The Geology of the Arabian-Nubian Shield (Regional Geology Reviews) 303072994X, 9783030729943

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Regional Geology Reviews

Zakaria Hamimi · Abdel-Rahman Fowler · Jean-Paul Liégeois · Alan Collins · Mohamed G. Abdelsalam · Mohamed Abd EI-Wahed Editors

The Geology of the ArabianNubian Shield

Regional Geology Reviews Series Editors Roland Oberhänsli, Institute of Earth and Environmental Sciences, University of Potsdam, Potsdam, Brandenburg, Germany Francois Roure, Geology-Geochemistry-Geophysics Department, French Petroleum Institute, Rueil-Malmaison, France Dirk Frei, Department of Earth Sciences, University of the Western Cape, Bellville, South Africa

The Geology of series seeks to systematically present the geology of each country, region and continent on Earth. Each book aims to provide the reader with the state-of-the-art understanding of a regions geology with subsequent updated editions appearing every 5 to 10 years and accompanied by an online “must read” reference list, which will be updated each year. The books should form the basis of understanding that students, researchers and professional geologists require when beginning investigations in a particular area and are encouraged to include as much information as possible such as: Maps and Cross-sections, Past and current models, Geophysical investigations, Geochemical Datasets, Economic Geology, Geotourism (Geoparks etc), Geo-environmental/ecological concerns, etc.

More information about this series at http://www.springer.com/series/8643

Zakaria Hamimi
Abdel-Rahman Fowler
Jean-Paul Liégeois
Alan Collins
Mohamed G. Abdelsalam
Mohamed Abd EI-Wahed Editors

The Geology of the Arabian-Nubian Shield

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Editors Zakaria Hamimi Department of Geology Faculty of Science Banha University Benha, Egypt Jean-Paul Liégeois Department of Geology and Mineralogy Royal Museum for Central Africa Tervuren, Belgium Mohamed G. Abdelsalam Boone Pickens School of Geology Oklahoma State University Oklahoma City, Oklahoma, USA

Abdel-Rahman Fowler Department of Geology United Arab Emirates University Al Ain, United Arab Emirates Alan Collins Department of Earth Sciences University of Adelaide Adelaide, SA, Australia Mohamed Abd EI-Wahed Department of Geology Tanta University Tanta, Egypt

ISSN 2364-6438 ISSN 2364-6446 (electronic) Regional Geology Reviews ISBN 978-3-030-72994-3 ISBN 978-3-030-72995-0 (eBook) https://doi.org/10.1007/978-3-030-72995-0 © The Editor(s) (if applicable) and The Author(s), under exclusive license to Springer Nature Switzerland AG 2021 This work is subject to copyright. All rights are solely and exclusively licensed by the Publisher, whether the whole or part of the material is concerned, specifically the rights of translation, reprinting, reuse of illustrations, recitation, broadcasting, reproduction on microfilms or in any other physical way, and transmission or information storage and retrieval, electronic adaptation, computer software, or by similar or dissimilar methodology now known or hereafter developed. The use of general descriptive names, registered names, trademarks, service marks, etc. in this publication does not imply, even in the absence of a specific statement, that such names are exempt from the relevant protective laws and regulations and therefore free for general use. The publisher, the authors and the editors are safe to assume that the advice and information in this book are believed to be true and accurate at the date of publication. Neither the publisher nor the authors or the editors give a warranty, expressed or implied, with respect to the material contained herein or for any errors or omissions that may have been made. The publisher remains neutral with regard to jurisdictional claims in published maps and institutional affiliations. Cover illustration: Panoramic view showing NW (to NNW) verging thrust-related folds in the middle carbonate unit (Shubayrim Formation) of the Fatima Group (post-amalgamation volcanosedimentary sequence). Southwest of Wadi Daf, Fatima Shear Belt, Western Arabia. Photo by Zakaria Hamimi This Springer imprint is published by the registered company Springer Nature Switzerland AG The registered company address is: Gewerbestrasse 11, 6330 Cham, Switzerland

This book is dedicated to all universities and geological surveys in the countries of the Arabian-Nubian Shield (ANS) who, regardless of their limited resources at times, continued to provide crucial logistical support that led to the successful completion of hundreds if not thousands of field excursions resulting in the collection of geoscientific data that helped bringing us to this level of understanding of the geodynamic evolution of the shield. It is dedicated to the spouses and children of the ANS Geoscientists, who condoned their absence numerous times in the field. It is dedicated to the army of unsung heroes of drivers, cooks, and field assistances, without their support field activities in the ANS would have been almost impossible. It is dedicated to the local inhabitance of the ANS, who always generously welcomed Geoscientists in their land to collect field data. It is dedicated to the pioneer Geoscientists, who laid the foundation for subsequent studies in the shield. Finally, it is dedicated to all who have contributed their efforts and insights into understanding the complex geological history of the ANS. Guest Editors

Preface

Why This Book? The Arabian–Nubian Shield (ANS) is a vast and exceptionally well-preserved expanse of crust and underlying sub-continental lithospheric mantle that was largely produced during the Neoproterozoic, with only restricted contribution from older lithospheric components, except in the extreme north-eastern part of the shield. With an exposed area of nearly 2 million km², and an additional 1 million km² buried beneath the great Cenozoic East African Rift System and its associated volcanic rocks, the shield represents one of the best-preserved Neoproterozoic orogenic belts on the planet and the largest preserved expanse of new Neoproterozoic lithosphere in Earth’s geological history. The ANS was formed between ca. 870 and 540 Ma and forms the northern part of the vast East African Orogen (EAO). Earlier Tonian juvenile crust is also found in Sudan, Madagascar, East Africa and Antarctica, which suggests that the Tonian–Cambrian ANS is really the continuum of subduction-accretion magmatism on the periphery of the Mesoproterozoic–Neoproterozoic African continents. The first period of complex tectonic history occurred between 870 and 640 Ma, which corresponds to the initiation of oceanic arcs and back-arcs followed by their accretion on top of each other within an oceanic tectonic environment or towards the peripheral margins of continents. This resulted in the formation of a series of juvenile Neoproterozoic tectonic terranes of various ages and an array of ophiolitic assemblages. The 640–540 Ma period subsequently witnessed the final continental amalgamation, with the collision with eastern continent including Neoproterozoic India. The ultimate development of post-collisional continental scale mega-shear zones linked to tectonic escape, and later the overprint of molassic basins and voluminous magmatism in the form of dike swarms, ultimately resulting in the emergence of the present appearance of the shield. While in the east, the former landmass that made up eastern Gondwana immediately east of the ANS in Arabia has been masked by younger Phanerozoic tectonic events including the opening of the Indian Ocean, to the west, continental expression of western Gondwana is still intact, represented by the Saharan metacraton. Although the last decade has witnessed significant progress in knowledge of this enigmatic geological feature, its relationship with the ANS and its evolution during the formation of the latter are still poorly known. Documentation of the juvenile character of the ANS can be credited to the excellent preservation of all stages of its development. This is largely because the final collision period did not involve extensive shortening and crustal thickening, largely preserving the preceding tectonic features. This left the ANS as one of the best-preserved examples of an accretionary orogen, from the initial oceanic stages leading to the main collisional orogenic stage, together with the opportunities for studies aimed at understanding the controls of pre-existing lithospheric structure of the Neoproterozoic terranes on the later geologic events that came after the orogenic phases. Understanding the core question of “lithospheric memory” is crucial for advancing our knowledge in both geodynamic evolution of the Earth from the deep geological time to present and for societal benefit through the understanding of mineralization processes. vii

