa Basin troughs are separated from each other by highs of Palaeozoic sediments in the form of arches on their western margins and faulted upthrown blocks on their eastern sides (Figs. 17, 19 and 20). These arches are composed mostly of the Palaeozoic Cape Supergroup (mainly Ordovician-Silurian Table Mountain quartzites and Devonian Bokkeveld slates) which are aligned along the grain of the Permo-Triassic Cape Fold Belt. Deep drilling offshore has intersected basement rocks (Table Mountain quartzites) only on the flanks of the basement arches and on basement highs. Onshore drilling in the Uitenhage Trough has everywhere terminated in Table Mountain quartzites, while in contrast, every hole drilled to basement in the Sundays River Trough bottomed in Bokkeveld black slates. The major bounding faults (Fig. 1), particularly the Gamtoos Fault and the St. Croix Fault, extend deep into the crust and probably have complex histories (Friedinger, 1986; Malan et al., 1990). Onshore the Gamtoos Fault has a throw of about three kilometres; offshore the fault plane can be traced to a depth of approximately 12 kilometres (5.5 seconds two-way time). Seismic profiles across the Gamtoos Fault illustrate its listric nature (Figs. 17, 18). Additional major faults in the Gamtoos Basin occur on the eastern flanks of the shallow St. Francis Arch. In the Algoa Basin the Uitenhage Fault, which divides the southernmost Uitenhage Trough into two half grabens, appears to be a late feature and was possibly developed during 1Atl to 6Atl times. The Port Elizabeth Trough was probably continuous at its northern end with the Uitenhage Trough prior to fault movement during the latest Valanginian to Hauterivian (1 At 1 to 6At 1). In both basins, subsidence and sedimentation were rapid during the pre-Kimmeridgian to Hauterivian (D to 6At 1), but later sediments (excepting the 13Atl to 15Atl canyon fills) are much thinner compared with the equivalent intervals previously described for the Pletmos and Bredasdorp Basins (Fig. 3). A simplified chronostratigraphic table is provided in Fig. 21. No detailed sequence-stratigraphic study has been undertaken for these basins because of the lack of economic success, limited thicknesses in the 1Atl to 15Atl interval and the prevalence of severe faulting in the pre-1At 1 section which makes seismic correlation difficult. Structural development of these basins has been summarised recently by Malan et al. (1990), and their hydrocarbon potential discussed by Broad (1989, 1990). Additional comments on the Algoa, Gamtoos and Pletmos basins and their

359

tectonostratigraphic evolution are given by Bate and Malan (1992). Basin evolution Rift tectonics and sedimentation (D to 1Atl) Development of the Gamtoos and Algoa Basins occurred during the Late Jurassic and earliest Cretaceous, but may well have been initiated in the Middle Jurassic. The oldest dated sediments encountered in drilling are Kimmeridgian, but substantial thicknesses near depocentres remain unexplored, particularly in the Gamtoos Basin where the depocentre contains approximately 7000 metres of undrilled Mesozoic section. Here basement horizon D exceeds 5.5 seconds two-way time (approximately 12 km) and it is possible that the undrilled rift sediments are of Middle to Late Jurassic age (see Fig. 17). Gamtoos Basin. Early sedimentation in the Gamtoos Basin shows substantial lateral variation in depositional environments. Kimmeridgian sedimentation on the flanks of the St. Francis Arch (borehole 16, Fig. 17) is characterised by a basal non-marine conglomeratic and red bed interval, overlain by shallow-marine interbedded sandstones and siltstones. In contrast, to the east, close to the Gamtoos Fault and near the Gamtoos Basin depocentre, upper slope black claystones with minor turbiditic sandstones accumulated at this time in lowered oxygen conditions. Microfaunas here are sparse, mostly agglutinated foraminifera, with Radiolaria always present and often occurring in floods. Several intervals show organic enrichment and have good potential as source rocks, even though they are buried to depths around 4000 metres. Minor gas shows occurred while drilling thin sandstones in the sequence (Broad, 1989, 1990). Portlandian sedimentation in the depocentre is a continuation of that of the Kimmeridgian interval, but shows a gentle shallowing and increase in oxygen levels upwards, corresponding to a slow increase up-section of benthonic microfaunas. On the flanks of the St. Francis Arch, Portlandian sedimentation (borehole 16) commenced with an abrupt deepening, and outer-shelf well-oxygenated claystones with diverse benthonic microfaunas were laid down over the Kimmeridgian shallow marine sandstones. During Portlandian to Berriasian times the area of sedimentation in the Gamtoos Basin appears to have enlarged considerably, both towards the St. Francis Arch (borehole 15, Fig. 17) and towards the Recife Arch (boreholes 20, 21, Fig. 18). At these localities environments of deposition deepen upward towards the top of the Berriasian. Poorly fossiliferous basal pebble beds and sandstones give way up-section to outer shelf, well oxygenated claystones

360 with a diverse microfauna. In contrast, boreholes in the central part of the basin show a steady shallowing-upward to outer shelf conditions by the top Berriasian, coupled with increasing sea-floor oxygen levels and increasingly diverse and abundant microfauna that for the first time shows an appreciable calcareous component in the benthonic foraminifera. Sediments remain predominantly grey claystones with minor sandstones. In the offshore Gamtoos Basin Valanginian sedimentation reflects trends established in the Portlandian and Berriasian. The Valanginian shows much more uniformity of depositional environment across the southern half of the basin than seen lower in the section. Diverse and abundant foraminiferal faunas indicate well-oxygenated conditions at the sea-floor, and sediments everywhere are middle to outer shelf claystones with rare sandstones. Only in the northern half of the offshore basin are more proximal depositional environments seen. Boreholes drilled in the north close to the Gamtoos Fault intersected shallow-marine (mainly innermost shelf) interbedded sandstones and claystones (borehole 20, Fig. 18), or, further northwest, transitional and estuarine sandstones with fewer claystones and occasional red bed intervals. An entirely different sedimentation regime prevailed during this period in the onshore part of the Gamtoos Basin, where over 3000 metres of redstained conglomerates grading up-section to sandstones and rare red beds with very rare shallowmarine claystones, were intersected (McLachlan and McMillan 1976). Marine foraminifera in one interval high in the sequence suggest a pre-lAtl late Valanginian age, but the majority of the interval is undated, though it may well range back to at least the Portlandian.

Port Elizabeth Trough. In the Algoa Basin the three troughs display different rift sedimentation histories (Fig. 21) and consequently they are described separately. In the Port Elizabeth Trough basement horizon D attains about 3.5 seconds two way time (approximately 6500 metres) (Figs. 4, 19). Boreholes 23 and 24 (Fig. 19) intersected rift sequences, similar to those in borehole 16 in the Gamtoos Basin. In borehole 24 a basal shallow-marine interval, regarded as ?Kimmeridgian age consists of shallow-marine to transitional claystones and sandstones, locally rich in lignite, and with very poor, marginal marine foraminifera faunas. This sequence shallows westwards to coarse, pebbly sandstones and claystones containing lignite and siderite spheres, with some red beds in borehole 23. A marked deepening of the depositional environment can be seen at the base of the Portlandian, with outer shelf conditions predominating thereafter. The Portlandian

I.K. McMILLAN et al. sediments are mainly claystones which display intervals of organic enrichment, with significant source rock potential in boreholes 23 and 24. Significant oil shows occurred in adjacent porous and permeable sandstones (Broad, 1990). Microfaunas are abundant, though the foraminifera faunas are not diverse. Through the Berriasian interval in the two boreholes a general shallowing-upward trend can be recognised from the foraminifera faunas, and lithologies consist of interbedded sandstones and claystones deposited on a well-oxygenated shelf. Sedimentation in the two more northerly Port Elizabeth Trough boreholes commenced in the mid-Berriasian: a characteristic that seems very similar to that seen on the flanks of the Gamtoos Basin, as discussed previously. The Berriasian to Early Valanginian intervals of these northerly wells consist of interbedded sandstones and claystones with foraminiferal faunas that indicate well-oxygenated inner shelf environments. Early Valanginian sediments are preserved only locally due to erosion and indicate that the sequence becomes progressively more sandstone rich, and transitional in depositional environment, towards 1At 1. Over the entire Port Elizabeth Trough the 1At 1 surface is intensively planed, so that the most complete sequences (?Late Valanginian) lie close to the Port Elizabeth Fault.

Uitenhage Trough. The Uitenhage Trough (Fig. 20) shows considerable lateral variation in lithologies through the D to 1Atl sequence. Drilling in the onshore portion of the trough, where horizon D attains a maximum depth of about 1800 metres (Fig. 4) has revealed basal non-marine red sandstones and conglomerates (Enon Formation, Fig. 21) overlain by fluvial sandstones (Swartkops Member). These are overlain by a marine and fluvial-influenced interval of greenish and grey claystones with sandstones, that contain Portlandian foraminifera (Colchester Member equivalent). This Colchester Member equivalent is overlain by non-marine red and minor green beds with sandstones (main body of the Kirkwood Formation) which is in turn overlain by shallow marine grey claystones and sandstones with a foraminifera fauna indicative of a Late Valanginian age (lower Sundays River Formation). Some localised organic enrichment occurs in the Colchester Member equivalent, but it has not been buried deeply enough for it to be a viable source rock. The sequence outcrops near Port Elizabeth and has been described by McLachlan and McMillan (1976, 1979). In 1908, one of the first boreholes drilled for oil, at Swartkops, just north of Port Elizabeth, failed to discover hydrocarbons, but located a thermal mineral water supply that supported a health spa for several years. Offshore in the Uitenhage Trough drilling has occurred only in the southern half, well

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away from the onshore portion. Basement horizon D here attains a maximum of 4 seconds two-way time (approximately 8000 metres depth) in the most distal part (Fig. 4). North of the Uitenhage Fault the sequence thickens towards the bounding St. CroixCoega Fault system, but, as seen on Fig. 20, the plane of the St. Croix Fault lies at a very low angle, and underlies perhaps half the width of the Uitenhage Trough (see Malan et al., 1990, for a discussion of the unusual multileaved fault plane geometry of the St Croix Fault). Because of this, the most complete earlier successions are developed along the western portion of the trough (borehole 27, Fig. 20), while the most complete later sections lie close to the St. Croix Fault plane. Depositional environments in the D to 1At 1 interval are unusually shallow when compared with the equivalent interval in the adjacent Port Elizabeth Trough. Kimmeridgian sedimentation close to the St. Croix Fault consists of inner shelf claystones with sandy stringers, and contains a relatively diverse benthonic foraminiferal fauna. Interpreted Kimmeridgian sediments intersected a little further west, however, as in borehole 27 (Fig. 20) contain non-marine red pebbly sandstones and conglomerates, overlain by shallow-marine lignitic claystones and marginal marine sandstones. Portlandian sediments close to the St. Croix Fault reflect hypersaline, mud fiat conditions and contain

foraminiferal faunas very similar to those seen in the Colchester Member equivalent of the onshore part of the Uitenhage Trough. Away from the St. Croix Fault (borehole 27), the Portlandian interval is partly non-marine. Sediments in this interval are either greenish grey claystones with sandy stringers, marginal marine sandstones, or non-marine red claystones and fluvial sandstones. To the west of the Uitenhage Fault, Portlandian sediments were also intersected in borehole 25: they are shallow-marine interbedded grey claystones and sandstones. Berriasian sedimentation everywhere shows a distinctly shallower aspect than seen in the underlying Portlandian and Kimmeridgian. It is characterised by non-marine red and green beds, and transitional or estuarine green claystones rich in siderite spheres, and common lignite. Foraminiferal faunas are very poor, but freshwater microfaunas and floras have been identified. Similar depositional environments persist into the Early Valanginian. A major marine transgression can be seen in the mid-Valanginian of the offshore Uitenhage Trough that can be equated to the KirkwoodSundays River Formation boundary recognised in the onshore Uitenhage and Sundays River troughs. Late Valanginian sediments comprise outer shelf claystones, with a rich microfauna that extends up to 1Atl level. Valanginian sediments intersected west

370 of the Uitenhage Fault (boreholes 25 and 26 on Fig. 20) are very sandy, accumulated in marginal marine environments, and may provide potential reservoirs. Because of the extremely variable depositional environments and the frequent changes from marine to non-marine conditions in the offshore Uitenhage Trough, microfaunal correlation is often difficult to apply, and seismic correlation has been more effective in interpreting geological history

Sundays River Trough. Outcrop and borehole data from this trough have been described in some detail (Rigassi, 1968; Dingle, 1969; Brenner and Oertli, 1976; Shone, 1976; Winter, 1973, 1979; McLachlan and McMillan, 1976, 1979; Valicenti and Stephens, 1984). Along the northern rim and around the Kirkwood Panhandle are volcanics of the Suurberg Group, dated as Callovian (McLachlan and McMillan, 1976). Nonmarine red and white conglomerates and sandstones of the Enon Formation are extensively developed in the north of the trough, and conformably overlie the Suurberg Group. In the trough depocentre basement (horizon D) has been intersected at its maximum depth of 4250 metres (Fig. 4) and consists of Devonian Bokkeveld sediments. They are overlain by less than 200 metres of Enon conglomerate over which are developed thin fluvial sandstones of the Swartkops Member (Winter, 1973). The Swartkops Member is overlain by the Colchester Member, sensu stricto (Fig. 21) which comprises interbedded grey and red claystones with sandstones. The Colchester Member attains 200 metres thickness and contains lacustrine microfauna and microflora. It is confined to the depocentre of the Sundays River Trough, and towards the margins it is replaced by red beds of Kirkwood Formation type. This interpretation supersedes that presented by McLachlan and McMillan (1976, Fig. 9) in which the Colchester Member was correlated with the marginal marine beds at Dunbrodie and Bezuidenhouts River: these latter beds are regarded as lying at the Kirkwood-Sundays River Formation transition. The Colchester Member shows intermittent organic enrichment and is clearly a viable source rock in parts of the trough as indicated by oil shows in boreholes 29 and 30 (Broad, 1990). Above the Colchester Member is a thick sequence of fluvial red and green claystones and sandstones (main body of the Kirkwood Formation), overlying which are marine grey claystones and sandstones of the lower Sundays River Formation that contain a diverse microfauna and macrofauna (Fig. 3). Foraminifera indicate that the base of the Sundays River Formation is about mid-Valanginian in age. The sequence onlaps the Palaeozoic high on the northern, upthrown side of the St. Croix Fault. The 1At 1 drift-onset unconformity marks a major

I.K. McMILLAN et al. change in sedimentation style in both the Gamtoos and Algoa basins. At most localities where a 1Atl to 6Atl (Latest Valanginian to Hauterivian) interval is preserved in the two basins, post-lAtl sediments tend to be deeper marine, and to have accumulated in a markedly less well-oxygenated environment. Horizon 1Atl is associated with the first downhole appearance of the foraminifera Lenticulina "coegaensis" amongst others, in the Pletmos and offshore Gamtoos basins. This datum lies at the top of the lower third of the Sundays River Formation in the onshore Sundays River and Uitenhage troughs of the Algoa Basin (Fig. 3). It is thus clear that 1Atl exists as a surface within the Sundays River Formation, but its character is scarcely recognisable on borehole logs, and seismic quality from the onshore Algoa Basin is extremely poor.