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The excellent preservation of all stages of its orogenic development, however, presents remarkable research opportunities for geoscientists, and it also introduces a complex and challenging geology in the ANS that requires the implementation of full range of Earth science techniques, including the most recent and advanced ones. It is remarkable to see how the advent of precise geochronology using U–Pb isotopic methods to date zircon and other phases has resulted in re-thinking early interpretations and allowed for new ideas to emerge. These ideas that originated from the shield are now flourishing around the World. To provide a solid basis for future studies, it is therefore time to bring together in one book both chapters that provide up-to-date syntheses of the current tectonic models for the evolution of the ANS, as well as chapters that present new geoscientific data that will further advance our understanding of the shield. Especially, the editors hope that this book will inspire the talented new generation of Arab and African Geoscientists and Geoscientists interested in the Precambrian of Africa to engage in research that might build on some presentations of the book or challenge others. The goal is to arrive at more accurate and comprehensive tectonic models for the shield. Advancing knowledge with the ANS is indeed dynamic, requires the continuous embrace of new and innovative ideas, and needs to be kept as an incessant process.

Contents The chapters of this book have been grouped according to five themes that recognize shared contexts among the book topics. These five themes include the nature and boundaries of the ANS; the juvenile and pre-Neoproterozoic components of the ANS; the deformation events, magmatism/volcanism and sedimentary basins evolution of the ANS; key petrological, mineralogical and geochronological aspects of the ANS; and Phanerozoic events that followed the ANS. The following well-known scientists and influential scholars (arranged in alphabetical order) contributed the five themes: Tamer Abu-Alam, Ahmed H. Ahmed, Tadesse Alemu, Salah Al-Khirbash, Khaled Al Selwi, Richard Armstrong, Paul D. Asimow, Mokhles K. Azer, Mahrous Abu El-Enein, Hamdy H. Abd El-Naby, Kamal Ali, Alan S. Collins, Nagy Shawky Botros, Adam J. Bumby, Osama Dessouky, Gaafar A El Bahariya, Morgan L. Blades, Ahmed Mohammed Eldosouky, Reda A. Y. El-Qassas, Khalid A. Elsayed Zeinelabdein, Abdel Moneim A. El-Dougdoug, Abdalla E.M. Elsheikh, Abdel-Rahman Fowler, Harald Fritz, Hind Ghanem, Moustafa E. Gharib, Geoff H. Grantham, Zakaria Hamimi, Hesham M. Harbi, Mahmoud Hassan, Christoph Hauzenberger, Mohamed Th. S. Heikal, Montasir A. Ibinoof, Ghaleb H. Jarrar, Peter R. Johnson, Mohamed S. Kamar, Petrus Le Roux, Jean-Paul Liégeois, Ahmed A. Madani, Ayman E. Maurice, Andrew S. Merdith, Nathan R. Miller, Hamed I Mira, Eiman A. Mohamed, Erdinc Oksum, Ali Farrag Osman, Luan Thanh Pham, Gehad M. Saleh, Hussam A. Selim, Mohamed G. Shahien, Ahmed Abu Sharib, Robert J. Stern, Adel A. Surour, Jacques Varet, Martin Whitehouse, Simon A. Wilde, Brian F. Windley and Basem Zoheir. We would like to thank all of them for their new insights into the geology of the ANS. The first theme of this book explores the ANS as entirety, beginning with Chap. 1: “The Arabian-Nubian Shield, An Introduction: Historic Overview, Concepts, Interpretations, and Future Issues” by P. R. Johnson, which deals with the comprehensive evolution and tectonic setting of the ANS in association with neighboring Proterozoic and older cratons. In Chap. 2: “The Boundary Between the Saharan Metacraton and the Arabian-Nubian Shield: Insight from Ediacaran Granites in the Nuba Mountains (Sudan): Geochemistry, Sr-Nd Isotopes and U–Pb SHRIMP Zircon Dating” by M. A. Ibinoof, A. J. Bumby, J.-P. Liégeois, G. H. Grantham, R. Armstrong and P. Le Roux, granite geochemistry is used to identify the tectonic boundary between the ANS and south-eastern Saharan metacraton. The southernmost extent of the ANS is recognized by H. Fritz and C. Hauzenberger, in Chap. 3: “The Southern Tip

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of the Arabian-Nubian Shield: The Arc Continent Transition” as lying in southern Madagascar to northern Mozambique, which is the southern limit of Neoproterozoic oceanic basins and arcs that were closed during collision of Neoproterozoic India with the African continents. In the following Chap. 4: “Tectonic Evolution of the Pan-African Belt in Western Ethiopia, Southern Arabian-Nubian Shield” by A. Tadesse-Alemu, three N–S terranes separated by shear zones are recognized in the Neoproterozoic exposures of western Ethiopia. The magmatic and structural histories of these terranes are broadly correlated with equivalent ANS events. At the northern extremes of the ANS, in Jordan, G. H Jarrar and H. Ghanem recount in Chap. 5: “Neoproterozoic Crustal Evolution of the Northernmost Arabian-Nubian Shield, South Jordan”, the evolution of the ANS from Tonian tonalitic arc magmatism at 787 Ma to Ediacaran alkaline/peralkaline volcanism and dyking at 580 Ma. In Chap. 6: “Geophysics of the Lithosphere of the Arabian-Nubian Shield” by A. M. Eldosouky, L. T. Pham, E. Oksum and R. A. Y. El-Qassas, potential field, magnetic and gravity data are analysed across the ANS to prepare detailed structural maps, Curie depth maps and lithosphere–asthenosphere boundary maps for the ANS. Global-scale glaciation events in the latest Tonian-early Cryogenian (Sturtian glaciation), late Cryogenian (Marinoan glaciation) and Ediacaran (Gaskiers glaciation) are discussed in Chap. 7: “Arabian-Nubian Shield Evolution and Snowball Earth” by N. R. Miller and R. J. Stern. The tectonic, atmospheric and bioevolutionary factors bearing on these panglacial events are explored. Finally, in this section A. S. Collins, M. L. Blades and A. S. Merdith place the ANS into a global plate tectonic context and use the geology of the shield to reflect large-scale plate tectonic changes occurring at the time in Chap. 8: “The Arabian-Nubian Shield Within the Neoproterozoic Plate Tectonic Circuit”. The second theme of this book presents studies of the important early events of the ANS, from Rodinia break-up to the island arc and subduction stages. In Chap. 9: “Early Ensimatic Stage of the Arabian-Nubian Shield” by M. Hassan, A. Fowler, O. Dessouky and T. Abu-Alam, the links between the earliest tectonic stages of the ANS and models of Rodinia break-up are considered. Evidence bearing on this relationship include structural, petrological and geochemical, and especially geochronological data. The role of terrane tectonics in the growth of the ANS accretionary orogen is related, in Chap. 10: “Terrane Accretion Within the Arabian-Nubian Shield” by A. F. Osman and A. Fowler. Two important dynamically similar processes discussed are “amalgamation” (terrane assembly in an oceanic setting) and “accretion” (terrane assembly at a continental margin). M. E. Gharib, A. E. Maurice and H. A. Selim, in Chap. 11: “Tonian–Cryogenian Volcanic Arcs of the Arabian-Nubian Shield”, provide detailed geochemical characterization of the Tonian–Cryogenian island arc metavolcanics that form a major component of the ANS, and compare the ANS arcs to modern arcs. They discuss the petrogenesis of the tholeiitic to calc-alkaline volcanics and correlate them with primitive and mature arcs, respectively. In Chap. 12: “Neoproterozoic Ophiolites of the Arabian-Nubian Shield” by M. G. Shahien, M. K. Azer and P. D. Azimow, the main features of the ANS ophiolites are described, particularly their age, metamorphic grades, alteration styles and associated mineralization. The geochemistry is argued to indicate a mainly forearc setting for the ANS ophiolites. A complete ophiolite sequence at Wadi Ghadir is the topic of Chap. 13: “Ghadir Ophiolite, Eastern Desert, Egypt: A Complete Sequence of Oceanic Lithosphere in the Arabian-Nubian Shield” by G. A. El Bahariya. In this chapter, the identity of mid-ocean ridge basalts (MORB) ophiolites as back-arc eruptives is suggested for the Egyptian Eastern Desert. Chapter 14: “Evidence for Mesoproterozoic Components in the Arabian-Nubian Shield” by H. H. Abd El-Naby discusses the important problem of the presence of pre-Neoproterozoic (Palaeoproterozoic and Archean) zircons in the sediments and juvenile igneous rocks of the ANS. The candidates for sources of these older zircons include unexposed older crustal sources underlying the shield and pre-Neoproterozoic neighboring cratons. The third theme of this book examines events subsequent to the arc volcanic—subduction stage and deals with deformation events associated with terrane collisions, and post-amalgamation volcanism and sedimentary basin formation. The impressive strike-slip Najd Fault System (NFS) is treated in Chap. 15: “Najd Shear System in the Arabian-Nubian

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Shield” by Z. Hamimi and A. Fowler. This chapter deals with a wide range of field, remote sensing and geophysical characteristics and models for the NFS, in addition to examining hydrogeological and mineralization aspects of the NFS. In Chap. 16: “Low-Angle Normal Shear Zones, Folds and Wrench Faults During the Post-Amalgamation Stage of the Arabian-Nubian Shield” by A. S. A. A. Abu Sharib, the post-accretion deformations are divided into “active” and “passive” responses to regional stress fields. Wrench faulting is active, while low-angle normal faulting is passive and accompanies gravitational extensional collapse. Reactivation of earlier structures and overlapping of structural events make a complex pattern of deformation in the late stages of the ANS. A. A. Surour compares several models for magmatic evolution of the ANS, in the light of age data, geochemistry and petrogenetic models and field characteristics and relations, in Chap. 17: “Arc accretion and calc-alkaline plutonism at the end of the subduction stage in the Arabian-Nubian Shield”, with emphasis on the gneissic core complexes. The important role of arc magmatism is emphasized in the formation of core/gneiss complexes of the ANS. In Chap. 18: Origin of the volcanic-arc signature in the late-orogenic granitoids of the Arabian-Nubian Shield, by B. Zoheir and A. Diab, the strikingly similar geochemistry between late-orogenic granitoids and magmatic arc rocks of the subduction stage is explained by high temperature melting of subduction-related lower crustal rocks, combined with mantle derived melts. Chapter 19: “Post-amalgamation Depositional Basins in the Arabian-Nubian Shield: The Hammamat Basins of Egypt” by A. Fowler and Z. Hamimi focusses on the Egyptian Eastern Desert molasse basins and makes a case for dividing these into “larger” and “smaller” Hammamat basins. There are apparent distinctions between the two groups, based on age, stratigraphy, petrology, geochemistry and probably basin subsidence mechanisms. The topic of Chap. 20: “Volcanism During the Post-accretion Stage of the Arabian-Nubian Shield” by M. K. Azer, P. D. Asimow and S. A. Wilde is post-accretion magmatism and associated volcanosedimentary sequences. In this chapter, the authors proposed that an early calc-alkaline phase in its final stages records evidence for mantle delamination, while the following alkaline/peralkaline phase reflects a change to extensional tectonic environment. The fourth theme of this book is dedicated to studies involving important minerals and occurrences of mineralization in the ANS. Chapter 21: “An Overview Study of Zircon Geochronology from Sinai Crystalline Basement: Implication for Crustal Evolution of Northern Arabian-Nubian Shield”, by M. M. Abu El-Enen and K. Ali, describes the magmatic zircon populations that extend back to 1030 Ma and correlate them with similar aged zircon populations in the clastic deposits, which are believed to have occupied arc-related basins. The authors also report on the metamorphic zircons and the two high-grade metamorphic events that formed them. The following Chap. 22: “Spectral Characteristics of Listvenites and Serpentinites Along Ophiolite-Decorated Megashears (Suture Zones) in the Arabian Shield Using ASD Fieldspec and Satellite Data” by A. A. Madani, H. M. Harbi, A. A. El-Dougdoug, A. A. Surour and Ahmed H. Ahmed, uses remote sensing data to map and distinguish serpentinites and listvenites in some areas of the ANS ultramafic rocks. Mineral phases controlling spectral reflectances are hydroxyl-phases and a range of characteristic carbonates in the listvenites. The broad topic of economic mineral occurrences in the ANS is comprehensively covered in Chap. 23: “Ore Deposits in the Arabian-Nubian Shield” by N. S. Botros. Among the deposits described are orogenic gold—volcanogenic massive sulphides (VMS), polymetallic and volcanic-hosted gold deposits; pegmatitic rate earth elements (REE) and tantalum ores, TiO2 deposits, band iron formations, Cu–Ni–Co sulphide and chromium ores, and Sn, W, Mo, Be and U enrichments. The mineralization occurrences in the Precambrian of Yemen are described in Chap. 24: “Evolution and Mineralization of the Precambrian Basement of Yemen” by S. Al-Khirbash, M. T. S. Heikal, M. Whitehouse, B. F. Windley and K. Al Selwi. The authors classify the deposits as magmatic (ultramafic complex related), magmatic hydrothermal (mafic metavolcanic, metasedimentary and shear zone hosted) and late hydrothermal (associated with pegmatites and late-stage plutons). Applications of remote sensing and GIS for mapping mineral and groundwater prospecting in NE Sudan are outlined

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in Chap. 25: “Application of Remote Sensing and GIS in Geological Mapping, Mineral Prospecting and Groundwater Investigations in the Arabian-Nubian Shield: Cases from the Red Sea Hills of NE Sudan” by K. A. E. Zeinelabdein, E. A. Mohamed and A. E. M. Elsheikh. Examples of successes using these methods include discovery of auriferous gossans and porphyry Cu deposits of the region, and lithological/structural maps useful for groundwater investigations are given. The final theme of this book reports studies of Phanerozoic events affecting the ANS. In Chap. 26: “Phanerozoic Minor Volcanics and Intrusives of the Arabian-Nubian Shield” by Gehad M. Saleh, Mohamed S. Kamar and Hamed I. Mira, locally uraniferous alkali ring complexes, Natash volcanics and felsite, trachyte and bostonite dykes, sheets and plugs are described. Also studied are Oligocene to Quaternary Red Sea rift-related mantle plume and continental flood volcanics in Saudi Arabia and Jordan. In Chap. 27: “Relationship of the Pan-African Tectonic Structures with the Opening of the Afar Triple Junction” by J. Varet, the characteristics and history of the Afar triple junction linking the Red Sea, Gulf of Aden and Main Ethiopian rifts are described, and the likely inheritance of the Afar triple junction geometry from earlier Neoproterozoic structures is debated. Benha, Egypt Al Ain, UAE Tervuren, Belgium Adelaide, Australia Oklahoma City, USA Tanta, Egypt