Transitional-early drift tectonics and sedimentation (1Atl to 13Atl) 1Atl to 6Atl interval. Only an incomplete record of sedimentation during the 1Atl to 6Atl (Latest Valanginian to Hauterivian) interval now remains, because of major uplift at 6Atl times and the consequent intense erosion thereafter until about 13Atl times (Early Aptian). Both in the Gamtoos and Algoa Basins, IAtl to 6Atl deposits are best preserved on the downthrown sides of the major bounding faults and in other structurally low areas. The I Atl to 6Atl interval is preserved only in the central Gamtoos Basin, from about the localities of borehole 15 in the west to boreholes 20 and 21 in the east (Fig. 1), and has been removed in the south. In the offshore Algoa Basin, 1At 1 to 6At 1 deposits are preserved only against the Uitenhage Fault (borehole 26, Fig. 20) and against the St. Croix Fault. Onshore they occur against the Coega Fault in the Uitenhage Trough, and in the depocentre of the Sundays River Trough. Everywhere the upper surface of this interval is intensively planed. In the northern part of the offshore Gamtoos Basin, and in the onshore Uitenhage and Sundays River troughs (the upper Sundays River Formation), 1Atl to 6At 1 sediments are composed of grey claystones with thin interbedded sandstones. Depositional environments indicate shelf conditions, with a well oxygenated sea floor resulting in diverse and abundant microfaunas. Sediments in the Sundays River Trough are markedly richer in sandstones than elsewhere: estuarine and transitional marine conditions occur along the northeastern margin of the trough. Southwards, depositional environments deepen to upper slope (perhaps deeper) in the southem Gamtoos Basin and in the Algoa Basin south of the St. Croix Fault. Here, sediments are clayier with occasional turbidite sandstone stringers. In addition, sea floor conditions were poorly oxygenated, and

LATE MESOZOIC SEDIMENTARY BASINS OFF THE SOUTH COAST OF SOUTH AFRICA

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microfaunas generally consist of Radiolaria, small numbers of agglutinated foraminifera and fish bone. However, immediately above 1Atl is an interval containing abundant agglutinated foraminifera (including the distinctive Trochamminoides sp. A) and Radiolaria. This interval is organically enriched and has some value as a source rock, although only locally is it sufficiently deeply buried for it to be of economic significance. In both the Gamtoos and Algoa basins there are structural features which can be attributed to tectonic strain along the Agulhas-Falkland Fracture Zone as continental separation took place. These include the Gamtoos Anticline, and reverse and strike slip faulting. The Gamtoos Anticline (Figs. 1, 18, 22) lies adjacent to the Gamtoos Fault and is a tight fold 30 kilometres long which plunges away from the fault plane at its northern and southern ends. Its asymmetric geometry precludes a rollover origin, and suggests formation as a result of uplift and compression during clockwise rotation of the Gamtoos Basin and Recife Arch, induced by drag caused by movement along the right-lateral Agulhas-Falkland Fracture Zone. Further evidence for a compressional phase of defor-

371

mation lies in the western Gamtoos Basin, where a dramatic oversteepened normal fault has been identified as an inversion feature (between boreholes 15 and 16, Fig. 17). Reverse faults are also known in the Gamtoos Basin, and some displace the Gamtoos Fault plane. Evidence for strike-slip displacement can be seen on the two-way time contour map of the Gamtoos Anticline (Fig. 22). The axis of the anticline is displaced by numerous strike-slip faults, most of which are fight lateral and appear to have originated at 1Atl times. This fault pattern implies that left lateral displacement took place along the Gamtoos Fault after development of the Gamtoos Anticline, as the Agulhas-Falkland Fracture Zone shearing process developed. Left lateral strike-slip movement is also expected to have occurred on the other major basin bounding faults in the southern offshore, but cannot be demonstrated from seismic data. In contrast, in the Algoa Basin, transitional rift tectonics seem to be evident in the faulting alone, rather than in the development of anticlinal structures. Nevertheless, a distinct cross-basin basement arch can be recognised on the horizon D contour map (Fig. 4) extending roughly east-west across the Port Elizabeth Trough, through boreholes 23 and 24 (Fig. 1). This basement high has influenced sedimentation patterns up to the level of 1Atl and is on trend with the highest part of the Gamtoos Anticline east of borehole 22, and it is possible that the two features are related in origin. No evidence has been found of compressional tectonics in the Uitenhage Trough. Tensional processes in the southern Uitenhage Trough at 6Atl times led to the development of the Uitenhage Fault, which has a throw of about 2000 metres.

6Atl to 13Atl period of erosion. During the 6Atl to 13Atl period (earliest Barremian to Early Aptian) two canyons were scoured - - a small one in the Gamtoos Basin close to the Gamtoos Fault (Fig. 17) which partly eroded the Gamtoos Anticline, and a much larger one (60 kilometres long, average of 30 kilometres wide and 1 kilometre deep) developed in the Uitenhage Trough (Fig. 23). This Algoa Canyon cuts across the Uitenhage Fault exposing basement on the upthrown side (Fig. 20). A probable arm of the canyon, now severed by late reactivation of the Port Elizabeth Fault, lies in the northern Port Elizabeth Trough. Erosion at 6At 1 locally cuts down well below 1Atl. A two-way time contour map of the canyon floors and hinterland illustrates the extent of these features (Fig. 23) and it is suggested that tectonic uplift was responsible for their origin. Comparison with the other basins. A comparison of sedimentation processes active during the 1Atl to 13Atl time period (latest Valanginian to

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LATE MESOZOIC SEDIMENTARY BASINS OFF THE SOUTH COAST OF SOUTH AFRICA Early Aptian) in the four major southern offshore basins reveals substantial differences. In the Bredasdorp Basin the 1A to 4A sequences are restricted to the central basin area, followed by a major 5Atl transgressive episode initiating sedimentation over wide areas along the basin's northern and eastern margins and in the Infanta Embayment, thus suggesting more rapid subsidence of these two basins at 5Atl time. Deposition of deep marine, poorly oxygenated turbiditic sandstones and black claystones was widespread throughout the Bredasdorp Basin from 1Atl to 13Atl. In the Pletmos Basin, pre-6Atl sediments are generally deep marine, poorly oxygenated and turbiditic, but in contrast post-6Atl sediments are generally better oxygenated. However, in the Gamtoos and Algoa basins, deep marine turbiditic sedimentation was developed up to 6Atl times, after which major uplift occurred that incised a major canyon system in the Algoa Basin and a smaller one in the Gamtoos Basin. No sedimentation occurred in these two basins until Late Aptian-Early Albian (post 13At 1), when the inactive canyons began to be infilled with marine claystone (Figs. 3, 21).

Drift tectonics and sedimentation (13Atl to present day) Canyon infilling commenced at or shortly after 13Atl (Late Aptian) and indicates the termination of erosion and sediment-conduiting into deep water. Where canyon-fill sediments have been intersected (boreholes 25, 26 and 27 on Fig. 20; 19 on Fig. 17) they are characterised by marine shallowing-upward sequences, mainly of claystones but with minor sandstones, with common lignite and dolomite stringers. The fill of the Gamtoos Canyon (borehole 19, Fig. 17) is dolomitic and lignitic, with phosphate pellets, "ooliths", rare foraminifera (none planktonic) and widespread though sparse Radiolaria. The depositional environment here is clearly an unusual one, and is interpreted as a slope accumulation. The fill of the Algoa Canyon is primarily claystone deposited in well-oxygenated conditions in uppermost slope to outermost shelf conditions at boreholes 25 and 26 (Fig. 20) and microfaunas are diverse and abundant. However, at the base of the canyon fill sequence in borehole 26, located close to the canyon axis, is a package of interbedded sandstones and claystones rich in lignite, but poor in microfaunas. Foraminiferal datings of this interval are ambiguous, but seismic correlation suggests a Late Aptian age. Borehole intersections of the proximal canyon fill, and also in the severed arm in the northern Port Elizabeth Trough, consist of middle to inner shelf claystones and minor sandstones that are rich in foraminifera, indicating well-oxygenated sea-floor conditions.

373

Termination of canyon sedimentation is marked by a seismic sequence boundary (Top Canyon) that correlates faunally with a horizon above 14Btl in the Bredasdorp Basin (see Fig. 3). Overlying the canyon fill is a sequence of Late Albian and possibly Early Cenomanian age that is strongly truncated by the 15Atl unconformity. This sequence is developed over the entire southern half of both the Gamtoos and Algoa basins, and marks the first return to widespread marine conditions since pre-6At 1 times (latest Hauterivian). Sediments in this interval in the southern Gamtoos Basin are inner to middle shelf glauconitic claystones with minor sandstone stringers. Foraminifera faunas, although indicating well-oxygenated conditions on the sea floor, tend to be poor, especially in the sandier, more proximal intersections. In the southern Algoa Basin this interval is everywhere much sandier, with considerable glauconite, even in the more distal boreholes such as 25 and 26 (Fig. 20) than is the case in the Gamtoos Basin. Again, depositional environments tend to be inner to middle shelf, and foraminiferal faunas are poor. It seems likely that the Algoa sandstone interval is at least partly equivalent to the sandstones of the 14Ctl to 15Atl interval in the Bredasdorp Basin. Thin Early Turonian glauconitic sandy clays overlie the 15Atl surface in the southern part of the offshore Gamtoos and Algoa basins. They accumulated in inner shelf depositional environments and lack the organic enhancement seen in equivalent intervals in the Pletmos, Infanta and Bredasdorp basins. In the northern offshore Gamtoos and Algoa basins 15Atl is compounded with horizon K, and mid-Coniacian or later deposits unconformably overlie Late Albian to Early Cenomanian rocks. The later Cretaceous (Coniacian to Maastrichtian) interval in both basins is markedly thinner than seen in the Pletmos or Bredasdorp basins, but is as complete as elsewhere. Sediments are shelf claystones with minor sandstone stringers, and contain rich microfaunas indicating well-oxygenated conditions. Following mild tectonic subsidence in the mid-Campanian, a major transgression led to deposition of the Late CampanianEarly Maastrichtian Igoda Formation and its equivalents, at Needs Camp, near East London (Klinger and Lock, 1978). It is clear that sediments of this interval were also laid down as a thin veneer over the onshore Gamtoos and Algoa Basins, but were stripped off by later erosion. The Late CampanianMaastrichtian (horizons X to L) interval has been recognised in all offshore Gamtoos and Algoa boreholes. Proximally it is composed of glauconitic claystones and sandstones that unconformably overlie pre-6Atl, Hauterivian rocks. Distally it consists predominantly of shelf claystone. Microfaunas are diverse and abundant and reflect normal sea-floor oxygen levels.

374 Cainozoic (post-horizon L) sediments consist of a thin Palaeocene glauconitic claystone sequence, u n c o n f o r m a b l y overlain by M i d d l e to Late E o c e n e biogenic limey clays, and, also in the distal G a m toos Basin, an unusually thick interval of Early Oligocene age (as seen in all boreholes on Fig. 17). Early Oligocene sediments are generally silty greenish clays with glauconite. Early M i o c e n e sediments (biogenic white limey clays) are confined to the most distal borehole drilled to date in the Algoa Basin (borehole 25, Fig. 20) and have not been recognised elsewhere in the G a m t o o s or Algoa basins. Overlying deposits are generally thin veneers of Pleistocene and H o l o c e n e age, but in borehole 25 a unique sequence of mid-Pliocene clays and earliest Pleistocene shelly sands have been recognised. This sequence reflects a r e n e w e d episode of subsidence restricted to the Algoa Basin and the western Port Alfred Arch. Equivalent deposits, mostly c o m p o s e d of shelly sands, have been described onshore as the upper Algoa Group (see McMillan, 1990b).

ACKNOWLEDGEMENTS

The authors are indebted to the m a n y Soekor geoscientists who, over the years, have contributed towards our present understanding of the geology of the southern offshore basins and their onshore components. Without their efforts this review could not have been written. This study has also been influenced by numerous discussions with our colleagues in the exploration teams and in specialist geology (petrography, geochemistry, palaeontology and sedimentology) which have led to lively debates, and have provided checks and balances for our own ideas. However, the authors accept full responsibility for the views expressed in this paper. We thank the m a n a g e m e n t of S O E K O R (Pty) Ltd for permission to publish this paper.

REFERENCES

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LATE M E S O Z O I C S E D I M E N T A R Y BASINS OFF T H E SOUTH COAST OF S O U T H A F R I C A Divergent/Passive Margin Basins. Am. Assoc. Pet. Geol. Mem., 48: 239-248. Espitali6, J. and Sigal, J., 1963. Contribution ~ l'6tude des foraminif~res (micropal6ontologie-microstratigraphie) du Jurassique Sup6rieur et du N6ocomien du Bassin de Majunga (Madagascar). Ann. Geol. Madagascar Fasc., 32: 9-100. Falvey, D.A., 1974. The development of continental margins in plate tectonic theory. Australian Petrol. Explor. Assoc. J., 14: 95-106. Fatti, J.L., Strauss, P.J., Stallbom, K.E., 1995. A three-dimensional seismic survey over the offshore F-A gas field. S. Afr. Geophys. Rev., 1: 1-22. Fouch6, J., 1990. The F-X prospect: submarine fan submarine slump? Abstract, Geocongress 90, Geol. Soc. S. Afr., Abstr. Vol., pp. 161-164. Fouch6, J., Bate, K.J. and Van der Merwe, R., 1992. Plate tectonic setting of the Mesozoic Basins, southern offshore, South Africa: a review. In: M.J. de Wit and I.G.D. Ransome (Editors), Inversion Tectonics of the Cape Fold Belt, Karoo and Cretaceous Basins of Southern Africa. A.A. Balkema, Rotterdam, pp. 33-48. Friedinger, P.J.J., 1986. A model for the formation of the Outeniqua Basin (Agulhas Bank) and possible plate tectonic implications. Tech. Rep. Joint Geol. Surv. S. Afr. Univ. Cape Town Mar. Geosci. Group, 17: 39-50. Gerrard, I. and Smith, G.C., 1982. Post-Paleozoic succession and structure of the southwestern African continental margin. In: J.S. Watkins and C.L. Drake (Editors), Studies in Continental Margin Geology. Am. Assoc. Pet. Geol. Mem., 34: 49-74. Gilbert, C.E., 1990. Sandstone distribution at the 14Atl level in the E-G area, south central Bredasdorp Basin: implications for a predictive depositional model. Abstract, Geocongress 90, Geol. Soc. S. Afr., Abstr. Vol., pp. 177-180. Gouws, H.J., 1990. The Ga-W prospect: a comparison of the seismic stratigraphic prognosis with drilling results. Abstract, Geocongress 90, Geol. Soc. S. Afr., Abstr. Vol., pp. 181-184. Haq, B.U., Hardenbol, J. and Vail, P.R., 1987. The new chronostratigraphic basis of Cenozoic and Mesozoic sea level cycles. In: C.A. Ross and D. Haman (Editors), Timing and Depositional History of Eustatic Sequences: Constraints on Seismic Stratigraphy. Cushman Found. Foraminiferal Res., Spec. Publ., 24: 7-13. Hill, S.J., 1990. Sandstones of the central Bredasdorp Basin depositional and diagenetic controls on reservoir quality. (Abstract) Geocongress 90, Geol. Soc. S.Afr., Abstr. Vol., pp. 682-685. Hodges, K. and Winters, S., 1990. Submarine fans in the central Bredasdorp Basin. Abstract, Geocongress 90, Geol. Soc. S. Afr., Abstr. Vol., pp. 686-689. Jeletzky, J.A., 1 9 8 3 . Macroinvertebrate palaeontology, biochronology and paleoenvironments of Lower Cretaceous and Upper Jurassic rocks, Deep Sea Drilling Hole 511, Eastern Falkland Plateau. In: J.H. Blakeslee and M. Lee (Editors), Initial Reports Deep Sea Drilling Project, 71:951-975. Jones, D.L. and Plafker, G., 1977. Mesozoic megafossils from DSDP Hole 327A and Site 330 on the Eastern Falkland Plateau. In: P.E Barker, I.W.D. Dalziel et al., Initial Reports Deep Sea Drilling Project, 36: 845-856. Kent, P.E., 1974. Continental margin of East Africa m a region of vertical movements In: C.A. Drake and C.L. Drake (Editors), The Geology of Continental Margins. Springer, New York, N.Y., pp. 313-320. Klinger, H.C. and Lock, B.E., 1978. Upper Cretaceous sediments from the Igoda River mouth, East London, South Africa Ann. S. Afr. Mus., 77: 71-83. Leith, M.J. and Rowsell, D.M., 1979. Burial history and temperature-depth conditions for hydrocarbon generation and migra-