Zakaria Hamimi Abdel-Rahman Fowler Jean-Paul Liégeois Alan Collins Mohamed G. Abdelsalam Mohamed Abd El-Wahed

Acknowledgements

The guest editors of this volume would like to thank the following reviewers (arranged in Alphabetical order) for their thoughtful reviews that improved the manuscripts: Ali M. A. Abdalla, Abdel-Aal Abdel-Karim, Mohamed Abdelkarim, Yasser Abdelrahman, Mohamed Abdelsalam, Tamer Abu Alam, Ahmed Abusharib, Ahmed H. Ahmed, Kamal Ali, Salah Al-Khirbash, Shoji Arai, Mulugeta Araya, Mohamed Arnous, Asfawossen Asrat, Asran M. Asran, Mokhles Azer, Abdelrahman Bendaoud, Enrico Bonatti, Nelson Boniface, Bernard Bonin, Robert Bussert, Daniel Clark-Lowes, Alan Collins, Mohammed El-Bialy, Baher El-Kalioubi, Samir Kamh, Mohamed El-Sharkawi, Ashraf Emam, Abdel-Rahman Fowler, Anke Friedrich, Zakaria Hamimi, Haral Fritz, Jean-François Ghienne, Lee A. Groat, Wael Hagag, Hisham Jahlan, Ghaleb Jarrar, Peter Johnson, Ibrahim Khalaf, Jean-Paul Liégeois, Stephen Mccourt, Abdelkader Moghazi, Franz Neubauer, Victoria Pease, Amin Beiranvand Pour, Juergen Reinhardt, Minghua Ren, Robert Stern, Kurt Stüwe, Adel Surour, Martin Whitehouse, Simon Wilde, Gezahegn Yirgu and Khairy Zaki.

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Contents

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The Arabian–Nubian Shield, an Introduction: Historic Overview, Concepts, Interpretations, and Future Issues . . . . . . . . . . . . . . . . . . Peter R. Johnson 1.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.2 ANS: Plate Tectonic Setting . . . . . . . . . . . . . . . . . . . . . . . . . . 1.3 ANS: Geologic Investigations . . . . . . . . . . . . . . . . . . . . . . . . . 1.3.1 Historic Developments . . . . . . . . . . . . . . . . . . . . . . . 1.3.2 Mineral Resources and Exploration . . . . . . . . . . . . . . 1.3.3 Geologic Mapping, Surveys, and Findings . . . . . . . . . 1.4 ANS: Nomenclature . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.5 ANS: External Boundaries . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.5.1 ANS Northwestern Margin . . . . . . . . . . . . . . . . . . . . 1.5.2 ANS Southern Margin . . . . . . . . . . . . . . . . . . . . . . . 1.5.3 ANS Eastern Margin . . . . . . . . . . . . . . . . . . . . . . . . 1.5.4 ANS Margins in the Southwestern Arabian Peninsula 1.5.5 ANS Margins in the Eastern Arabian Peninsula . . . . . 1.5.6 ANS Northern Extent . . . . . . . . . . . . . . . . . . . . . . . . 1.6 ANS: Internal Divisions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.6.1 Volcanic Arcs; Where, How Many, and Origin? . . . . 1.6.2 Mineral Belts: An Aid to Metallogenic Analysis and Exploration Targeting . . . . . . . . . . . . . . . . . . . . . . . . 1.6.3 Structural or Tectonic Belts and Domains . . . . . . . . . 1.6.4 Divisions in the Egyptian Eastern Desert . . . . . . . . . . 1.6.5 Tectonostratigraphic Terranes . . . . . . . . . . . . . . . . . . 1.6.6 Isotopic Domains . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.6.7 Inherited Zircons . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.7 Concluding Comments . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

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The Boundary Between the Saharan Metacraton and the Arabian Nubian Shield: Insight from Ediacaran Shoshonitic Granites of the Nuba Mountains (Sudan): U–Pb SHRIMP Zircon Dating, Geochemistry and Sr–Nd Isotope Constraints . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Montasir A. Ibinoof, Adam J. Bumby, Jean-Paul Liégeois, Geoff H. Grantham, Richard Armstrong, and Petrus Le Roux 2.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.2 Granitoids in the ANS and the SmC and Change in Stress Regime . . . . 2.3 The SmC–ANS Boundary in Sudan . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.4 Geological Setting and Petrography . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.5 U–Pb SHRIMP Dating of Zircon . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.5.1 Analytical Techniques . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.5.2 Results . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

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Geochemistry of the Late-Orogenic Granitoids . . . . . . . . . . . . . . . . . 2.6.1 Analytical Techniques . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.6.2 Major and Trace Elements Chemistry . . . . . . . . . . . . . . . . 2.7 Sr and Nd Isotopes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.7.1 Analytical Techniques . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.7.2 Rb–Sr and Sm–Nd Isotopic Results . . . . . . . . . . . . . . . . . . 2.8 Discussion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.8.1 Lithochemistry and Structural Characterization of the Nuba Post-collisional Plutons . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.8.2 The Source of the Parental Magma of the Nuba Post-collisional Plutons . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.8.3 Timing of Magmatism, Metamorphism and Regional Correlations . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.8.4 Implications for the Boundary of the Saharan Metacraton in Southern Sudan . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.9 Conclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

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The Southern Part of the Arabian–Nubian Shield in Kenya and Tanzania . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Harald Fritz and Christoph Hauzenberger 3.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.2 Geological Overview . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.3 Lithology, Formation Ages and Tectonic Setting . . . . . . . 3.3.1 Sobo Unit . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.3.2 Galana Unit . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.3.3 Sagala Unit . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.3.4 Kinyjki Unit/Kasigau Group . . . . . . . . . . . . . . . 3.3.5 Kurase Group/Eastern Granulites Metasediments 3.3.6 Eastern Granulites Metamagmatics . . . . . . . . . . 3.4 Isotopic Constraints . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.5 Tectonics and Metamorphism . . . . . . . . . . . . . . . . . . . . . 3.6 Interpretation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.7 Discussion and Conclusions . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

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Tectonic Evolution of the Pan-African Belt in Western Ethiopia, Southern Arabian-Nubian Shield . . . . . . . . . . . . . . . . . . . . . . . . . Tadesse Alemu 4.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.2 Regional Geologic Setting . . . . . . . . . . . . . . . . . . . . . . . . . . 4.3 Lithology . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.3.1 Gimbi Terrane . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.3.2 Nejo Terrane . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.3.3 Asosa Terrane . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.3.4 Tulu Dimtu–Baruda Belt . . . . . . . . . . . . . . . . . . . 4.3.5 Intrusive Rocks . . . . . . . . . . . . . . . . . . . . . . . . . . 4.3.6 Young Volcano-Sediments . . . . . . . . . . . . . . . . . . 4.4 Structures . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.4.1 Gimbi Terrane . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.4.2 Nejo Terrane . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.4.3 Asosa Terrane . . . . . . . . . . . . . . . . . . . . . . . . . . .