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tion on the Agulhas Bank, South Africa. Geokongres 77, Geol. Soc. S. Afr. Spec. Publ., 6:205-217. Light, M.P.R., Garcia-Victoria, E., Turner, J. and McLachlan, I.R., 1982. Petroleum geology of the F-A gas field, offshore South Africa. Abstract, 3rd Symp. Sediment. Div., Geol. Soc. S. Afr., Johannesburg, 1982, Abstr. Vol., pp. 69-72. Maier, J.J., 1990. The complexities of the sedimentary and reservoir models in the Superior Structure. Abstract, Geocongress 90, Geol. Soc. S. Afr., Abstr. Vol., pp. 325-327. Malan, J.A. and Viljoen, J.H.A., 1990. Mesozoic and Cenozoic geology of the Cape South Coast. Guidebook Geocongress '90, Geol. Soc. S. Afr., P03: 1-81. Malan, J.A., Martin, A.K. and Cartwright, J.A., 1990. The structural and stratigraphic development of the Gamtoos and AIgoa Basins, offshore South Africa. Abstract, Geocongress 90, Geol. Soc. S. Afr., Abstr. Vol., pp. 328-331. Marot, J.E.B. and McLachlan, I.R., 1982. Petrography of the reservoir sandstones of the gas-bearing F-A structure, offshore South Africa. Abstract, 3rd. Symp. Sediment. Div., Geol. Soc. S. Afr., Johannesburg, 1982, Abstr. Vol., pp. 74-78. Martin, A.K. and Hartnady, C.J.H., 1986. Plate tectonic development of the south west Indian Ocean: a revised reconstruction of East Antarctica and Africa. J. Geophys. Res., 91: 47674786. Martin, A.K., Hartnady, C.J.H. and Goodlad, S.W. 1981. A revised fit of South America and south central Africa. Earth Planet. Sci. Lett., 54: 293-305. Martin, A.K., Goodlad, S.W., Hartnady, C.J.H. and Du Plessis, A., 1982. Cretaceous palaeopositions of the Falkland Plateau relative to southern Africa using Mesozoic seafloor spreading anomalies Geophys. J. R. Astron. Soc., 71: 567-579. McLachlan, I.R., 1977. Palynological assessment of the petroleum source-rock potential of the Mesozoic and Cenozoic sediments on the coastal margin of South Africa. Abstract, Geokongres 77, Geol. Soc. S. Afr., Abstr. Vol., pp. 60-63. McLachlan, I.R., Brenner, P.W. and McMillan, I.K., 1976. The stratigraphy and micropalaeontology of the Cretaceous Brenton Formation and the PB-A/I well, near Knysna, Cape Province Trans. Geol. Soc. S. Afr., 79: 341-370. McLachlan, I.R. and McMillan, I.K., 1976. Review and stratigraphic significance of southern Cape Mesozoic palaeontology. Trans. Geol. Soc. S. Afr., 79: 197-212. McLachlan, I.R. and McMillan, I.K., 1979. Microfaunal biostratigraphy, chronostratigraphy and history of Mesozoic and Cenozoic deposits on the coastal margin of South Africa. Geokongres 77, Geol. Soc. S. Afr. Spec. Publ., 6: 161-181. McMillan, I.K., 1990a. Foraminiferal definition and possible implications of the major mid-Cretaceous (Albian to Coniacian) hiatuses of southernmost Africa. Abstract, Geocongress 90, Geol. Soc. S. Afr., Abstr. Vol., pp. 363-366. McMillan, I.K., 1990b. A foraminiferal biostratigraphy and chronostratigraphy for the Pliocene to Pleistocene Upper AIgoa Group, eastern Cape, South Africa. S. Afr. J. Geol., 93: 622-644. Mutti, E. and Normark, W.R., 1987. Comparing examples of modern and ancient turbidite systems: problems and concepts. In: J.K. Leggett and G.G. Zuffa (Editors), Marine Clastic Sedimentology Concepts and Case Studies. Graham and Trotman, pp. 1-38. Norton, 1.O. and Sclater, J.G., 1979. A model for the evolution of the Indian Ocean and the breakup of Gondwanaland. J. Geophys. Res., 84: 6803-6830. Partridge, T.C. and Maud, R.R., 1987. Geomorphic evolution of Southern Africa since the Mesozoic. S. Afr. J. Geol., 90: 179-208. Pferdek~imper, H.W., Ranoszek, M., Holmes, L.C., Burger, C.A.J. and Wenham, M., 1992. E-M volumetric evaluation m 1992. Soekor Report, 57 pp. (unpublished).

376 Powell, C.McA., Johnson, B.D. and Veevers, J.J., 1980. A revised fit of East and West Gondwanaland. Tectonophysics, 63: 13-29. Quennell, A.M., McKinlay, A.C.M. and Aitken, W.G., 1956. Summary of the geology of Tanganyika, Part 1. Introduction and stratigraphy. Geol. Surv. Tanganyika, Mem., 1: 1-264. Questiaux, J.-M., Strauss, P.J., Unstead, P. and Stallbom, K.E., 1985. Geophysical and geological appraisal of the F-A gas field. Soekor Report, 26 pp. (unpublished). Rigassi, D.A., 1968. Preliminary report on the geology and oil prospects of the Sundays River Basin. Report, Petroconsultants, S.A., Geneva, 35 pp. (unpublished). Rigassi, D.A., 1970. Cretaceous of the Cape Province, Republic of South Africa. Report, Petroconsultants, S.A., Geneva, 82 pp. (unpublished). Rigassi, D.A. and Dixon, G.E., 1972. Cretaceous of the Cape Province, Republic of South Africa. In: Proc. Ibadan Univ. Conf. on African Geology, December 1970, pp. 513-527. Rust, I., 1983. Excursion Guidebook to Sedplett '83. Univ. Port Elizabeth, Geol. Surv. S. Afr., 91 pp. Scrutton, R.A. and Du Plessis, A., 1973. Possible marginal fracture ridge south of South Africa. Nature, 242:180-182. S6goufin, J., 1978. Anomalies magn6tiques mesozo'l'ques dans le bassin de Mozambique. C. R. Acad. Sci., Paris, Ser. D, 291: 109-112. Shone, R.W., 1976. The sedimentology of the Mesozoic Algoa Basin. M.Sc Thesis, Univ. Port Elizabeth, pp. 1-48 (unpublished). Silva, G.A., 1966. Sobre o ocorr6ncia do Jur~issico marinho no norte de Moqambique. Estud. Gerais Univ. Moqambique, 3: 61-68. Smith, A.G. and Briden, J.C., 1977. Mesozoic and Cenozoic Paieocontinental Maps. Cambridge University Press, Cambridge, London, New York, 63 pp. Strauss, P.J., Schreuder, S., Holmes, L.C., Willis, R., Steyn, A.V.C., Rathan, D.G.L., 1990. F-A dry gas initially in place sensitivity evaluation. Soekor Report, 41 pp. (unpublished). Tankard, A.J., Jackson, M.P.A., Erikson, K.A., Hobday, D.K.,

I.K. M c M I L L A N et al. Hunter, D.R. and Minter, W.E.L., 1982. Crustal Evolution of Southern Africa. Springer-Verlag, New York, N.Y., 523 pp. Turner, J.R., 1990. The geological evolution of the Upper Jurassic syn-rift interval within the Infanta Embayment, offshore Republic of South Africa. Abstract, Geocongress 90, Geol. Soc. S. Afr., Abstr. Vol., pp. 558-561. Vail, P.R., Bubb, J.N., Hatlelid, W.G., Mitchum, R.M., Sangree, J.B., Thompson, J.B., III, Todd, R.G. and Widmier, J.M., 1977. Seismic stratigraphy and global changes of sea level, parts 1 to 11. In: G.E. Peyton (Editor), Seismic Stratigraphy m Applications to Hydrocarbon Exploration. Am. Assoc. Pet. Geol. Mem., 26: 49-212. Valicenti, V.H. and Stephens, J.M., 1984. Ostracods from the Upper Valanginian and Upper Hauterivian of the Sundays River Formation, Algoa Basin, South Africa. Rev. Esp. Micropaleontol., 16:171-239. Van der Merwe, R. and Fouch6, J., 1992. Inversion tectonics in the Bredasdorp Basin, offshore South Africa. In: M.J. de Wit and I.G.D. Ransome (Editors), Inversion Tectonics of the Cape Fold Belt, Karoo and Cretaceous Basins of Southern Africa. A.A. Balkema, Rotterdam, pp. 49-60. Van Wagoner, J.C., Mitchum, R.M., Jr., Posamentier, H.W. and Vail, P.R., 1987. Seismic stratigraphy interpretation procedure using sequence stratigraphy, part II: key definitions of sequence stratigraphy. In: A.W. Bally (Editor), Atlas of Seismic Stratigraphy. Am. Assoc. Pet. Geol., Studies Geol., 27:11-14. Wickens, H. de V., 1989. The sedimentology of the C-to-D sequence in the E-M/E-S/E W gasfield area, northern Bredasdorp Basin. Soekor Report, 24 pp. (unpublished). Winter, H. de la R., 1973. Geology of the Algoa Basin, South Africa. In: G. Blant (Editor), Sedimentary Basins of the African Coasts. 2nd Part: South and East Coasts. Assoc. Afr. Geol. Surv., Paris, pp. 17-48. Winter, H. de la R., 1979. Application of basic principles of stratigraphy to the Jurassic-Cretaceous interval in southern Africa. Geokongres 77, Geol. Soc. S. Afr. Spec. Publ., 6: 183196.

Chapter 14

Puzzling Questions in the Simple History of a Continent

KENNETH J. H S 0

The geology of the Africa is simple. As the volume editor wrote: the sedimentary basins of Africa are largely of two types, sag basins and failed rifts. When I accepted the appointment of the series editor of the project Sedimentary Basins of the World, I acquired a contractual previlege to write in my style a summary at the end of each volume to interpet the geology on the basis of my understanding. But the geology of Africa seems so straight forward, the articles are so well-written, and the volume is so well edited that a summary by the series editor could be superfluous. After reading through the manuscripts, however, I have the impression that the simplicity could be deceiving. There are still possibilities for alternative interpretations, and there are still puzzling questions in the geologic history of Africa. Of those I have chosen three to discuss in this chapter. (1) Where was the the northern plate-margin of Africa? (2) What is a retro-arc foreland basin? (3) Is the East African Rift destined to become an aulacogen? Not being very knowledgable about Africa, I shall not be able to offer a definitive synthesis. I shall attempt, however, to discuss those questions on the basis of my understanding of the geologic history of Africa. At the risk of offending the authors and the volume editor, I venture to offer some ideas which might eventually help the solution of those puzzling problems.

WHERE WAS THE NORTHERN PLATE-MARGIN OF

AFRICA? Two of the great proponents of the theory of continental drift came from the Southern Hemisphere: W. Carey from Australia and A. Du Toit from South Africa. The fact is not surprising when we appreciate that the geology of the two southern

continents can be best understood by the unifying theory, postulating an accretional phase during the Palaeozoic, leading to the formation of the supercontinent Pangaea, and a fragmentation phase during the Mesozoic and Cenozoic, leading their present present configuration. The sag basins of Africa owed their genesis, on the whole, to the Palaeozoic accretion, and the failed rifts were formed during the later fragmentation. Few challege the conclusion that the Karoo Basin of South Africa is a late Palaeozoic foreland basin (Chapter 12), and none have ever questioned that East African basins are intracratonic rifts (Chapter 9). There are, however, unresolved problems where the basins include sediments deposited during both phases of the tectonic evolution of Africa. The problem is particularly acute when one interpretes the geology of the sedimentary basins of Northwest Africa (Chapters 1 and 2), because the controversy on the positioning of the northern margin of the African plate boundary can lead to different geological interpretations. While I do not question the "fundamental truth about the geology of Arabia and North Africa (that those) lands [... ] together once formed the southern shores of the great Tethyan Ocean" (Chapter 1), many of us do wonder where was the shoreline exactly? Selley wrote (Chapter 1): It is 'believed [ . . . ] that the Atlas Fold Belt accreted onto the African Shield during late lateral crustal movement, as the Tethyan Ocean closed into the Mediterranean Sea (Clifford, 1986). [... ] The Atlas Mountains are newcomers, and thus not really a part of mainland Africa. Having been a student of the Alpine-Mediterranean tectonics for quarter of a century, I beg to differ, and suggest that the geology of Africa is more easily understood if the Atlas Mountains are not newcomers? Is it possible that the Atlas terrane was a part of Mesozoic Africa and had been a part of

African Basins. Sedimentary Basins of the World, 3 edited by R.C. Selley (Series Editor: K.J. Hsti), pp. 379-382. 9 1997 Elsevier Science B.V., Amsterdam. All fights reserved.

380 Africa even before the late Palaeozoic deformation during the accretional phase of the African tectonic history? The kinematics of the movement of Africa relative to Europe since the Early Jurassic (180 Ma) is well known, being constrained by the seafloorspreading data of the North Atlantic: there has been first a left-lateral displacement of some 3000 km during the first 90 million years, followed by a counterclockwise rotation during the second 90 million years. Did Atlas moved with Africa, or was the terrane European until it was "accreted onto the Africa Shield during late lateral crustal movement."? The present plate-boundary between Africa and Europe lies north of the High Atlas; the Tell Atlas and the Rif are a part of the European plate. In my latest interpretation of the Alpine-Mediterranean tectonics, I postulated that the plate was bounded until Miocene by an island-arc margin: the Mediterranean Ridge was the buffed eastern segment of the arc, and the Tell-Atlas/Rif was the western segment of the arc (Hsti, 1996). The latter collided with Africa to form the High Atlas and the Saharan Atlas. Using the terminology of tectonic facies, the Tell-Atlas/Rif is a rhaetide, or the overriding block, which has been thrust onto the African margin; the sedimentary strata north of the Sahara Flexure constitute an alemanide, having been pushed out of the underthrust margin to form a Cenozoic foreland fold belt. The northern Africa margin could not have been a passive-margin during the Mesozoic and Cenozoic. The eastward movement of Africa must have taken place along a trancurrent fault or a zone of transcurrent faults. The Sahara Flexure was one transcurrent fault near the plate-margin. I have not studied the geology of the Atlas, but my colleague John Warme has. He found evidence for left-lateral movement along the Sahara Flexure, but the displacement cannot be very large. The transtensional stress induced by movement led to the formation of a series of rhomb-shaped basins in the Atlas Mountains (Warme, 1988). The master-fault along which most of the 3000 km displacement took place is not the Sahara Flexure, but should be located north of the High Atlas; it must be buried now under the overriding rhaetide of Rif and Tell Atlas. The hypothesis of a transcurrent margin between the European and African plates provides an explanation of the absence of an ophiolitic celtide between the Rif/Tell Atlas and the High Atlas, because there should have been little seafloor-spreading on a transcurrent margin. The Tethys Ocean, even if its oceanic lithosphere did extend as far west as northwest Africa, must have been a narrow seaway beteen the Atlantic and the Mediterranean before the Cenozoic arc-continent collision. Any ophiolite

KENNETH J. HSU

melange that might be present, like the transcurrent plate-boundary, is now buried under the rhaetide of Rif and Tell Atlas. My hypothesis gives a portrait of the Mesozoic North African margin more akin to the Pacific than to the Atlantic margin of North America There should have been a coastal barrier, if not a coast range, which served to restrict the Triassic marginal seas of northwest Africa and to isolate them at times from the Tethys, permitting thus the deposition of Triassic and Lower Jurassic salts in northwest Africa (Chapter 2). Thick Palaeozoic strata are present in the Northwest African basins (Chapter 2). Situated south of the Palaeozoic foreland fold-and-thrust belt of the Atlas, those sediments have also undergone foreland deformation during the late Palaeozoic orogenesis, when Africa was sutured onto North America.

WHAT IS A RETRO-ARC FORELAND BASIN?

The Karoo Basin was "classified as a foreland basin, since it contains a thick flysch-molasse wedge which flanks the front of a mountain chain and wedges out northward over the adjacent craton [... ] The basin constitutes a retro-arc basin as defined by Dickinson (1974), situated behind a magmatic arc and associated fold-thrust belt (Cape Fold Belt) produced by northward subduction of oceanic lithosphere located south of the arc" (Chapter 2). I have no quarrel with this interpretation except an implication that the genesis retroarc basin was related to the forearc subduction of oceanic lithosphere. In my review of global tectonics (Hsti, 1996), I found that the so-called retro-arc foreland basins are, in every instance, a successor basin of collapsed back-arc basin. The theory of orogenesis by the process of back-arc basin collapse was proposed by Ian Dalziel (1981). Circum-Pacific forearc subduction caused the formation of back-arc basins behind frontal arcs. The Mesozoic subduction of the Pacific plate down an east-dipping Benioff Zone under the Chilean Andes led, for example, to the genesis of a backarc basin in Argentine. The ocean lithosphere under the back-arc basin was, however, subducted during the early Tertiary down a west-dipping Benioff Zone, and the consumption of the ocean floor, i.e., the collapse of the backarc basin, caused the orogenesis of the Southern Andes. A fiysch-molasse sequence was then filling the foredeep where the backarc basin had collapsed. What Dickinson once called a retro-arc foreland basin is in fact a successor basin of a collapsed back-arc basin. The archiepelago model of orogenesis finds an actualistic analogue in the tectonic evolution of the

PUZZLING QUESTIONS IN THE SIMPLE HISTORY OF A CONTINENT South China Sea (Hsti, 1994). The back-arc basin, situated behind the Banda-Sunda arc, is a back-arc basin; it was formed by seafloor spreading during the Oligocene and early Miocene time. Since middle Miocene, oceanic lithosphere under the South China Sea has been subducted under the inner wall of the Manila trench at a rate of a few centimeters per year. One can visualize the complete consumption of the ocean floor in about 50 million years time, when the Philippine Arc will collide with the South China margin. The collapsed back-arc basin of South China Sea, much smaller at that time,would receive detritus derived from the mountains formed by the collision, and a flysch-mollasse sequence will be deposited in that foreland basin. Applying the back-arc basin collapse model to interpret the geology of South Africa, the Cape Supergroup of the Cape Folded Belt can be considered the shallow marine sedimentary cover of an island arc; the African margin south of the Cape was an active, not a passive margin, during the Palaeozoic. The subduction of ocean lithosphere down a northdipping Benioff Zone under the active margin caused the formation of a back arc basin north of the Cape Arc. During the late Palaeozoic, the floor of that back-arc basin was subducted down a Benioff Zone dipping south under the arc. The subsequent collapse of the back-arc basin led to an arc (Cape)-continent (Africa)collision, when the backarc basin was converted into a foreland basin. The glacio-marine sediments of the Dwyka Group were laid down in this initially deep basin, before it was eventually filled up by the thick flysch-molasse sequence of the Karoo Supergroup. It seems pedantic to argue on basin-nomenclature, but a name is a name is more than a name. According the scheme proposed for this series, the so-called retro-arc basin is a composite of an isostatically adjusted, extensional, back-arc basin (1.1.4) and an isostatstically not adjusted, compressional, foreland basin (2.1.1). Retro-arc basin finds no place in the genetic classification of The Sedimentary basins of the World. We allow each author the freedom to speak his own dialect, or use his own language. My job, as the series editor, is to enhance the unity and harmony of the whole series through a calibration of the meaning of the terms.

IS EAST AFRICAN RIFT DESTINED TO BECOME AN

AULACoGEN? In his discussion of the origin of the East African Rift, Frostick distinguished mantled-generated active rifts and lithosphere-generated passive rifts (Chapter 9). Lithospheric stretching could be a consequence of an actively rising asthenosphere, or an active

381

stretching of lithosphere could induce a passive rise of asthenosphere. Doming, arching, and uplift on a regional scale was considered by Frostick characteristic of active rifting, whereas the the first topographic expression of passive rifting should be subsidence, "initiated as a result of tension in the lithosphere which stretches and thins it" (Chapter 9). Citing the evidence of early doming, Frostick considered thus the East Africa Rift a good example of mantle-generated rifting. I take exception to this interpretation. There could be a chicken-and-egg type of argument. A lithospheric stretching causing the first thinning caused at about the same time also the asthenosphere to rise. The process continued then for some time: the lithosphere continued to be stretched while the asthenosphere continued to rise, so that we never see any evidence of subsidence. In other words, the chronological precision of our geological evidence does not permit a determination whether the lithospheric thinning or the topographical doming came first. Theoretical considerations can throw some light on the problem. If asthenospheric rising is the triggering mechanism, why did it start in the first place and why did it stop? If lithospheric stretching is the triggering mechanism, why did it start in the first place and why did it stop? Or is it to stop? Asthenosphere rises along the axis of seafloor spreading: asthenosphere has been rising, for example, under the Mid-Atlantic Ridge for 180 million years at least. The process continues because of the positive feedback mechanism of a convection cell. After a spreading ridge becomes the site of a sink of asthenospheric heat, the thinning of the crust above the ridge will cause the ridge-axis to remain as the site of the heat-sink. A steady-state convection cell is thus established, as suggested by the linear rate of seafloor spreading. It would take a major external perturbance to disturb the steady state. The crust of East Africa "has been put under stress a number of times in its long history," and the East African Rift had "its inception in the Oligocene"(Chapter 9). But there is no evidence for a steady-state convection cell generated by an actively rising asthenosphere under East Africa. On the contrary, the rise of the asthenosphere has been a stop-and-go process. In contrast to the continuous flow of asthenosphere, the faulting induced by lithospheric stretching is a stop-and-go process. Rifting will take place only when the gradually accumulating extensional stress in the lithosphere reaches a magnitude greater than its critical yield strength. After its release by earthquake faulting, the extensional stress in the crust can start to accumulate again until the critical yield stress is again exceeded.

382 Where could the extensional stress come from to initiate the stop-and-go rifting? The answer lies in an analysis of the stress and strain orientations. The major rift basins of East Africa vary in orientation from almost N - S to N W - S E and N E - S W (Chapter 12). A N - S orientation is more or less parallel to sea-spreading direction of the Indian Ocean south of the continent, and it is also parallel to the direction of compression of the Alpine-Mediterranean system. Rock-deformation experiments indicate that extensional fractures can form under compression in a direction parallel to the axis of compression, and that shear fractures can form under compression in directions forming an acute angle bisected by the axis of compression. Applying this principle to interpret the African tectonics, I suggest that the East African Rift was formed because a N - S compression, equivalent to an E - W extension, has been induced when the northward movement of Africa was hindered by its collision with the Europeans plate. The N - S trending rift valleys are extensional fractures, and the N W - S E and the N E - S W trending rift valleys are surface expressions of shear fractures, all induced by a N - S compression. The stop-and-go mechanism of rifting suggests a complicated kinematic history of the movement of Africa, and the variety of mechanisms of crustal deformation in response to the various crustal stresses induced by the external forces of compression. If my interpretation is correct, the East African Rift will never become a major ocean. On the other hand, one cannot consider the structure a "failed rift or an aulacogen which will slowly die in the next few million years" (Chapter 12). The Rift never had an ambition to become an ocean and cannot be said to have failed, the rifting will "stop and go" as Africa continues to march northward.

K E N N E T H J. HSO

I have indicated that the volumes of the series of the Sedimentary Basins of the World will not be philately albums; they will not be collections of random observations. They will be pieces of mosaic to be pieced together for a unifying theory of global tectonics. I have indicated in my introduction to the South Pacific Volume that there is a parallelism in the pattern of orogenic deformation in China and in South Pacific: China, as a part of Eurasia, has undergone a billion years of amalgamation, while South Pacific is still in an earlier phase of accretion. Taking such a viewpoint, one sees in the geology of Africa the same manifestation of the same fundamental processes in a different stage of development. Africa has undergone through a fragmentation stage, and it is being pushed toward Eurasia to its ultimate destiny of a place in a supercontinent.

REFERENCES

Clifford, A.C., 1986. African o i l - past, present and future. In: M.T. Halbouty (Editor), Future Petroluem Provinces of the World. Am. Assoc. Pet. Geol. Mem. Dickinson, W.R., 1974. Plate tectonics and sedimentation. In: W.R. Dickinson (Editor), Tectonics and Sedimentation. Soc. Econ. Paleontol. Mineral., Spec. Publ., 22: 1-27. Dalziel, I.W.D., 1981. Back-arc extension in the southern Andes. A review and critical appraisal. Trans. R. Soc. London, Ser. A, 300: 319-335. Hsti, K.J., 1994. Tectonic facies in an archipelago model of intra-plate orogenesis. Geol. Today, 4: 190-294. Hsti, K.J., 1996. Geology of Switzerland: Introducing the Tectonic Facies Concept. Princeton University Press, Princeton, NJ. Warme, J.E. (Editor), 1988. Evolution of the Jurassic High Atlas Rift, Morroco: Transtension, Structural and Eustatic Controls on Carbonate Facies. Tectonic Inversion: Guidebook Am. Assoc. Pet. Geol. Field Seminar. Publ. No. 9, Exploration Geol. Inst., Colorado School of Mines, Golden. CO.

A u t h o r Index *

Abd-Allah, A.M. 85 Abdallah, A.M. 48, 49, 53, 55, 60, 82 Abdel Aal, A. 43, 82 Abdel Salam, Y. 124, 132, 134, 140, 147 Abdel Shafie, M. 132, 133, 140, 141,148 Abdin, S. 58, 82 Abdine, S.A. 85 Abdolrahman, K. 26 Abdula. H.H. 149 Abed, A.M. 10, 16 Abomo, 172 Abul Nasr, R.A. 71, 72, 82 Aburawi, R.M. 82 Adamson, D.A. 119, 124, 137, 140-146, 147, 149

Agagu, O.K. 156, 171 Ahmed, E 106, 147, 148 Aissaoui, D.M. 83, 84 Aitken, W.G. 376 AI Far, D.M. 56, 82 Alam, M. 98-100, 101 Ali, H.O. 115, 133, 134, 147 Alidou, S. 102 Allam, A.M. 61, 82 Allen, J.R.L. IX, XI, 152, 163, 165, 171, 197 Allen, P.A. IX, XI Alli Kassim, M. 232 Allix, P. 154, 171 Allsopp, H.J. 268 Allsopp, H.L. 258, 267 Almond, D.C. 106, 113, 147 Alzouma, K. 102 Amireh, B.S. 16 Anderson, A.M. 272, 276, 279, 280, 282, 313,315

Anderson, H.M. 257, 267 Anderson, J.M. 257, 267, 274, 277, 313 Anderson, R.E. 148 Andrew, G. 106, 123, 124, 132, 147 Anon. 115, 116, 118, 135, 147 Araujo, D. 277, 315 Ar6vian, A. 101 Aristova, K.Ye. 103 Arnauti, A. 85 Arruda, A.A. 26 Ashuri, O. 26 Aslanidis, P. 208 Austin, J.A. 186 Avbovbo, A.A. 89, 90, 101, 166, 171,172 Awad, G.H. 55, 82 Awad, G.M. 43, 45, 57, 61, 62, 65, 82

* Page references to text are in roman type, to bibliographic entries in italics.

Ayeed, M.A. 148 Ayoola, E.O. 101 Bachellen, W.D. 20, 26 Badenhorst, EP. 246, 267 Bailey D.K. 191,207 Baird, D.W. 21, 26 Baker, B.H. 113, 147, 191, 195, 207 Balducci, A. 20, 26 Ball, J. 137, 145, 147 Bally, A.W. 194, 207, 322, 374 Banerjee, S. 13, 16, 26 Barberi, E 195, 207 Barker and Associates 311,313 Barnes, S.U. 215, 217, 227, 230, 232 Barr, E.T. 36 Ban', E T. 28-30, 36, 80, 82 Barthel, K.W. 57, 71, 82 Bassett, R.L. 99, 102 Basson, W.A. 277, 313 Bate, K.J. 359, 374, 375 Baudet, J. 93-95, 101 Baumgartner, T.R. 174, 186 Bayoumi, A.I. 44, 82, 83 Beall, A.O. 49, 54, 83 Beamish, G.W.J. 322, 344, 374 Beard, L.S. 313 Beaumont, C. 89, 102, 323, 324, 374 Bebout, D.G. 29, 36 Behr, S.H. 290, 294, 313 Bellini, E. 10, 11, 16, 21, 25, 26, 83 Bellion, Y. 90, 91,102 Belmonte, Y. 176, 177, 186 Beltrand, M.D. 214, 215, 227, 232 Ben-Avraham, Z. 324, 374 Benelli, F. 26 Benkhelil, J. 102, 153-155, 171 Bennacef, A. 5, 16 Bennet, G. 329, 374 Benomran, O. 70, 83 Benson, J.M. 374 Bergggren, W.A. 157, 171 Bernasconi, A. 71, 83 Bernoulli, D. 84 Berry, L. 137, 147 Bertrand-Sarfati, J.D. 89, 102 Beuf, S. 10, 16 Beukes, N.J. 302, 304, 309, 313 Bezan, A. 83 Bhattacharyya, D.P. 49, 53-55, 60, 83, 85

Bigarella, J.J. 303, 313 Bignot, G. 85 Biju-Duval, B. 16 Binks, R.M. 31, 36 Bishop, W.E 68, 71, 83

Black, R. 91,102, 194, 195, 207 Blair, T.C. 197, 200, 207 Blarez, E. 232 Blignault, H.J. 316 Blignaut, J.J.G. 291,313 Blondeau, A. 85 Bonatti 191 Bonnefous, J. 67, 68, 83 Bosellin, A. 211,227, 232 Bosworth, W. 45, 83, 194-196, 207 Botha, B.J.V. 299, 300, 302, 303, 313, 316

Bottcher, R. 57, 59, 82, 83 Boudouresque, L. 99, 101,102 Bouju, J.P. 182, 186 Boukhary, M. 85 Bouma, A.H. 284, 285, 313 Boureau, E. 95, 102 Bourgeois, J. 313 Bowen, B.E. 196, 202, 207 Bowen, R. 4, 16, 54, 83 Bowin, C. 172 Bowitz, J. 84 Brady, T.J. 36 Braithwaite, C.J.R. 29, 36 Brakenridge, G.R. 313 Branson, J.C. 233 Braun, J. 323, 324, 374 Bredell, J.E. 310, 313 Brennan, P. 28, 36 Brenner, P. 370, 374 Brenner, P.W. 375 Briden, J.C. 85, 315, 341,376 Brink, A.H. 176, 177, 186 Brink, G.J. 322, 345, 357, 374 Bristow, J.W. 267 Broad, D.S. 320, 329, 341, 359, 360, 363-364, 369, 370, 374 Broderick, T.J. 314 Brognon, G. 182, 183, 186 Brown, F.H. 203, 207 Brown Jr., L.E 322, 374 Brown, L.E 184-186, 186 Browne, S.E. 113, 147 Bubb, J.N. 208, 376 Bullard, E.C. 152, 171, 173, 186 Burgeois, R.H. Jr. 278 Burger, A.J. 238, 267 Burger, C.A.J. 375 Burgis, M. 187, 207 Burke, K.C. 165, 166, 171 Burollet, P.F. 12, 16, 22, 23, 26 Buscaglione, L. 215, 232 Busch, D.A. 291,313 Bush, D.A. 159, 171 Busrewil, M.T. 35, 36

384 Bustin, R.M. 171,171 Butzer, K.A. 207 Butzer, K.W. 141, 144, 147 Byramjee, R. 22, 23, 26 Cadle, A.B. 290, 291,293, 294, 313 Cairncross, B. 289-291,293, 313 Campbell, N.D.H. 36 Cannon, R.T. 217, 218, 227, 232 Carmingnani, L. 232 Cartwright, J.A. 324, 374, 375 Carvalho, M.S. 93, 102 Carey, W. 379 Casanova, J. 188, 208 Caswell, P.V. 218, 232 Cerling, C.E. 203, 207 Cerling, T.E. 147, 149, 206, 207 Chaltellier, J. 212, 232 Chapman, D.S. 148 Chapman, G.R. 194, 207, 209 Chen, W.P. 194, 207 Chiarelli, A. 19, 26 Chorowitz, J. 193, 194, 207 Chowdhary, L.R. 46, 73, 76, 77, 83 Christie, A.D.M. 294, 301,302, 313, 316 Chukwura, P.I. 171, 172 Chumakov, I.S. 77, 81, 83 Cifelli, R. 172 Cita, M.B. 84 Clark-Lowes, D.D. 5, 16 Clarke-Lowes, D.D. 23, 26 Clauer, N. 252, 267 Clifford, A.C. IX, XI, 4, 16, 17, 21, 26, 31, 36, 152, 171, 173, 174, 177, 180, 181, 183, 186, 380, 382 Clifford, H.J. 28, 36 Clifford, T.N. 267 Clift, W.O. 232 Coal Commission 258, 267 Cococcetta, V. 85 Coe, R.A. 268 Coffin, M.E 232 Cohen, A. 203, 207 Cole, D.I. 275-277, 291, 293, 294, 297, 305, 310, 312, 313, 316 Colley, B.B. 24, 26, 36 Collinson, J.W. 308, 313 Colliston, W.P. 317 Collomb, G.R. 12, 16, 22, 23, 26 Combaz, A. 186 Compston, W. 148 Coniglio, M. 75, 76, 83, 84 Cooke, H.B.S. 144, 147 Cooper, J.A.J. 316 Coppens, Y. 187, 207 Corbet, T. 136, 149 Cordry, E.A. 152, 172 Cox, K.G. 314, 315 Cramez, C. 186 Curry, J.J. 37 Curtis, P.C. 193, 195, 208 Curtis, R.L. 277, 315 D'Amico, C. 211,232 d'Hoore, J.L. 263, 267 Dakkak, M.W. 48, 49, 51, 53, 83 Dakshe, A. 83 Dalziel, I.W.D. 380, 382