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4.4.4 Tulu Dimtu–Baruda Belt . 4.4.5 NW-Trending Structures . . 4.5 Discussion and Summary . . . . . . . . 4.5.1 Structural Evolution . . . . . 4.5.2 Tectonic Interpretation and of Terranes . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . 5

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Regional Correlation . . . . . . . . . . . . . . . . . . . . . . . . . . . 103 . . . . . . . . . . . . . . . . . . . . . . . . . . . 106

Neoproterozoic Crustal Evolution of the Northernmost Arabian-Nubian Shield, South Jordan . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Ghaleb H. Jarrar and Hind Ghanem 5.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.2 Aqaba Complex . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.2.1 Metamorphic Suites . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.2.2 Gabbroids and Granitoids . . . . . . . . . . . . . . . . . . . . . . . . . 5.3 Araba Complex . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.3.1 Safi Group . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.3.2 The Araba Mafic Suite . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.3.3 Humrat-Feinan-Mubarak Suite (HFMS) . . . . . . . . . . . . . . . 5.3.4 Aheimir Volcanic Suite (AVS) . . . . . . . . . . . . . . . . . . . . . 5.4 Ediacaran Dike Swarms . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.5 Summary and Conclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

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Lithospheric Structure of the Arabian–Nubian Shield Using Satellite Potential Field Data . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Ahmed M. Eldosouky, Luan Thanh Pham, Reda A. Y. El-Qassas, Zakaria Hamimi, and Erdinc Oksum 6.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.2 Data . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.3 Methodology . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.3.1 EHGA Method . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.3.2 Parker–Oldenburg Method . . . . . . . . . . . . . . . . . . . . . 6.4 Results . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.4.1 Edge Results . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.4.2 Moho and LAB Results . . . . . . . . . . . . . . . . . . . . . . . 6.5 Discussion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.6 Conclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

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Evolution of the Arabian Nubian Shield and Snowball Earth . . . . . . . . Nathan R. Miller and Robert J. Stern 7.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.2 Snowball Earth and Late Neoproterozoic Glaciation . . . . . . . . . . . . 7.3 Overview of ANS and EAO Development . . . . . . . . . . . . . . . . . . . 7.4 Expected Manifestations of Neoproterozoic Glaciations in the ANS and EAO . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.5 Evidence for Glaciation in the ANS . . . . . . . . . . . . . . . . . . . . . . . . 7.5.1 Evidence for Tonian (c. 780–755 Ma) Glaciation . . . . . . . 7.5.2 Evidence for Sturtian Glaciation (  717–659 Ma) . . . . . . 7.5.3 Evidence for Marinoan (Onset  650–639, to  635 Ma) and Ediacaran (  580–550 Ma) Glaciation . . . . . . . . . . . 7.6 Conclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

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The Arabian–Nubian Shield Within the Neoproterozoic Plate Tectonic Circuit . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Alan S. Collins, Morgan L. Blades, and Andrew S. Merdith 8.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8.2 A Full-Plate Tectonic Reconstruction . . . . . . . . . . . . . . . . . . . . . . . 8.3 The ANS Within the Northern ‘East African Orogen’ . . . . . . . . . . . 8.4 The Mozambique Ocean, Azania and Afif–Abas . . . . . . . . . . . . . . 8.5 The Eastern Margin of the EAO (NW India to Oman) . . . . . . . . . . 8.6 The Arabian–Nubian Shield (ANS) . . . . . . . . . . . . . . . . . . . . . . . . 8.7 The Western Margin of the EAO (the Eastern Saharan Metacraton) . 8.8 Discussion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8.9 Conclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Early Ensimatic Stage of the Arabian-Nubian Shield . . . . . . . . . . . . Mahmoud Hassan, Abdel-Rahman Fowler, Osama Dessouky, and Tamer Abu-Alam 9.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 9.2 Rodinia Pre-rifting Configurations . . . . . . . . . . . . . . . . . . . . . . 9.2.1 SWEAT Model . . . . . . . . . . . . . . . . . . . . . . . . . . . . 9.2.2 “Missing-Link” Model . . . . . . . . . . . . . . . . . . . . . . . 9.2.3 Other Models . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 9.3 Rodinia Rifting . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 9.4 Origins and Configurations of the Mozambique Ocean—Rift or Remnant? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 9.5 Remnants of Rodinia Within the ANS . . . . . . . . . . . . . . . . . . . 9.5.1 Rodinia Continental Signatures Within the ANS . . . . 9.5.2 Mozambique Ocean Signatures Within the ANS . . . . 9.6 Conclusions and Open Questions . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

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10 Terrane Accretion Within the Arabian-Nubian Shield . . . . . . . . . . . . . Ali Farrag Osman and Abdel-Rahman Fowler 10.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 10.2 The Rationale . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 10.2.1 Terrane Terminology . . . . . . . . . . . . . . . . . . . . . . . . . . 10.2.2 Formation of Terranes . . . . . . . . . . . . . . . . . . . . . . . . . 10.3 Formation and Amalgamation of Gondwana Supercontinent . . . . . 10.3.1 Neoproterozoic Pan-African Orogeny . . . . . . . . . . . . . . 10.3.2 The East African Orogen (EAO) . . . . . . . . . . . . . . . . . . 10.4 The Arabian-Nubian Shield (ANS) . . . . . . . . . . . . . . . . . . . . . . . 10.4.1 Terrane Analysis, Amalgamation, and Accretion of the ANS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 10.4.2 Assembly of the ANS and Crustal Growth . . . . . . . . . . 10.5 Tectonic Models of the ANS . . . . . . . . . . . . . . . . . . . . . . . . . . . . 10.5.1 Earlier Models of Evolution of ANS . . . . . . . . . . . . . . . 10.5.2 Recent Models of ANS Evolution: Arc-Accretion Model 10.6 Faults and Shear Zones in the ANS . . . . . . . . . . . . . . . . . . . . . . . 10.7 Ophiolite Zones as Evidence for Sutures in the ANS . . . . . . . . . . 10.7.1 Arc–Arc Sutures . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 10.7.2 Arc-Continent Sutures . . . . . . . . . . . . . . . . . . . . . . . . . 10.7.3 The Eastern Margin of the ANS . . . . . . . . . . . . . . . . . . 10.7.4 The Western Margin of the ANS . . . . . . . . . . . . . . . . .

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10.8 Post-accretionary Structures . . . . . . . . . . . . . . . . . . . 10.9 Post-amalgamation Basins of the NE Arabian Shield 10.10 Conclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