A U T H O R INDEX Damotte, R. 102 Darracott, B.W. 113, 148 Darwin, C. IX Darwish, M. 55, 82, 83 Darwish, Y.A. 83 Daukoru, E.M. 152, 162, 164, 172 Davies, C.P.N. 344, 345,374 Davies, D. 208 Dawoud, A.S. 148 De Beer, J.H. 269, 305, 313 De Charpal, O. 16 De Heinzelin, J. 80, 81, 83, 141,147 De Lapparent, A.E 96, 102 De Ruiter, P.A.C. 173, 184, 185, 186 De Swardt, A.M.J. 310, 315, 322-324, 329, 374 De Wit, M.J. 305, 307, 313 Degens, E.T. 201,208 Deibis, S. 58, 82 Delteil, J.R. 153, 155, 171, 186 Demaison, G. 186 Deroo, G. 16, 26 Deynoux, M. 17, 26 Dickinson, W.R. 281,283, 305, 313, 380, 382

Dikouma, M.S. 89-92, 94, 96, 98, 102 Dingle, R.V. 152, 153, 172, 305, 309, 313, 322-324, 329, 348, 370, 374 Dixon, G.E. 348, 376 Doherty, S. 374 Doornkamp, J.C. 128, 143, 147 Dott Jr., R.H. 278, 313 Doughri, K. 26 Dreimanis, A. 314 Drezet, H. 101 Druckman, Y. 55, 83 Du Plessis, A. 324, 375, 376 Du Toit, A.L. 277, 288, 305, 313, 379 Du Toit, S.R. 322, 324, 374 Dualeh, A. 215, 227, 232 Dubois, D. 101,102 Duncan, A.L. 267 Duncan, A.R. 314 Duncan, R.A. 267, 314 Dunn, L.G. 49, 49, 53, 60, 83 Duval, B. 182, 186 Eales, H.V. 267, 303, 304, 314 Edwards, J.D. 374 Ejedawe, J.E. 169, 171,172 Ekweozor, C.W. 171,172 El Adindani, A. 82 E1Ageed, A.I. 106, 147 E1 Amin, A.M. 147 E1 Amin, A.S. 134, 147 E1 Aref M. 82 E1Badri, O. 124, 147 E1 Badry, O. 58, 61, 64, 65, 85 E1 Baz, F. 39, 83 El Boushi, I.M. 124, 132, 133, 140, 147, 148

E1 Fetouh, M.A. 53, 85 E1 Gaby, S. 40, 83 El Gezeery, M.N. 43, 48, 58, 61, 70, 83 E1 Gezeery, N. 43, 45, 83 E1 Haddad, A. 75, 76, 83 E1 Hawat, A.S. 30, 36, 43, 56, 58-61, 68, 71, 72, 79, 80, 83

E1Rabaa, S.M. 106, 147, 148 E1 Shazly, E.M. 39, 40, 60, 67, 73, 77, 84 E1 Shinnawy, M.A. 62, 84 E1Tohami, M.S. 130, 135, 148 E1 Zarka, M.H. 45, 58, 61, 62, 70-73, 84 E1-Arnauti, A. 80, 81, 83 EI-Nakhal, H.A. 172 El-Worfalli, H.O. 68, 83 Elliot, D.H. 308, 314 Elliot, T. 282, 314 Emery, K.O. 152, 153, 172 Erickson, A. 84 Erikson, K.A. 376 Eriksson, K.A. 258, 267, 291,314, 316 Eriksson, P.G. 302, 314 Erjavec, J.L. 313 Erlank, A.J. 267 Ernesto, M. 268 Espitali6, J. 16, 26, 323, 375 Eugster, H.P. 148, 188, 198, 207 Evamy, B.D. 152, 166, 168, 169, 172 Evans, R. 174, 186 Everett, J.E. 171, 186 Fabre, J. 102 Fahmi, N. 82 Fairhead, J.D. 31, 36, 113, 147, 190-193, 206, 207 Falcon, R.M.S. 258, 267, 310, 314 Falvey, D.A. 232, 323, 324, 375 Fantozzi, P. 232 Farid, M. 83 Fastovsky, D.E. 93, 102 Fatti, J.L. 329, 375 Faure, H. 94, 102, 121, 148 Fayose, E.A. 171 Fazzuoli, M. 215, 232 Fediuk, E 35, 37 Ferguson, D.S. 207 Ferguson, R.C. 313 Fitch, F.J. 314 Fleck. R.J. 106, 148 Fondeur, C. 171 Forsythe, R. 305, 314 Fouch6, J. 323, 329, 347, 375, 376 Francis, P.W. 112, 148 Frankl, E.J. 152, 172 Fredricks, P.E. 99, 102 Frets, D.C. 256, 258, 267 Freyer, E.E. 268 Friedinger, P.J.J. 359, 375 Frostick, L.E. 1 8 7 , 189, 190, 194, 196-198, 200, 201,203-206, 207 Fullagar, P.D. 4, 16 Furon, R. 89, 102, 259-261,267 Furter, EJ.J. 313 Galloway, W.E. 159, 172 Garcia-Victoria, E. 375 Garea, B.B. 29, 36 Gariel, O. 16 Garrison, R.E. 65, 84 Gass, I.G. 106, 111,148 Gasse, F. 140, 141,144-146, 147, 148 Gaudet, J.J. 187, 207 Gawthorpe, R.L. 194, 207 Genik, C.J. 89, 90, 90, 102 Gentle, R.1. 329, 374

AUTHOR INDEX Geological Map of South West Africa/ Namibia 240, 267 Geological Survey 259, 262, 267 Geophysics and Strojoexport 114-116, 130, 148 Gerrard, I. 322, 375 Gevers, T.W. 246, 267 Ghuma, M.A. 4, 16 Gibbs, A.D. 193, 194, 207 Gilbert, C.E. 347, 375 Gillespie, R. 147 Giori, I. 26 Girdler, R.W. 113, 148 Glen, J.M. 268 Glenn, C.R. 65, 84 Gold Fields 260, 267 Goodlad, S.W. 375 Goudarzi, G. 24, 26 Goudarzi, G.H. 22, 24, 26, 32, 36 Goudie, A.S. 207 Gough, D.I. 313 Gouws, H.J. 357, 375 Govean, EM. 45, 85 Gram, P.H. 207 Gravenor, C.P. 274, 314 Green, C.M. 190, 207 Greenwood, W.R. 148 Greigert, J. 89, 91, 92, 95-101,102 Grobler, N.J. 250, 267 Groenewald, G.H. 295, 296, 299-301,314 Groschel-Becker, H. 186 Grove, A.T. 187, 207 Grund, R. 36 Guerrak, S. 12, 16 Guiraud, R. 43, 84, 89, 90, 95, 102 Guj, P. 256, 267 Gumati, Y.D. 31, 33, 36 Gunn, R.H. 119, 123, 128, 139, 148 Giinzel, A. 267 Haberyan, K.A. 187, 208 Hadley, D.G. 148 H~ilbich, I.W. 268, 305, 309, 314 Halfman, J.D. 208 Hall, B. IX, XI Halstead, L.B. 98, 102 Hampton, M.A. 285, 315 Hamyouni, E. 29, 36 Handford, C.R. 99, 102 Hankel, O. 221,232 Hansen, C.L. 141,147 Hantar, G. 57, 58, 60, 66, 84 Haq, B.U. 159, 162, 167, 172, 375 Hardenbol, J. 172, 375 Haremboure, J. 172 Harms, J.C. 75, 76, 80, 84 Harms, J.G. 81, 84 Hartley, R.W. IX, XI Hartnady, C.J.H. 323, 324, 374, 375 Hasel, J.E.P. 232, 232 Hashad, A.H. 40, 84 Hassan, A. 149 Hassan, EA. 137, 148 Hassan, M.A. 84 Hastings, D.A. 103 Hatlelid, W.G. 208, 376 Haughton, S.H. 218, 232, 259-261, 267, 282, 314

385 Hea, J.P. 28, 36 Heath, D.C. 260, 267 Hecky, R.E. 187, 201,208 Hedberg, R.M. 238, 240, 242, 244, 245, 247, 248, 250-258, 262, 267 Helba, A.A. 82 Hempstead, N. 180, 186 Henderson, N.B. 193, 207 Hendriks, E 57-60, 62, 65, 70, 84, 85 Henry, B. 208 Henry, G. 246, 267 Hepworth, J.V. 113, 148 Herman-Degen, W. 71, 82 Hermina, M. 64, 71, 84 Higham, M. 208 Hill, S.J. 345, 347, 375 Hiller, N. 163, 172, 299, 300, 309, 314 Hirtz, P. 186 Hobday, D.K. IX, XI, 288, 290, 292-294, 298, 300, 305, 313, 314, 316, 317, 376 Hodges, K. 344, 345, 347, 375 Hodgson, ED.I. 260, 267 Hollister, C.D. 157, 171 Holmes, A. 121,148 Holmes, L.C. 375, 376 Hooper, P.L. 267 Hooper, P.R. 303, 314 Horsthemke, E. 258, 267 Hospers, J. 169, 172 Houbolt, J.J.H.C. 165, 172 Houesson, A. 102 Howell, EC. 207 Hsti, K.H. 79-81, 84 Hsti, K.J. 76, 77, 80, 84, 380, 381,382 Hubler, S.L. 207 Hughes, M.J. 268 Hugo, P.J. 238, 255, 257, 259, 260, 262-264, 267 Hunt, J.A. 232 Hunter, D.R. 316, 376 Hunting Technical Services 130, 133, 134, 148 Hurley, A.M. 85, 315 Hurst, H.E. 124, 137, 148 Hurst, J.M. 194, 207 Hussein, M.T. 148 Hutchinson, D.R. 16 Ibrahim, H. 232 Ibrahim, M.W.I. 33, 36 Inman, K.E 313 Innes, J. and 267 Isaac, G.L. 147, 207 Iskander 135 Issawi, B. 48, 84 Ivannikov, A. 102 Jackson J. 194, 208 Jackson, M.P.A. 186, 316, 376 James, N.E 75, 76, 83, 84 Jaque, M. 21, 26 Jaujon, M. 102 Jeletzky, J.A. 323, 375 Jenkins D. 55-57, 60, 62, 63, 84 Johnson, B.D. 376 Johnson, G.D. 203, 208 Johnson, M.R. 269, 278, 281-283, 295-303, 305-308, 314

Johnson, T.C. 193, 208 Johnstone, D.W. 232 Jollands, A. 374 Jonathan, D. 26 Jones, B. 97, 102 Jones, B.F. 148 Jones, D.L. 323, 375 Jones, W.B. 103 Jonker, J.P. 277, 315 Jordaan, M.J. 284, 287, 314, 316 Jungslager, E.H.A. 374 Jux, U. 4, 16, 48, 54, 83, 84 Kaaya, C.Z. 208 Kabesh, M.L. 106, 148 Kajato, H.K. 223-226, 230, 231,232 Kallenback, H. 58, 59, 84 Kamen-Kaye 230 Kamerling, P. 172 Kanes, W.H. 31, 36 Karanjani, F.M. 232, 232 Karcz, I. 55, 84 Karkanis, B.G.Y. 111,148 Karson, J.A. 193, 195, 208 Keeley, M.L. 41, 43, 49, 50, 84 Keenan, J.H.G. 374 Keheila, E.A. 72, 85 Keller, G. 99, 102 Kendall, R.L. 137, 146, 148 Kent, P.E. 222, 225, 227, 232, 323, 375 Khalifa, M.A.G. 72, 85 Khalil, B. 16 Khalil, M.H. 45, 84 Khalil, N.A. 48, 49, 84 Khan, A. 193, 208 Kheiralla, M.K. 109, 111, 124, 132, 148 Kidd, R.B. 84 Kilian, C. IX, XI, 4, 16, 101,102 King, B.C. 194, 208 Kingsley, C.S. 278, 280-282, 297, 314 Klemme, H.D. 211,232 Klerkx, J. 40, 84 Klinger, H.C. 373, 375 Klitgord, K.D. 4, 16, 21 Klitzsch, E. 10-14, 16, 21, 22, 25, 26, 39, 41-43, 47-57, 59, 60, 62-64, 66, 67, 73, 74, 84, 85, 106, 109, 148 Klotchko, V. 102 Knapp, W.A. 172 Knepp, R.A. 313 Knobel, J. 101 Knox, G.J. 166, 172 Kogbe, C.A. 15, 16, 89, 94, 95, 97-101, 102

Kostandi, A.B. 53, 53, 61, 84 Krasheninnikov, V.A. 99, 102 Kreuser, T. 191,208, 230, 232 Kr6ner, A. 106, 113, 148, 238, 242, 252, 267

KrUger, T.L. 243, 249, 267 Kuenen, P.H. 279, 314 Kulbicki, G. 16 Kuznir, N.J. 206, 208 Labuschagne, L.S. 310, 313 Ladd, J. 309, 314 Lambiase, J.J. 191,208 Lang, J. 101,102

386

AUTHOR INDEX

Langbein, W.B. 141,148 Laronne, J.B. 197, 208 Larsen, P.H. 194, 208 Larson, R.L. 309, 314 Lawson, A.C. 137, 148 Le Blanc Smith, G. 258, 267, 289-291, 293, 294, 314 Le Roex, A.P. 268 Le Roux, J.P. 299, 310, 312, 314, 316 Le Theoff, B. 102 Leakey, M.D. 208 Leakey, M.G. 203, 208 Leakey, R.E.F. 187, 203, 207, 208 Ledendecker, S. 258, 267 Legrand, P.H. 4, 16, 18, 26 Lehacek, T. 267 Lehner, P. 173, 184, 185, 186 Leith, M.J. 314, 322, 375 Lejal-Nicol, A. 49, 84 Lemmer, W.M. 284, 285, 314 Lemoigne Y. 95, 102 Lewy, Z. 63, 84 Light, M.P.R. 322, 329, 344, 375 Lindberg, F.A. 313 Linstrtim, W. 295, 299, 300, 313, 314 Linton, R.E. 218, 233 Lippard, S.J. 207 List, F.K. 83, 84 Lister, G.S. 208 Livingstone, D.A. 137, 143, 146, 148, 193, 208 Lock, B.E. 282, 303, 305, 314, 373, 375 Logan, C.T. 267 Logar, J.F. 173, 175-177, 180, 181,186 Lombaard, A.F. 256, 267 Long, R.E. 193, 208 Loock, J.C. 284, 287, 288, 308, 316, 317 Lorenz, J. 22, 26 Lofty, H.L. 83 Loule, J.P. 172 Loupekine, I.S. 198, 208 Loutit, T.S. 172 Lowe, D.R. 314 Luger, P. 65, 84 Lundin, E.R. 183, 186 M'Rabet, A. 85 Mabrook, B.M. 133, 148 Macdonald, M. 148 Macdonald, R. 113, 148, 195, 208 MacLeod, I.N. 232 Maggliore, P.R. 26 Maguire, P.K.H. 193, 208 Maher, C.E. 36 Maiden, K.J. 267 Maier, J.J. 355, 375 Maisey, J.G. 93, 102 Makhlouf, I.M. 16 Malan, J.A. 324, 348, 359, 361-362, 369, 374, 375

Malmberg, G.T. 132, 133, 140, 141,148 Maluski, H. 186 Mamgain, V.D. 37 Manderscheid, G. 12, 16 Mansour, S.E. 83-85 Marin, H. 267 Marinho, M. 232 Markwort, S. 208