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11 Tonian/Cryogenian Island Arc Metavolcanics of the Arabian-Nubian Shield . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Moustafa E. Gharib, Ayman E. Maurice, and Hussam A. Selim 11.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 11.2 Island Arc Volcanism of the Northern Nubian Shield . . . . . . . . . . . . 11.2.1 Island Arc Volcanic Rocks of the Eastern Desert of Egypt . 11.2.2 Island Arc Volcanic Rocks of Sinai . . . . . . . . . . . . . . . . . . 11.2.3 Geochemistry and Petrogenesis of Island Arc Volcanics in the Eastern Desert of Egypt . . . . . . . . . . . . . . . . . . . . . 11.3 Island Arc Volcanism in the Arabian Shield . . . . . . . . . . . . . . . . . . . 11.4 Island Arc Volcanism in the Southern Nubian Shield . . . . . . . . . . . . 11.4.1 Island Arc Metavolcanics of Sudan . . . . . . . . . . . . . . . . . . 11.4.2 Island Arc Metavolcanics of Ethiopia . . . . . . . . . . . . . . . . 11.4.3 Island Arc Metavolcanics of Eritrea . . . . . . . . . . . . . . . . . 11.5 Conclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 12 Neoproterozoic Ophiolites of the Arabian-Nubian Shield . . . Mohamed G. Shahien, Mokhles K. Azer, and Paul D. Asimow 12.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 12.2 Tectonic History of the Arabian-Nubian Shield . . . . . . 12.3 Components of the ANS Ophiolites . . . . . . . . . . . . . . . 12.4 Geochronology . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 12.5 Protolith and Geodynamic Setting . . . . . . . . . . . . . . . . 12.6 Alteration and Metamorphism . . . . . . . . . . . . . . . . . . . 12.7 Examples of ANS Ophiolites . . . . . . . . . . . . . . . . . . . 12.7.1 Saudi Arabian Ophiolites: Jabal Al-Wask . . . 12.7.2 Egyptian Ophiolites: Wadi Ghadir . . . . . . . . . 12.8 Mineralization of the ANS Ophiolites . . . . . . . . . . . . . 12.8.1 Gold . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 12.8.2 Chromitite . . . . . . . . . . . . . . . . . . . . . . . . . . 12.9 Discussion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

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13 Ghadir Ophiolites, Eastern Desert, Egypt: A Complete Sequence of Oceanic Crust in the Arabian-Nubian Shield . . . . . . . . . . . . . . Gaafar A. El Bahariya 13.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 13.2 Geological Setting . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 13.2.1 Ghadir Ophiolite Sequence . . . . . . . . . . . . . . . . . . 13.2.2 Ghadir Ophiolitic Mélange . . . . . . . . . . . . . . . . . . 13.2.3 Dismembered Ophiolites . . . . . . . . . . . . . . . . . . . . 13.3 Geochemistry and Tectonic Setting . . . . . . . . . . . . . . . . . . . 13.3.1 Geochemical Characteristics . . . . . . . . . . . . . . . . . 13.3.2 Ophiolitic Affinity and Tectonic Setting . . . . . . . . . 13.3.3 Comparison with Other Ophiolites . . . . . . . . . . . . 13.4 Petrogenesis and Tectonic Evolution . . . . . . . . . . . . . . . . . . 13.5 Concluding Remarks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

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331 332 332 333 334 334 334 338 338 340 341 341

xx

14 Evidence for Mesoproterozoic Components in the Arabian-Nubian Shield . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Hamdy H. Abd El-Naby 14.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 14.2 Early Ideas on the Existence of Pre-Pan-African Crust in the ANS 14.3 Mesoproterozoic or Older Xenocrystic and Detrital Zircons in the ANS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 14.4 Nature of the Older Cratonic Areas Flanking the ANS . . . . . . . . . 14.5 Significance of the Mesoproterozoic Rocks in Crustal Evolution of the ANS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 14.6 Conclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

Contents

. . . . . 343 . . . . . 343 . . . . . 344 . . . . . 348 . . . . . 350 . . . . . 352 . . . . . 353 . . . . . 353

15 Najd Shear System in the Arabian-Nubian Shield . . . . . . . . . . . . . . . . . Zakaria Hamimi and Abdel-Rahman Fowler 15.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 15.2 The Arabian-Nubian Shield Within the East African Orogen . . . . . . 15.2.1 The Najd Orogeny in the Tectonic Frame of the Arabian Shield . . . . . . . . . . . . . . . . . . . . . . . . . . . 15.3 The Najd Fault System . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 15.3.1 The Main Najd Megashears . . . . . . . . . . . . . . . . . . . . . . 15.4 Evidence for the Timing of Deformation on the Najd Fault System . 15.5 Kinematic History of the NFS . . . . . . . . . . . . . . . . . . . . . . . . . . . . 15.6 Metamorphic Conditions of Najd Shearing . . . . . . . . . . . . . . . . . . . 15.7 Strain Analysis Across the Najd Shear Zones . . . . . . . . . . . . . . . . . 15.8 Remote Sensing and Geophysical Studies of the Najd Fault System 15.9 Brittle Evolution of the Najd Fault Zone . . . . . . . . . . . . . . . . . . . . 15.10 Role of Transpression in the Evolution of the Najd Fault System . . 15.11 N-S to NE-SW and Other Trending Dextral Strike-Slip Faults of the Southern Saudi Shield . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 15.12 Discussion of Najd as a Regional System . . . . . . . . . . . . . . . . . . . 15.12.1 Along-Strike Trends in NFS Characteristics . . . . . . . . . . . 15.12.2 Lithospheric Structure of the NFS . . . . . . . . . . . . . . . . . . 15.12.3 Najd Shear Corridor in Egyptian Eastern Desert . . . . . . . . 15.12.4 Tectonic Models for the NFS . . . . . . . . . . . . . . . . . . . . . 15.12.5 Brittle Reactivation of the NFS and Seismicity . . . . . . . . 15.12.6 Najd-Related Mineralization . . . . . . . . . . . . . . . . . . . . . . 15.12.7 Najd Role in Hydrogeological Systems . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

. . . . 359 . . . . 359 . . . . 360 . . . . . . . . . .

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361 363 365 367 368 369 370 371 374 374

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378 380 380 381 382 382 385 385 386 387

16 Low Angle Normal-Sense Shear Zones, Folds and Wrench Faults During the Post-Amalgamation Stage of the Arabian-Nubian Shield . . . . . . . . . . . Ahmed S. A. A. Abu Sharib 16.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 16.2 Geologic Setting . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 16.3 Post-Amalgamation Events . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 16.3.1 Taphrogenic Event . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 16.3.2 Wrenching . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 16.3.3 Post-Amalgamation Shortening . . . . . . . . . . . . . . . . . . . . . . 16.4 Discussion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 16.4.1 Extension: Motives, Evidence, and Deformation vs. Collapse Extension . . . . . . . . . . . . . . . . . . . . . . . . . . . . 16.4.2 Wrenching . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

. . 393 . . . . . . .

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393 394 394 394 397 399 402

. . 402 . . 407

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16.4.3 16.4.4

Shortening: Convergence- and Transpression-Related . . . . . . The Eastern Desert Low Angle to Sub-Horizontal Shear Zone: A Shortening-(Tectonic) versus Extensional Collapse-(Non-Tectonic) Related Fabric! . . . . . . . . . . . . . . . 16.4.5 Inversion Tectonics and Basins Inversion . . . . . . . . . . . . . . 16.4.6 Complications of Tectonic Interpretations: A Regional Single Versus Multi-models . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 16.5 Conclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

. . 408

. . 409 . . 412 . . 412 . . 412 . . 413

17 Arc Accretion and Calc-Alkaline Plutonism Finalizing the Subduction Stage in the Arabian–Nubian Shield, with Emphasis on the Gneissic Core Complexes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Adel A. Surour 17.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 17.2 Age of Arc Collisions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 17.3 Syn-Collisional (Arc-Related) Older Granitoids . . . . . . . . . . . . . . . . . 17.4 Gneissic Domes: Continental Windows (Infrastructure) versus Arc-Related Plutonism and Metamorphism . . . . . . . . . . . . . . . 17.5 Arc-Related Gneissic Complexes and Metamorphic History in the Sinai Peninsula, with an Emphasis on the Wadi Kid Environ . . . . . . . . . . . 17.6 Terrane Accretion in the Arabian Shield . . . . . . . . . . . . . . . . . . . . . . 17.7 Conclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

. . . .