Marot, J.E.B. 322, 344, 375 Marsh, J.S. 303, 314 Marsh, T.S. 267 Martin, A.K. 323, 324, 375 Martin, H. 238, 243, 246, 256, 258, 267 Martini, J.E.J. 281,296, 308, 314 Martyn, J.E. 207 Marzouk, I. 84 Mascle, J. 171, 172, 211,227, 232 Mason, T.R. 289, 314, 316 Massa, D. 10, 11, 16, 21, 26 Masson, P. 186 Matheis, G. 89, 102 Mathew, D. 288, 315 Maud, R.R. 329, 375 Maurin, J-C. 43, 84 Mbede, E.I. 230, 231,232 McBride, E.F. 302, 315 McDonald, K.C. 60, 62-64, 85 McDougall, I. 124, 128, 141,148 McGarva, A.M. 49, 84 McGregor, D.S. 20, 26 McKee, E.D. 22, 23, 26, 109, 148 McKenzie, D.P. 191, 192, 194, 208 McKinlay, A.C.M. 376 McLachlan, I.R. 184, 185, 186, 272, 275-277, 313, 315, 322-324, 329, 344, 348, 357, 360, 369, 370, 374,

Morgan, P. 40, 47, 53, 69, 84 Morley, C.K. 194, 208 Morton, W.H. 148 Mougenot, D. 232 Moussine Pouchkine, A. 102 Moustafa, A.R. 43, 45, 82, 84, 85 Mula, A.H.G. 148 Muller, C. 84 Munn, S.G. 208 Muntingh, A. 184-186, 186, 374 Murat, R.C. 154, 172 Musrati, H. 36 Mutti, E. 345, 375 N.C.R. 139, 148 Nagati, M. 231,232 Nairn, A.E.M. 31, 33, 36, 83, 232 Nairn, E.M. 84 Neev, D. 41, 84, 85 Nel, L. 315 Nelson, R.A. 85 Netherwood, R.E. 45, 85 Netterberg, F. 263, 268 Newton, A.R. 313, 374 Normark, W.R. 345, 375 Norton, I.O. 323, 375 Nwachukwu, J.I. 171,172 Nwajide, C.S. 154, 172

375

McMillan, I.K. 322, 323, 329, 347, 348, 358, 360, 369, 370, 374, 375 Me'hes, K. 102 Medani, A.H. 116, 148 Megerisi, M.F. 37, 83 Meister, E.M. 24, 26 Melieres, F. 84 Memeisy, M.Y. 56, 62, 64, 68, 69, 84 Merabet, O. 95-98 Merki, P. 169, 172 Meyer, A. 186 Meyers, J.B. 179, 186 Miall, A.D. 94, 102, 201,208 Michum, R.M. 85 Middleton, G.V. 285, 315 Milad, G. 85 Miller, J.A. 314 Miller, R. McG. 238, 239, 242-244, 246, 247, 252, 256-258, 260, 267, 268 Milner, S.C. 268 Minter, W.E.L. 316, 376 Missallati, A. 83 Mitchell, A.H.G. 308, 309, 315 Mitchell, C. 315 Mitchum, R.M. 172, 208 Mitchum Jr., R.M. 376 Mohamed, A.M. 116, 135, 148 Mohamed, I.I. 147 Mohr, P. 195, 208 Mohsen, S.M. 83 Molinas, E. 101 Molloy, F.A. 172 Molnar, P. 194, 207, 208 Momper, J.A. 268 Mongin, D. 96, 102 Monie, P. 186 Montadert, L. 84, 171 Moody, R.T.J. 93, 94, 96-100, 102, 103 Moreau, C. 102

O'Connor, J.M. 268 Obelliane, J.M. 101 Oelofson, B.W. 277, 315 Oertli, H.J. 374 Okoh, S.U. 171,172 Okoye, N.V. 171,172 Oloviera, M.A.M. 26 Omara, S.M. 76, 80, 85 Omatsola, M.E. 166, 172 Orife, J.M. 164, 171,172 Ortiz, E.F. 26 Osahon, G.A. 101 Otto, C.J. 136, 149 Ousmane, B. 102 Oxburgh, E.R. 191,208 Pacca, I.G. 268 Pagel, M. 186 Park, R.G. 206, 208 Parker, D.H. 97-99, 103 Parrish, J.T. 277, 315 Parsons, M.G. 29, 37 Partridge, T.C. 329, 375 Patriat, P. 171 Patton, T.L. 45, 85, 208 Paverd, A.L. 314 Pavoni, N. 31, 37 Peacock, D.C.P. 194, 208 Peck, N. 83 Pellaton, C. 102 Pendexter, C. 29, 36 Perrin, M. 268 Pesce, A. 35, 37 Peterson, R.M. 20, 26 Petrosyants, M.A. 103 Petters, S.W. 99, 103, 154-156, 159, 166, 167, 171, 172 Pferdek~imper, H.W. 344, 375 Phillips, J. 172

AUTHOR INDEX Phillips, P. 124, 137, 148 Philobbos, E.R. 45, 71, 85 Pierobon, E.S.T. 21, 26 Pinheiro, H.G. 267 Plafker, G. 323,375 Playfair, K.J. IX, XI Pliny IX, XI Poliani, G. 83 Pollack, H.N. 148 Pomeyrol, R. 67, 85, 109, 148 Pommier, G. 20, 26 Popoff, M. 154, 171 Porada, H. 267 Posamentier, H.W. 172, 208, 376 Potgieter, G.J.A. 284, 315 Pougnet R. 89, 91, 92, 95-97, 101,102 Powell, C.M. 308, 316 Powell, C.McA. 323, 376 Powers, P.W. 206, 207 Pr6vot, M. 268 Purser, B.H. 45, 77, 78, 80, 84, 85 Pyre, A. 214, 215, 222, 227, 232 Pyre, J.T.O. 229, 232 Quennell, A.M. 323,376 Questiaux, J.-M. 344, 376 Qureshi, I.R. 106, 148 R.E.G.W.A. 116, 133, 148 R.W.C. 149 Rabinowtz, P.D. 227, 228, 232 Radier, H. 89, 98, 103 Radwan, I.A. 70-73, 84 Rais-Assa, R. 217, 218, 232 Ramos, V.A. 308, 315 Rankama, K. 242, 267 Ranoszek, M. 375 Rathan, D.G.L. 376 Rayner, R.J. 316 Reading, H.G. 315 Reeves, C.V. 192, 207, 217, 232 Regnoult, J.M. 169, 172 Rehacek, J. 314 Reid, I. 187, 189, 190, 194, 196-198, 200, 201,203, 205-206, 207 Reijers, T.J.A. 167 Reischmann, T. 148 Rene, P.R. 258, 268 Rettig, S.L. 148 Reyment, R.A. 96, 98, 103, 153, 172 Rhodis, H.G. 123, 133, 134, 149 Richardson, J.L. 147 Richter A. 40, 85 Riek, E.F. 297, 315 Rigassi, D.A. 348, 370, 376 Rizzini, A. 73, 76-80, 85 Roberts, D.L. 288-291, 293, 294, 312, 313,315,316

Roberts, G.P. 194, 208 Roberts, J.M. 28, 37 Robson, D.A. 45, 85 Rogers, J.J.W. 4, 16, 193, 208 Rognon, P. 16, 148 Rona, P.A. 186 Rosendahl, B.R. 172, 186, 193, 194, 196, 201,208 Ross, D. 180, 186 Rossouw, P.J. 278, 315

387 Rowlands, P.H. 172 Rowsell, D.M. 310, 315, 322, 329, 374, 375

Russeger, J. 13, 16 Rust, I.C. 287, 303, 315, 348, 376 Rust, U. 263, 265, 268 Ruxton, B.P. 106, 124, 149 Ryan, P.J. 279, 281,287, 315, 316 Ryan, W.B.E 77, 84, 85 Ryberg, P.T. 313 SACS (South African Committee for Stratigraphy) 249, 260, 262, 268, 315 Sadig, A.A. 106, 148 Saeed, E.M. 116, 134, 149 Said, E.M. 149 Said, R. 39, 45, 48, 49, 70, 73, 76, 81, 85, 128, 141, 145, 149 Saint-Marc, P. 102 Salama, M.N. 123, 130, 132-134, 149 Salama, R.B. 106, 109, 111-115, 119, 120, 122-124, 129-135, 137-140, 146, 149 Salem, M.J. 79, 83 Salem, R. 69, 73, 75-77, 80, 85 Sanad, S. 76, 80, 85 Sanavely, P.D. 84 Sander, S. 193, 208 Sanderson, D.J. 194, 208 Sanford, R.M. 27, 28, 37 Sangree, J.B. 208, 376 Santacroce, R 207 Santogrossi, P.A. 322, 374 Sarg, J.E 172 Sassi, F.P. 232 Schamel, S. 31, 36 Schandelmeier, H. 40, 41, 45, 47, 51, 53-56, 65, 73, 85 Schermerhorn, L.J.G. 268 Schild, R. 141,149 Schlich, R. 227, 232 Schmidt, D.L. 148 Scholtz, D.L. 305, 315 Schoton, H. 16 Schreuder, S. 376 Schroeder, J.H. 84 Schull, T.J. 109, 110, 113-115, 118, 135, 136, 149 Schurmann, H.M.E. 41, 85 Sclater, J.G. 323, 375 Scott, R.W. 45, 85 Scrutton, R.A. 376 Searle, R.C. 193, 208 S6goufin, J. 323, 376 Seilacher, A. 6, 16, 48, 51, 85, 315 Selley, R.C. 3, 5, 16, 30, 33, 34, 37, 79, 80, 85 Sellwood, B.W. 45, 85 Seme Obomo, R. 169, 172 Semkiwa, P.M. 208 Semtner, E. 16 Sestini, G. 70, 85 S.G.E.P. 149 Shafie, A.I. 123, 149 Shanmugan, G. 279, 316 Shaw, J. 315 Shelmani, M.A. 43, 55, 56, 61, 68, 72, 83,85

Shepherd, E 267 Shone, R.W. 315, 370, 376 Short, K.C. 152, 166, 172, 174, 186 Shukri, N.M. 137, 149 Siebrits, L.B. 284, 287, 288, 315 Siesser, W.G. 313, 374 Sigal, J. 323, 375 S.I.K.R. 119, 120, 132, 149 Silva, G.A. 323, 376 Simiyu Siambi, W.M.N. 232 Simpson, E.S.W. 232 Sims, K.W. 207 Slevin, A. 232 Sly, P.G. 149 Smellie, J.L. 305, 315 Smith, A.G. 54, 68, 85, 171, 186, 308, 315, 341,376 Smith, A.M. 263, 265, 268, 293, 297, 315,316

Smith, G.C. 322, 375 Smith, J.P. 24, 26, 32, 36 Smith, R.H.M. 270, 315 Snavely, P.D. 71, 85 Snelson, S. 322, 374 Stihnge, P.G. 238, 242, 243, 246-250, 252, 268 Soliman, M.A. 83, 85 Soliman, S.M. 53, 58, 61, 64, 65, 85 SONATRACH 18, 20 Sougy, J. 26 Specht, T.D. 196, 208 Spence J. 225, 232 Squyres, C.H. 13, 14, 16, 25, 26, 39, 42, 43, 49, 54, 83, 84 Stacher, P. 166, 168, 171, 172 Stallbom, K.E. 375, 376 Stanistreet, I.G. 267, 314 Stauble, A.J. 152, 166, 172, 174, 186 Stavrakis, N. 277, 295, 299, 300, 309, 314,315

Stear, W.M. 296, 297, 315 Steel, R.J. 196, 198, 207 Stephens, J.M. 370, 376 Stem, R.J. 148 Stevaux, J. 16 Stewart, I. 208 Steyn, A.V.C. 376 Stow, D.A.V. 279, 316 Strauss, P.J. 329, 341,342, 344, 375, 376 Street, F.A. 140, 141,147, 148, 207 Strougo, A. 71, 71, 72, 85 Strydom, D. 316 Strydom, H.C. 277, 316 Stuart, G.W. 206, 207 Sultan, I.Z. 62, 84 Sutcliffe, P.J.C. 93, 94, 96-100, 102, 103 Sweeney, J.F. 89, 102 Taha, I. 43, 83 Taha, S. 46, 73, 76, 77, 83 Taher, M. 83 Talbot, M.R. 202, 208 Taleb, T. 83 Tallon, P.W.J. 196, 208 Tankard, A.J. IX, XI, 269, 270, 301,304, 305, 316, 322, 376 Taquet, P. 93, 94, 103

388

A U T H O R INDEX

Taverner-Smith, R. 289, 290, 293, 295, 314-316

Taylor, G.K. 315 Tehrani, R. 83 Tempere, C. 26 Temple, P.H. 128,, 143, 147 Terblanche, J.C. 284, 287, 288, 316 Terry, C.E. 30, 37 Theron, A.C. 279, 303,316 Theron, J.C. 297, 309, 313, 316 Theron, J.N. 163, 172, 296, 299, 316 Thiebaud, C.E. 45, 85 Thomas, D. 21, 24, 26, 27, 37 Thomas, D.S.G. 260, 268 Thomas, R.G. 103 Thompson III, J.B. 376 Thompson, S. 85 Thomson, J. 209 Thomson, S. 208 Thorpe, R.S. 148 Thunell, R.C. 71, 72, 82 Thusu, B. 85 Tiercelin, J.J. 197, 198, 208 Timonine, L. 102 Tinney, C.B. 310, 316 Tissot, B. 10, 16, 20, 26, 177, 185, 186 T.N.O. 133, 134, 149 Todd, R.G. 208, 376 Toens, P.D. 299, 310, 312, 314, 316 Toth, J. 136, 137, 149 Trichet, J. 102 Trofimov, D.M. 95, 99, 102, 103 Trompette, R. 26 Truckle, P.H. 187, 208 Truswell, J.F. 279, 316 Tsumeb Corporation 250, 268 Turk, T.M. 21, 26 Turner, B.R. 11, 16, 23, 25, 26, 297, 298, 300-302, 309, 316 Turner, J. 375 Turner, J.P. 180, 186 Turner, J.R. 355,376 Turner, J.R.T. 329, 341,374 Turpie, A. 233 Uchupi, E. 172 Unstead, P. 376 U.S. Naval Oceanographic Office 165, 172

Vail, J.R. 34, 37, 106, 107, 109, 112, 113, 116, 148, 149

Vail, P.R. 67, 68, 85, 172, 193, 201,208, 347, 375, 376 Valery, P. 171 Valicenti, V.H. 370, 376 Vallier, T.L 232 Van Andel, T.J.H. 174, 186 Van der Merwe, R. 329, 376, 375 Van der Spuy, A. 316 Van Der Westhuizen, W.A. 277, 316 Van Eeden, O.R. 277, 303, 313, 316 Van Houten, EB. 14, 16, 25, 26, 31, 37, 39, 56, 60, 62-64, 67, 68, 85 Van Vuuren, C.J. 288-291,293-296, 316 Van Wagoner, J.C. 159, 172, 322, 376 Van Wyk, N.J.S. 374 Van Zijl, J.S.V. 313 Varet, J. 207 Vavra, C.L. 313 Veevers, J.J. 308, 316, 376 Venter, J.J. 304, 317 Verrier, G. 182, 183, 186 Verweij, J.M. 134, 136, 137, 149 Vezzani, E 85 Viljoen, J.H.A. 278, 284, 308, 316, 324, 348, 375 Vincens, A. 188, 208 Virlogeux, P. 232 Visser, J.N.J. 103, 269, 270, 272, 274, 275, 277, 279, 284, 287, 288, 301-303, 305, 308, 309, 316, 317 Viterbo, I. 68, 85 Von Brunn, V. 272, 275, 305, 314, 317 Von Hohnel, L. 208 Vondra, C.F. 196, 202, 207 Vos, R.E. 290, 317 Wadsworth, M.J. 35, 36 Wahab, S.A. 72, 85 Wahdan, L. 149 Walgenwitz, F. 184, 186 Walker, B.R. 30, 36, 80, 82 Walker, R.G. 281,317 Walsh, J.J. 194, 208 Walter, R. 218, 233 Ward, J. 5, 16, 23, 26 Ward, W.C. 60, 62-64, 85 Warme, J.E. 380, 382 Washbourn-Kamau, C. 147 Watterson, J. 194, 208 Watts, D.R. 308, 314 Weber, K.J. 152, 163, 166, 172 Weeger, A.A. 28, 29, 36

Weissbrod, T. 53, 85, 109, 149 Welke, H.J. 252, 268 Wellington, J.H. 263, 268 Wendorf, E 141, 146, 149 Wenham, M. 375 Westermann, G.E.G. 218, 233 White, E.I. 98, 103 Whiteman, A.J. 16, 67, 85, 106, 111, 123, 124, 132, 137, 145, 147-149, 164, 172 Wickens, H. de V. 184, 185, 186, 278, 279, 283-287, 317, 316, 317, 344, 374,376

Widmier, J.M. 208, 376 Willcocks, N. 137, 145, 149 Williams, A.J. 149 Williams, F. 118, 141, 143, 147, 149 Williams, H.R. 103 Williams, J.J. 28, 30, 37 Williams, L.A.J. 194, 209 Williams, M.A.J. 118, 124, 128, 137, 140--146, 147-149 Willis, R. 376 Windley, B.E 89, 103 Winter, H. de la R. 304, 317, 322, 347, 370, 376 Winter, M.F. 290, 292, 317 Winters, S. 344, 345, 375 Wipplinger, P.E. 297, 312, 313 Wolfgang, S. 233 Woller, E 35, 37 Wonger, R. 186 Wood, J.M. 93, 103 Wopfner, H. 208 Wozny, E. 98, 102 Wray, J.L. 75, 76, 79-81, 84, 85 Wright, J.B. 89, 90, 103 Wright, L.D. 282, 317 Wright, R. 84 Wycisk, P. 47, 50-52, 54-57, 59, 63, 64, 66, 84, 85 Yitshak, Y. 208 Young, D. 84 Youssef, E.A.A. 80, 85 Youssef, M.I. 65, 85 Yuretich, R.F. 149 Zaborski, P.M.P. 98, 103 Zagaar, A.M. 37 Zak, I. 55, 84 Zawada, P.K. 284, 285, 317 Zawiskie, J.M. 313

Geographic, Tectonic and Stratigraphic Index

1Atl 323, 325 5Atl/6Atl 325 6Atl 323 13Atl 325 15Atl 325 Abakaliki subbasin 153 Abenab Subgroup 243, 244 Abinky Formation 93 Abong M'Bang 31 Abrahamskraal Formation 295 Abu Ballas Formation 58, 59, 67 Abu Gabra Formation 109, 135 Abu Gabra trough 121, 136 Abu Gharadig 43, 45, 57, 81 Abu Gharadig basin 45, 47, 48, 65 Abu Habil trough 129 Abu Madi Formation 77 Abu Qada Formation 62 Abu Thora Formation 49 Acacus Sandstone 10 Adelaide Subgroup 295, 309 Ader Doutchi Series 101 Adigrat Formation 215 Afikpo subbasin 153 African Shield 176 Agadez Group 92 Agala Sandstones 154 Agbada facies 163 Agbada Formation 156, 163, 166, 168 Agbani Sandstones 156 Agedabia (Marsa Brega) rifts 31 Agulhas Arch 329 Agulhas Bank 329 Agulhas-Falkland fracture zone 309, 319 Ahnet 17, 19, 20 Air 92 Air Massif 89, 90 Ajali Formation 159 Akata 163 Akata facies 163 Akata Formation 156, 166 AI Alamain Formation 58 AI Faiyum basin 71 AI Jabal al Akhdar 61, 71 AI Jabal al Akhdar trough 43 Alamain Formation 67 Albert rift 143 Alcarcesse Member 92 Algoa 319 Algoa Group 374 Algoa Basin 358 Allocyclic events 159 Amal Formation 28, 109, 114, 135 Ameki Formation 167

Ameki "group" 156 Amezroun Formation 94 Anambra Basin 156, 159 Anambra subbasin 153 Andoni 241 Andoni Formation 261,266 Angola basin 182,~"185 Angola/Cuanza basin 180 Anole Formation 215 Anrado Formation 217 Anyeli Formation 94 Anza Graben 217 Aouinet Ouenine 25 Aouinet Ouenine Formation 12, 21 Aptian evaporites 177, 180 Araba Formation 48, 49 Arabia 4, 5, 113 Arabian plate 73 Arabian shield 67 Arabo-Nubian 25 Arida Formation 30 Aroma graben 116 Assedjefer 25 Assouas Formation 93 Asu River Group 154 Aswan 39 Atacorian Series 91 Ataqa Formation 53 Atbara graben 129 Atlas Fold Belt 3 Atlas Mountains 12, 13, 17, 20 Augila Formation 30 Auros 241 Auros Formation 243, 245, 266 Awe Formation 154 Azile Formation ?? Babanusa trough 115 Back-arc basins V Baggara basin 135 Baggara graben 115 Bahi Formation 29 Bahr El Arab rift 114, 118, 120, 128, 130, 146 Baidabo Formation 215 Balfour Formation 295 Bamboesberg Member 301 Bara basin 133 Bara trough 115 Baraka Formation 111 Bassin de L'Oued Azazouk 89 Batanga Formation 177 Bazuzi 33 Beaufort Group 260, 283, 295, 329 Bechar basin 12, 13

Beiseb 241 Beiseb Formation 261,263, 266 Beletutuen Formation 217 Belmont Formation 299 Benin facies 163 Benin Formation 166 Benin Hinge Line 153 Bentiu Formation 109, 111, 114, 135 Benue Rift 96 Benue Trough XI, 89, 153, 154, 156 Berere Member 95 Berg Aukas 241 Berg Aukas Formation 266 Bilma graben 89 Bima Sandstone 154, 174 Binem Formation 25 Binga Limestone 184 Blue Nile rift 116, 129, 132, 146 Blue Nile rift basin 124 Bogma area 53 Bokkeveld 359 Bokkeveld Group 163, 329 Bombouaka Sandstones 91 Borak Formation 95 Brava Formation 214 Brazil 93 Break-up 323 Breaks, 4th or 5th order 162 Bredasdorp 319 Britskraal Shale Member 282 Bucomazi Formation 180, 182 Buen Series 91 Burgersdorp Formation 299, 309 Busul Member 215 Cabo Ledo ridge 183 Cainozoic Rift System 146 Calabar Flank 153 Calabar Hinge Line 153 Calanscio arch 25 Calcaires Blancs 97 Cameroon 173 Cameroon Volcanic Line 169, 173, 174, 183 Campos basin of Brazil 176 Canyons 162, 164 Cape Arc 381 Cape Fold Belt 269, 305, 359, 381 Cape Supergroup 269, 305, 319, 329, 359 Carlsberg Ridge 227 Catumbela Limestone 184 Catumbela Limestone Formation 183 Central Swamp 169 Chad 90 Chad Basin 89

390 Chain fracture zone 153 Cheffadene Formation 96 Chela Formation 181, 185 Chieun Formation 14, 25 Chuos 241 Chuos Formation 238, 245 Clarence Formation 258 Clarens Formation 302, 309 Clocolan Dome 309 Coal 53, 56, 310 Coastal Swamp depobelts 169 Coccolith Formation 68 Cocobeach Group 176 Colchester 348 Colchester Member 323, 3609 Collapsed back-arc basin 380 Collingham Formation 278, 308 Condensate 319 Condensate field 170 Congo basin 173, 185 Continental Intercalaire 13, 19, 92, 95 Continental Mesozoic 13, 14, 21 Continental Post-Tassilian 13, 25 Continental Terminal 101 Cross River Delta 152, 168, 169 Cuanza (or Kwanza) basin 182 Cuanza basin 173 Cuanza River 182 Cuvo Formation 182 Cyrenaica 68, 80 Cyrenaica platform 56, 67 Dabla Series 93 Daggaboursnek Member 296 Dahomeyian Series 91 Dakhla 40, 57, 60, 65, 81 Dakhla Basin 47, 48, 57-60, 65, 66, 68 Dakhla Formation 65 Dakhla Shale Formation 70 Damara Sequence 238 Dange Formation 98, 99 Darfur Group 109, 111, 135 Davie fracture zone 211 Deleb Member 215 Delta 80 Dembaba 25 Depobelts 166 Depocentres 166 Desouqy 50 Dhiffah Formation 50 Diapiric structures 170 Dibba Formation 30 Dinder trough 125 Djado 21 Djado Basin 95 Dolerite intrusions 304 Dondo Group 182 Doula Basin 169, 173, 174 Drakensberg Group 303, 309 Driekoppen Formation 300 Drift cells 157, 165 Drift onset 323 Drifting 319 Dukamaje 97 Dukamaje Formation 98 Duruma Series 217, 218 Duwi Formation 65 Dwyka 52, 54, 241,258

GEOGRAPHIC, TECTONIC AND STRATIGRAPHIC I N D E X Dwyka basin 308 Dwyka Formation 256, 257, 260, 266 Dwyka Group 270, 274, 329, 381 Early Cretaceous 323 East African Rift 187, 221, 381 Eastern Cape basin 184 Ecca 259 Ecca Group 260, 275, 329 Ekismane Formation 95 El Arab Formation 58 Elandshoek 241 Elandshoek Formation 248, 249, 251,266 Elandsvlei Formation 272 Elizabeth Trough 358 Elliot Formation 302, 309 En Nassame 95 Enon Conglomerate 323 Enon Formation 360 Equatorial Guinea 173 Escalator regression 166 Esna Shale 45 Esna Shale Formation 70, 71 Estcourt Formation 295 Etjo 241 Etjo Formation 258, 260, 266 Etosha Pan 262 Eurasia 55, 60, 68, 71, 81, 82 Eustatic sea-level changes 167 Evaporites 12, 19 Faiyum basin 45, 65 Falaise de Tiguedi 94 Falkland 324 Falkland Islands 308, 309 Falkland Plateau 323 Falkland Plateau Basin 323 Farak Formation 96 Ferfer Formation 217 Flood basalts 303 Foreland V Foreland basin 305 Fort Brown Formation 281,308 Freretown limestone 218 Gabon basin 173, 185 Gabredarre Formation 217 Gaghboub 43 Gamba Formation 99, 177, 181, 185 Gamtoos 319 Gao-Ansongo Trough 91, 98 Garadoua Formation 100 Garbaharre Formation 215 Gargaf arch 12, 21, 22 Gas 310, 319 Gauss 241 Gauss Formation 243, 245, 266 Gazal Formation 111 Genetic sequences 159 Ghadames 20, 21 Gharadig 62 Gharandal Group 45 Gharian 34 Gialo 29 Gialo Limestone Formation 30 Gilf Kebir Formation 59 Gindy basin 70, 71 Gir Formation 30

Glacial 51, 54 Glaciation 7, 10, 53, 54, 81 Gondwana IX, 269, 305, 319 Gondwanaland IX, 15, 211, 221,227, 303 Gondwanide Orogen 305 Gongola subbasin 153 Goufat Series 92 Greater Ughelli depobelt 169 Gr6s a Tigillites 92 Gr6s Argileux du Moyen Niger 101 Gr6s d'In Azaoua 92 Gr6s de Timesgueur 92 Growth fault-related traps 171 Growth faults 166, 169 Gruis Dolomite Member 248 Guinea Current 165 Gulf of Suez 39, 45, 53-55, 70-73, 75 Gulf of Suez basin 61, 62, 76, 82 Gumbro Series 217 Gumburo Group 214 Gundumi Formation 95 Gwandu Formation 99, 101 Hagfa 31 Hagfa Shale Formation 29 Hammada basins 25 Haouaz 25 Haruj al Aswad 34 Hassaouna Formation 22, 25 Hercynian IX, 20, 43, 51, 55, 66 Hercynian conformity 44 Hercynian Orogeny 12, 19, 53, 54 High Atlas 380 Hofra Formation 28 Hoggar 54 Hoggar Massif 20 Hon 31 Hon Graben 34 Htittenberg 241 Htittenberg Formation 249, 250, 266 Hydrocarbon traps 170 Hydrocarbons 319 Hydrogeological closed basins 146 Iabe Group 181 Idnwe Sandstone Member 301 Igdaman Group 98 Igoda Formation 373 Ikang Trough 153 Ilhas Formation 93 Illizi 21 Illizi/Ghadames 17, 20 I11o Formation 95 I11o Group 97 Imo Shale 156, 166 In Jinjira Formation 100 Incised valley 162 India 227 Infanta Embayment 319 Infra-Cambrian 4, 21 Irhazer Group 92, 93 Inhuca Formation 181 Ironstones 12 Ituk High 153 Izegouandane Formation 92 Jabal 39 Jabal Maghara 56

391

GEOGRAPHIC, TECTONIC AND STRATIGRAPHIC INDEX Jabal Uweinat 40, 42, 49, 54, 64, 81 Jebel Eghei 35 Jebel Kamil 40 Jesomma Formation 217 Jordan 73 Judith River Formation 93 Kafra Graben 95 Kalahari Sequence 238, 260, 263, 266 Kalambaina Formation 99, 100, 156 Kalkrand Basalt Formation 256 Kalkrand Formation 241,266 Kambe Formation 218 Karkar Formation 217 Karoo IX, 43, 191,238, 329 Karoo Basin 269, 305, 380 Karoo Sequence 256, 257 Karoo Supergroup 269, 329, 381 Karoo Trough 269, 305 Karroo 218, 221,227, 230, 231 Karroo Series 215 Katberg Formation 299 Keana Formation 154 Khalifa Formation 29 Kibiongoni Beds 218 Kirkwood Formation 355, 360 Kiseiba Formation 65 Kombat 241 Kombat Formation 252, 253, 266 Kookfontein Formation 285 Koonap Formation 295 Kufra IX, 17, 31 Kufra basin 14, 24, 25, 40, 42 Kurzuk 31 Laingsburg Formation 279, 308 Lambarene horst 177 Landana Formation 182 Late Jurassic 323 Le Accident sud-Atlasian 20 Libya 4, 20 Lidam Formation 29 Lindi Rift Basin 225 Lithostratigraphy 166 Loeme Formation 181 Logbaba Formation 174 Lower Cuvo Formation 182 Luculla Formation 180, 182 Madagascar 217, 218, 227, 323 Madiela Formation 177 Madingo Group 182 Maghrabi Formation 62 Maieberg 241 Maieberg Formation 246, 247, 251,266 Majiya Chumvi Formation 218 Major ground-water basins 105 Majunga Basin 323 Makurdi Sandstones 154 Mamu Formation 98, 159 Mamu-Ajali Sandstones 156 Manzeras Sandstone Formation 218 Marada 31 Marada Formation 30 Mariakani Formation 218 Marmarica Formation 75 Marsa Brega trough 30 Masajid Formation 57

Matolani Formation 218 Matruh 57, 62 Matruh basin 43 Mavuma Formation 181 Maximum flooding surfaces 162 Mayaputi Member 302 Mbizane Formation 270 Mediterranean Ridge 380 Melez 25 Melez Chogranne Shale Formation 22 Memouniat 25 Memouniat (Ordovician) Sandstone 24 Memouniat Sandstone Formation 22 Mengo Formation 180 Merca Formation 214 Messak Sandstone 21 Messinian 82 Metlaoui Group 71 Middleton Formation 295 Moghra Formation 73 Molteno Formation 260, 301,309 Molybdenum 310 Moordenaars Member 311 Mossel Bay 329 Mossgas 319 Mourizidie Formation 21 Mousseden Formation 92 Mouydir 17, 20 Mozambique 323 Mozambique Basin 323 Mozambique Channel 323 Mto Mkuu Formation 218 Muiden Group 252 Mulden Group 238, 266 Mundek Formation 174 Mungo Formation 174 Murzuk basin IX, 17, 21, 22 Mustahil Formation 217 Mut Formation 65 Nabis 241 Nabis Formation 238 Nanka Formation 156, 159 Naphthenic, nonwaxy medium crude 171 Naqus Formation 49, 53, 54 Natal Group 305 Natal Trough 269, 305 Neo-Tethys 227 Ngerengere Beds 218 Niger Delta XI, 151, 152, 159, 163 Niger Delta Basin 151-154, 174 Niger Delta crude 171 Niger-Benue drainage 152 Nigeria Escarpment 165 Nile 76 Nile Delta 43, 73, 75, 81 Nile Delta basin 45, 77 Nkporo cycle 159 Nkporo depositional cycle 156 Nkporo Formation 98 Nkporo Group 158 Nkporo Shale 156 Normandien Formation 295 North Sea 31 Northern Delta depobelt 168 Northern Plate Margin of Africa 379 Nosib Group 238, 252, 266 Nsongezi Series 143