18 Origin of the Volcanic-Arc Signature in Late-Orogenic Granitoids from the Arabian–Nubian Shield . . . . . . . . . . . . . . . . . . . . . . . . . . Basem Zoheir and Aliaa Diab 18.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 18.2 Tectono-Magmatic Evolution of the ANS . . . . . . . . . . . . . . . 18.3 Geochemical Characteristics . . . . . . . . . . . . . . . . . . . . . . . . . 18.4 Discussion and Conclusive Remarks . . . . . . . . . . . . . . . . . . . 18.4.1 Constraints from Zircon Hf Isotope Composition . . . 18.4.2 Hypothetical Geodynamic Model . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

. . . . . . .

. . 421 . . 421 . . 423 . . 425 . . 428 . . . .

428 433 434 434

. . . . . . . . 439 . . . . . . .

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19 Post-amalgamation Depositional Basins in the Arabian-Nubian Shield: The Hammamat Basins of Egypt . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Abdel-Rahman Fowler and Zakaria Hamimi 19.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 19.2 Geological Setting . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 19.3 Stratigraphy . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 19.4 Sedimentology, Facies, Environment . . . . . . . . . . . . . . . . . . . . . . 19.5 Provenance . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 19.6 Petrography . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 19.7 QFL Sedimentary Tectonic Environment . . . . . . . . . . . . . . . . . . . 19.8 Geochemistry . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 19.9 Structure of the Basins—the Effects of the Eastern Desert Shear Zone (EDSZ) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 19.10 Strain Studies . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 19.11 Basal Unconformity, Boundary Faults, Basin Shape . . . . . . . . . . . 19.12 Internal Deformation Structure—Folds and Faults . . . . . . . . . . . . 19.13 Metamorphism . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 19.14 Geochronology . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

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439 440 440 444 446 447 448

. . . . . 451 . . . . . . . .

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451 454 454 455 456 458 460 461

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464 465 467 468 469 470

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19.15 Discussion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 19.15.1 Relations of the Hammamat Basins to Each Other . . . . . . . 19.15.2 Information Based on Basin Size, Width and Thickness of Sedimentary Fill . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 19.15.3 Relations Between Basin Parameters and Degree of Basin Inversion for a Simple Graben Model . . . . . . . . . . . . . . . . 19.15.4 Subsidence and Inversion of the Larger Hammamat Basins 19.15.5 Tectonic Setting of the Hammamat Basins . . . . . . . . . . . . . 19.15.6 Future Trends . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 19.16 Conclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 20 Volcanism During the Post-accretionary Stage of the Arabian–Nubian Shield . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Mokhles K. Azer, Paul D. Asimow, and Simon A. Wilde 20.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 20.2 Geotectonic Evolution of the Arabian–Nubian Shield . . . . . . . . . . 20.3 General Classification of the ANS Volcanic Rocks . . . . . . . . . . . . 20.4 Post-accretionary Volcanic Sequences in the ANS . . . . . . . . . . . . 20.4.1 Early Calc-Alkaline Volcanic Sequences . . . . . . . . . . . . 20.4.2 Late Alkaline/Peralkaline Volcanic Sequences . . . . . . . . 20.5 Ages of Post-accretionary Volcanic Sequences in the ANS . . . . . . 20.6 Collision of East and West Gondwana . . . . . . . . . . . . . . . . . . . . . 20.7 Existence of Old Continental Crust Beneath the ANS . . . . . . . . . . 20.8 Examples of Post-accretionary Volcanic Sequences in the ANS . . 20.8.1 Early Calc-Alkaline Volcanic Sequences . . . . . . . . . . . . 20.8.2 Late Alkaline/Peralkaline Volcanic Sequences . . . . . . . . 20.9 Geodynamic Significance of ANS Post-accretionary VolcanoSedimentary Successions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 20.10 Conclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

. . . 471 . . . 471 . . . 471 . . . . . .

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474 474 476 477 478 479

. . . . . 485 . . . . . . . . . . . .

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485 487 489 490 490 491 492 494 495 496 496 511

. . . . . 520 . . . . . 522 . . . . . 523

21 An Overview Study of Zircon Geochronology from Sinai Precambrian Basement: Implications for Crustal Evolution of Northern Arabian-Nubian Shield . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Mahrous M. Abu El-Enen and Kamal A. Ali 21.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 21.2 Regional Geology of the Sinai Precambrian Basement . . . . . . . . . . . . 21.3 Data Compilation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 21.3.1 Magmatic Crystallization Ages . . . . . . . . . . . . . . . . . . . . . . 21.3.2 Maximum Depositional Age . . . . . . . . . . . . . . . . . . . . . . . . 21.3.3 Metamorphic Age . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 21.3.4 Inherited Zircon . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 21.4 Discussion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 21.4.1 Comparison of Zircon Ages from the Sinai Peninsula and Arabian-Nubian Shield . . . . . . . . . . . . . . . . . . . . . . . . . 21.4.2 Implications for the Evolution of the ANS . . . . . . . . . . . . . . 21.5 Conclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

. . 535 . . . . . . . .

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535 536 539 539 541 542 542 543

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543 543 546 554

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xxiii

22 Spectral Characteristics of Listvenites and Serpentineites Along Ophiolite-Decorated Megashears (Suture Zones) in the Arabian Shield Using ASD Fieldspec and Satellite Data . . . . . . . . . . . . . . . . . . . . . . . . . Ahmed A. Madani, Hesham M. Harbi, Abdel Moneim A. El-Dougdoug, Adel A. Surour, and Ahmed H. Ahmed 22.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 22.2 Geologic Setting . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 22.2.1 Jabal Ess Area . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 22.2.2 The Bir Umq Ophiolite Complex . . . . . . . . . . . . . . . . . . 22.2.3 Jabal Jizah . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 22.2.4 Jabal Tays—Al Amar Area . . . . . . . . . . . . . . . . . . . . . . . 22.3 Materials and Methods . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 22.3.1 Portable ASD FieldSpec Data Collection . . . . . . . . . . . . . 22.4 Results . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 22.4.1 Description of ASD Fieldspec Profiles . . . . . . . . . . . . . . . 22.4.2 Listvenites Discrimination Using Band Ratio Technique . . 22.5 Discussion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 22.5.1 Jabal Ess Area . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 22.5.2 Bir Umq Area . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 22.5.3 Jizah Area . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 22.5.4 Jabal Tays—Al Amar Area . . . . . . . . . . . . . . . . . . . . . . . 22.6 Conclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