Nsukka Formation 159 Nubian 13, 25 Nubian Sandstone 13, 25, 55, 67, 79, 105 Nubian Sandstone Formation 109 Nubian shield 40 Numidian Sandstone 95 Oceanic gulfs V Offshore depobelts 169 Ogooue River 173 Ogwashi-Asaba Formation 166 Oil 310 Olukonda 241 Olukonda Formation 261,266 Ombalantu 241 Ombalantu Formation 260, 261,263, 266 Orange River 173, 184 Orange River basin 173, 184 Ordovician glaciation 20 Otavi Group 238, 243, 252, 266 Oti Shale Formation 91 Otterburn Formation 299 Ouan Casa 25 Ouan Casa Formation 11 Oued Mya 20 Ougarta 19 Owambo 241 Owambo Basin 237 Owambo Formation 252, 254, 256 Owamboland 266 Owelli Formation 159 Owelli Sandstones 156 Owen fracture 225 Owen fracture zone 211,227 Palaeochannel fills 171 Palaeocurrents 55, 60 Palaeodrainage 169 Palaeo-Nile 75, 80, 81 Palaeo-Nile Delta 79 Pan-African 211 Pan-African Event 106, 109 Pangea 55, 66 Paraffinic waxy crude 171 Parasequence sets 159, 166 Passive margins V Pelusium Line 41 Petroleum reserves 170 Pietermaritzburg Formation 288, 309 Pinda Group 181 Pleistocene 4 Pletmos 319 Pluto's Vale Member 280 Point Indienne Formation 180 Polignac 13 Polignac basin 20 Poortjie Member 311 Port Elizabeth 360 Precambrian 20 Prince Albert 241 Prince Albert Formation 256, 257, 266, 275, 308 Purpre d'Ahnet 21 Qiba Member 302 Qoz Dango trough 122 Quianga Group 183 Quseib Formation 53

392 Quseir Formation 65 Rakhlyat Formation 65 Rasthof Member 247, 251 Ravinement surface 162 Red Sea 39, 60, 61, 65, 70, 71, 77, 82, 111, 113, 11,5 Reggane 17, 1 20 Reggane basiw 12 Reguba-Samah 33 Reguibat massif 19 Reservoirs for hydrocarbon accumulation 105 Rift V Rift basins 319 Rift basins of the Sudan 105 Rifting 319 Rima Group 97-99 Rio del Rey basin 169 Ripon Formation 280, 308 River Atbara 146 River Atbara rift 116, 124 River Atbara rift basin 133 River Congo 173 River Cuanza 173 River Nile 82 River Nile Sandstone 109 Rollover structures 170 Ruaga Limestone 29 Sabaya Formation 59, 60, 62 Sabil Formation 29 Sail Formation 50 Sag El Naam graben 118 Sag El Naam trough 121 Sahabi 29 Sahabi Channel 81 Sahara Flexure 380 Saharan Atlas 380 Saline lakes 146 Salinity crisis 80 Samah 28 Sanaga River 173 Sand pinch-outs 171 Santa Formation 93 Sarir (Faragh, Calanscio) Sandstone 68 Sarir Sandstone 28 Satal Formation 29 Second-order cycle 162 Sequence stratigraphy 159, 166 SErie Argilo-Sableuse Lignites 101 SEries Argileuse 99 SEries Siderolithique 101 Shagra trough 116 Sharaf Formation 109, 111 Sinai 39, 55, 61, 62, 72, 73, 79, 80 Singa subbasin 133 Sirt basin 68, 70-72, 80, 81 Sirte X Sirte basin 15, 17, 27 Sirte Shale Formation 29 Siwa 43, 81 Siwa basin 47, 49, 53 Six Hills Formation 57-59 Skoorsteenberg Formation 284, 308 Soekor 322

GEOGRAPHIC, TECTONIC AND STRATIGRAPHIC I N D E X Sokoto basin 95, 97, 98 Sokoto cycle 156 Sokoto Embayment 156 Sokoto Group 98 Somali Coastal Basin 214 Somalia 323 South Africa 173 South America 324 South Atlantic Ocean 173 Southern Cape Conductive Belt 305 Stacked reservoirs 171 Stratigraphic traps 171 Stromatolites 80 Submarine 165, 166 Sudanese 113 Sudanese rift structures 105 Sudanese rift system 111, 114 Sudanese rift 134 Sudd graben 114, 115, 130 Sudr Formation 66 Sundays River Formation 323 Sundays River Trough 358 Superior Fault 347, 329 Suurberg Group 323 Swartkops 360 Swartkops Member 360, 370 Syrian Arc 61, 64, 65, 68, 69, 82 System 113 Table Mountain Group IX, 329 Table Mountain quartzites 359 Tadrart 50 Tadrart Sandstone Formation 11, 25, 51 Taleh Formation 217 Taloka 97 Talrass Formation 95 Tamaguilelt Formation 100 Tannezuft Shale Formation 10, 22, 25 Tanzania 323 Tanzania Coastal Basin 221 Taref Formation 63 Tarkastad Subgroup 299 Taru Formation 218 Tassilian discordance IX Tchirezrine Formation 93 Teekloof Formation 295 Tegama Group 92, 93 Teloua Sandstone Formation 92 Terzikasan Formation 93, 94 Tethyan 19 Tethyan Ocean IX, 7, 11, 17, 22, 25 Tethys IX, 3, 22, 55, 56, 59-61, 66, 68, 69, 76, 80, 81, 96 Tethys ocean 43 Thebes Formation 71 Tibesti Mountains 35 Tibesti-Sirte 21, 25 Tibesti-Sirte arch 24, 27, 28, 31 Tierberg Formation 283, 308 Tiguedi Series 93 Tihemboka mole 21 Tillite 51, 54 Timimoun 13 Tin Sakan Formation 94 Tindouf 17, 19, 20 Tindouf basin 17

Toca Formation 180 Transgressive-regressive cycles 158 Trenches V Triassic Salt basin 20 Tripoli-Gabes basin 68, 70 Trumpeters Member 280 Truncation 171 Tschudi 241 Tschudi Formation 252, 253, 266 Tsomo Member 302 Tsumeb Subgroup 245 Tuenza Group 182 Turbidity currents 165 Type 1 sequence boundary 162 Uanai Member 215 Uitenhage Trough 358 Umm Bogma Formation 52 Umm Ruwaba basin 133-135 Umm Ruwaba graben 115, 123, 140 Upper Corvo 185 Upper Cuvo Formation 182 Upper Nile 57, 65 Upper Nile basin 65, 71 Urandab Formation 215 Uranium 310 Uweinat 39 Vandji Formation 180 Varianto 241 Varianto Formation 238, 242 Verkykerskop Formation 300 Vermelha Formation 181 Vertical cyclic stacking 166 Vischkuil Formation 278, 308 Volksrust Formation 295, 309 Vryheid Formation 289, 309 Wadi Abdul Malik 54 Wadi El Kej graben 116 Wadi El Kuu rift 116, 146 Wadi Malik Formation 51 Wagadi Series 92 Wajee Shale Formation 100 Walvis ridge 183, 184 Wata Formation 62 Waterford Formation 282, 286, 309 Waw en Namus 34 Western Desert 53, 57, 58, 60, 64-66, 68, 71, 72, 75, 76, 79, 80, 82 Western Desert basins 50 White Nile basin 133, 134 White Nile graben 116 White Nile rift 115, 129, 133, 146 Whitehill Formation 277, 308 Wonderfontein Member 280 Wurno 97 Wurno Formation 98 Yola subbasin 153 Zella (Tagrifet) 31 Zella (Tagrifet) trough 35 Zelten-Waha ridge 33

Minerals, Petroleum, Rocks and Fossils Index

Abu Madi 45 Abu Qir gas fields 45 Acritarchs 319 Agulhas Current 322 A1 Alamain field 58 Alemanide 380 Alluvial fans 146, 341 Alpine-Mediterranean tectonics 380 Amal Field 28 Amia 97 Anemiasporites sp. 95 Angola 173 Anhydrite 29, 76, 77, 96, 217 Anhydrites 221 Aniopyge cf. prerassulata 53 Anoxic 345 Anthracite 310 Anthrophycus 49 Archiepelago model of orogenesis 380 Asteriacite gugelhupf 51 Atlantic-type passive margin 322 Atlas Fold Belt 379 Augila field 28 Bahi 29 Banded iron 242 Bantiu oil field 115 Basalts 34 Berriasian 360 Bifungites 49

Bifungites fezzanensis 50, 51

Coaly 218, 222, 230 Coffinite 311

Collenia 12 Collophane 82 Collophanite 276 Compressional 371 Conglomerates 360 Coniacian 347 Conophyton 243, 249

Conostichus broadhedi 51 Cretaceous anoxic events 185 Cruziana 6, 8, 10, 48, 49, 282, 284, 286-288 Cruziana rouaulti 49 Cyclicity 357

Halite 198, 263, 277

Dadoxylon 13, 95 Dadoxylon (Araucarioxylon) 95 Dahra B 29 Deep sea drilling project 323 Deep-marine 319 Deformation 371 Deltas and fans 146 Deltoidonautilus molli 99 Devonian 329, 370 Diamictite 242, 270, 272, 274 Diatoms 198, 202, 205 Diatomites 198, 205 Dinoflagellates 319 Diplocraterion sp. 65 Divergent margin basins 322 Drag 371

Biostratigraphic 325 Bombacacidites sp. 101 Buried saline lakes, sabkhas or playas 105 Buttinia adnreevi 99

Echinopsis cf. Friryi 99 Echiperiporites icacinoides 1O1

Callianassa 97

Eremopteris whitei 51

Callovian 370

Camerotoechia spp. 50, 51 Campanian 347 Canyons 344, 373 Ceralites 55

Ceratiopsis granulostriatum 99 Chalk 15 Channels 344 Chara 97

Chenolophonidites costatus 101

Gas 310, 344 Gas fields 344 Gialo field 30 Glacial 242, 258, 266, 305, 308 Glauconite 17, 29, 97, 215, 277, 289 Glauconitic 30, 61, 63, 65, 71,294 Glenfica field 184 Globigerina triloculinoides 99 Glossopteris 285, 295 Glossopteris flora 269 Graben 341,348 Greensands 4, 15 Gypsum 80, 98, 100, 140, 198, 212, 217, 231,255, 261

Harlania 8, 10, 49 Hassi er R'Mel 13, 20 Hassi Messaoud 10, 20 Hassi Messaoud oil field 13 Hauterivian 355, 359 Hauterivian-Barremian 323 Heat flow 31 Heavy minerals 206 Highly saline ground-water bodies 146 Highstand systems tract 345 Hofra Field 28 Holocene 347 Horizon D 348 Hydrobia 14, 25 Igneous activity 329

Inoceramus 61,344

Economic deposits 206 E1 Wastani 45 Embergerixylon 95 Erosion 347 Eustatic 322 Evaporite 80, 255, 266 Evaporites 30, 76, 81, 140, 221, 231, 341 Evaporitic deposits 196, 198, 206

Exogyra 61 Exogyra columba 96 Exogyra olisiponensis 96

Intisar (formerly Idris) oil fields 30 Inversion 31, 34, 371 Iron ore 21 Jebel es Soda 34 Jordisite 312 K/Ar 323 Kaolin 232 Kaolinite 212 Kaolinitic 232 Kerogen 29, 185, 230 Kimberlite 309 Kimmeridgian 323 Kudu gas field 184, 185 Kungulo oil field 181 Kyanite 232

Cicatricosisporites dorogensis 101 Cimonia sudanensis 99 Cladophlebis 13 Classopollis sp. 95

F-A platform 322 Fan-channel 345 Fans 345 Fenestrillina aft. omaciata 49 Fluvial 341 Foraminifera 319, 357

Laffiteina 97 Laffiteina bibensis 99

Coal 258, 277, 291, 293, 294, 302, 310, 312

Galinda field 184

Lamprophyres 329 Laramide

Chronostratigraphic 325

Lacustrine 370

394 Late Oligocene 329 Latest Valanginian 323 Lava 34 Leiotriletes sp. 95 Lepidodendron 12, 25 Lepidodendron sp. 11 Lepisosteus 97 Leptostraphia magnifica 49 Leucoxene 232 Libycoceras afikpoensis 98 Libycoceras ismaeli 97 Lingula sp. 58, 61 Lithosphere-generated passive rifts 381 Lithospheric stretching 381 Lockhartia haemei 99, 100 Longapertiles microbaculatus 99 Lophoctenium 285 Lowstand 322 Lowstand system tracts 345 Maastrichtian 347 Malongo North field 182 Malongo oil field 181 Malongo West field 182 Mantled-generated active rifts 381 Mesembrioxylon 95 Mesosaurus 277 Mesosaurus stereosternum 277 Mesosaurus tenuidens 308 Messla 28 Miocene 347 Mississippi-Valley-type mineralisation 252 Moeritherium 101 Molybdenite 312 Molybdenum 310, 312 Monocraterion 7 Mosasaurus Shales Mulenvos field 184 Multistructural system of rifts 105 Namibia 173 Nasser (formerly Zelten) field 29 Natrocarbonatites 195 Neolobites 95 Neolobites vibrayeanus 95, 96 Nereites 278, 284 Nigericeras 96 Nigericeras lamberti 96 Notocaris tapscotti 277 Nummulites gezahensis 71 Oil 310, 345 Oil fields 182 Oligocene 347 Oncolites 263 Ooliths 373 Operculinoides bermudezi 99 Operculinoides sp. 100 Orbitolina 226

MINERALS, PETROLEUM, ROCKS AND FOSSILS I N D E X Organic enrichment 359 Ostracods 319 Ostrea 61 Oxfordian 323 Palaeo-placers 212 Palaeocene 347 Palaeocurrent 6, 22, 25, 94, 98, 281,284, 296, 297, 299, 300, 309 Palaeozoic 359 Paleonodante fischeri 218 Paravascocerascauvini 96 Periretisyncolpites giganteus 99 Periretisyncolpites magnosagenatus 99 Perotriletes sp. 95 Phosphate 17, 29, 61, 69, 97, 101,295 Phosphate pellets 373 Phosphatic 65, 71, 72, 99, 100, 276, 278 Phyllotheca 295 Plagiogmus 285 Planolites 284, 288 Platyracheela cf. mesastrials 49 Pleistocene 347 Plesiolampas cf. saharae 99 Plesiolampas saharae 100 Pletmos faults 329 Pollen 319 Polypterus 97 Polytaxis 51 Portlandian 323, 359 Posidonia 218, 224, 230 Productella cf. hallina 49 Protopodocarpoxyolon 95 Pycnodus bowerbankii 1O0 Radiolaria 357 Raguba 28 Red beds 355 Regressive clastic wedge 152 Retidisporites nigeriensis 99 Retro arc 381 Retro-arc basin 380 Reverse faults 371 Rhaetide 380 Rhizocorallium 58 Right lateral movement 319, 371 Rollover 371 Rutile 232

Sabellarifex 7 Sabkha 341 Sag 196 Sag basins IX Sag-type basin 191 Saline ground-water body 146 Salt 173, 182, 211,212 Salt diapirs 217 Sandstones 360 Santonian 347 Sarir Field 28

Schizaster 100 Schizaster (Linthia) sudanensis 99, 100 Scolithos 7, 8 Scoyenia 288 Skolithos 48, 49, 282, 284, 287, 288 Source rock 10, 25, 31, 45, 206, 345, 360, 371 Steevesipollenites sp. 95 Straimonocolpites microcanilis 99 Strata-bound 206 Striamonocolpites anastomosus 99 Striatapollis bellus 101 Strike-slip 371 Stromatolites 201, 243, 245, 248, 249, 251,263 Stromatolitic 203 Subsidence rates 357, 358 Takula oil field 181 Taphrogenic subsidence 153 Tassilian discordance 5, 6 Tell-Atlas/Rif 380 Teredo 57 Tertiary 329, 347 Tertiary deltas 151 Thalassinoides 65 Thenardite 263 Theodessia aft. hungerfordi 49 Tigillites 7, 8, 10, 11, 12 Tillite 245, 258, 289 Tillite fill 256 Tobias field 184 Topographical doming 381 Torbanite 294, 310 Trachytic 329 Transform 322 Tricolpopollenites sp. 95 Trochamminoides 357 Trona 198, 263 Tsumeb Mine 256 Tuff 329 Turbidite 345 Turbidity 344 Turonian 347, 358

Umfolozia 279 Undichna bina 279, 282 Unio 96 Unity field 111 Uraninite 311 Uranium 258, 310, 311 Valanginian 355 Varves 258 Velascoensis Event 71

Waagenocencha montepelierenses 53 Zircon 232 Zoophycos 284

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