. . . . 559

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23 Ore Deposits in the Arabian-Nubian Shield . . . . . . . . . . . . . . . . . . . . . . . Nagy Shawky Botros 23.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 23.2 Tectonomagmatic Evolutions of the ANS . . . . . . . . . . . . . . . . . . . . 23.3 Petrological Assemblages Characteristic for Mineral Deposits in the ANS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 23.3.1 Relicts of Oceanic Crust . . . . . . . . . . . . . . . . . . . . . . . . . . 23.3.2 Island Arc Assemblages . . . . . . . . . . . . . . . . . . . . . . . . . . 23.3.3 Mafic–Ultramafic Rocks . . . . . . . . . . . . . . . . . . . . . . . . . . 23.3.4 Felsic Plutonic Rocks . . . . . . . . . . . . . . . . . . . . . . . . . . . . 23.4 Mineral Deposits in the ANS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 23.4.1 Mineral Deposits Associated with Mafic–Ultramafic Assemblages . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 23.4.2 Mineral Deposits Associated with Metamorphosed Arc Volcanic-Volcaniclastic Rocks and Associated Arc-Related TTG Assemblages . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 23.4.3 Mineral Deposits Associated with Felsic Association . . . . . 23.5 Discussion and Conclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 24 Evolution and Mineralization of the Precambrian Basement of Yemen Salah Al-Khirbash, Mohamed Th. S. Heikal, Martin J. Whitehouse, Brian F. Windley, and Khaled Al Selwi 24.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 24.1.1 General Regional Setting . . . . . . . . . . . . . . . . . . . . . . . 24.1.2 History of Geological Investigations . . . . . . . . . . . . . . . 24.2 Precambrian Basement Distribution and Terrane Interpretation . . . 24.2.1 Basement Distribution . . . . . . . . . . . . . . . . . . . . . . . . . 24.2.2 Terrane Definitions . . . . . . . . . . . . . . . . . . . . . . . . . . . .

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559 560 560 561 562 563 564 564 565 565 568 574 574 575 577 577 579 580

. . . 585 . . . 586 . . . 587 . . . . . .

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589 589 589 589 591 592

. . . 592

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598 611 620 623

. . . . . 633

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633 633 635 637 637 638

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24.2.3 Proposed Wider Terrane Correlations . . . . . . Mineralization of the Precambrian Basement of Yemen 24.3.1 Metallic Mineralization . . . . . . . . . . . . . . . . . 24.3.2 Non-metallic Mineralization . . . . . . . . . . . . . 24.4 Conclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 24.3

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644 647 647 653 654 654

25 Applications of Remote Sensing and GIS in Geological Mapping, Mineral Prospecting and Groundwater Investigations in the Arabian-Nubian Shield: Cases from the Red Sea Hills of NE Sudan . . . . . . . . . . . . . . . . . . Khalid A. Elsayed Zeinelabdein, Eiman A. Mohamed, and Abdalla E. M. Elsheikh 25.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 25.2 Geological Mapping . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 25.2.1 Preamble . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 25.2.2 Office Work Phase . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 25.2.3 Field Work Phase . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 25.2.4 Post-field Phase . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 25.2.5 Geological Mapping of Kadaweib Area, Red Sea Hills, NE Sudan . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 25.3 Mineral Prospecting . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 25.3.1 Background . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 25.3.2 Mineral Resource of the Red Sea Hills of Sudan . . . . . . . . . 25.3.3 RS and GIS Prospecting in the Red Sea Hills of NE Sudan . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 25.4 Groundwater Investigations . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 25.4.1 Preamble . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 25.4.2 Rainwater Harvesting . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 25.4.3 Investigations for Groundwater in Fractured and Weathered Basement Rocks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 25.5 Conclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

. . 678 . . 683 . . 683

26 Phanerozoic Minor Volcanics and Intrusives of the Arabian-Nubian Shield . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Gehad M. Saleh, Mohamed S. Kamar, and Hamed I. Mira 26.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 26.1.1 Alkaline Ring Complexes Including Associated Volcanics 26.1.2 Gabal Elba and Gabal Shendeib Tertiary Alkaline Ring Complexes, Egypt (Case Study) . . . . . . . . . . . . . . . . . . . 26.1.3 Geological Tertiary Basaltic Rocks of the Arabian-Nubian Shield . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 26.2 Wadi Natash, South-Eastern Desert, Egypt of the Arabian-Nubian Shield . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 26.2.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 26.2.2 Geological Setting of Wadi Natash Rocks . . . . . . . . . . . . 26.2.3 Geology of the Rings . . . . . . . . . . . . . . . . . . . . . . . . . . . 26.2.4 Geochemistry . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 26.2.5 Radioelements Distributions . . . . . . . . . . . . . . . . . . . . . . 26.3 Felsite Dykes and Plugs of the Arabian-Nubian Shield . . . . . . . . . . 26.3.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 26.3.2 Felsites at Um Safi Area, CED, Egypt . . . . . . . . . . . . . . . 26.3.3 Felsites at Wadi Ranga . . . . . . . . . . . . . . . . . . . . . . . . . .

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659 662 662 662 663 663

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663 667 667 668

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672 676 676 676

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702 702 702 703 704 706 707 707 708 710

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26.3.4 26.3.5 26.3.6 26.3.7

Felsites at Wadi El-Miyah . . . . . . . . . . . . . . . . . . . . . . . . Felsites at Wadi Shait . . . . . . . . . . . . . . . . . . . . . . . . . . . . Felsites at Gabal Atalla . . . . . . . . . . . . . . . . . . . . . . . . . . . Geochemistry of Felsites Rock from the Eastern Desert, Egypt . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 26.3.8 Radioelements Distributions . . . . . . . . . . . . . . . . . . . . . . . 26.4 Trachyte and Bostonite Dykes and Concordant Sheets . . . . . . . . . . . 26.4.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 26.4.2 El Atshan Area, CED, Egypt . . . . . . . . . . . . . . . . . . . . . . 26.4.3 Wadi Kareem Volcanic Rocks, CED, Egypt . . . . . . . . . . . 26.4.4 Nasb El Qash Volcanic Rocks, CED, Egypt . . . . . . . . . . . 26.4.5 Um El Khors Area, CED, Egypt . . . . . . . . . . . . . . . . . . . . 26.4.6 Geochemistry . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 26.4.7 Um Doweila Bostonite, SED, Egypt . . . . . . . . . . . . . . . . . 26.4.8 Gabal Um Domi, SED, Egypt . . . . . . . . . . . . . . . . . . . . . . 26.4.9 Gabal Umm Salatit Area, CED, Egypt . . . . . . . . . . . . . . . . 26.4.10 Origin of Radioactivity at Bostonites and Trachytes Dykes, Eastern Desert of the Arabian-Nubian Shield . . . . . . . . . . . 26.4.11 Mineralogical Features . . . . . . . . . . . . . . . . . . . . . . . . . . . 26.5 Conclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 27 Relationship of the Pan-African Tectonic Structures with the Opening of the Afar Triple Junction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Jacques Varet 27.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 27.2 The Afar Triple Junction and Plume . . . . . . . . . . . . . . . . . . . . . . 27.2.1 The Afar Depression . . . . . . . . . . . . . . . . . . . . . . . . . . 27.2.2 The Surrounding Ethiopian Plateaus . . . . . . . . . . . . . . . 27.2.3 Pre-rift Sequences Along Afar Margins . . . . . . . . . . . . . 27.3 The Process of Development of the Continental Break-up in Afar . 27.3.1 Nature of the Main Ethiopian Rift: The Continental Rifting . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 27.3.2 Mantle Plume(s) and Development Steps . . . . . . . . . . . . 27.3.3 Determinant Tectonic Structures of the Pan-African Orogeny . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 27.3.4 Axial Ranges, Present Oceanic Axis of Spreading . . . . . 27.4 Transverse Volcanic Ranges and Marginal Central Volcanoes . . . . 27.5 Conclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

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About the Editors

Zakaria Hamimi is a Tectonics and Structural Geologist, currently at Benha University in Egypt, and has worked previously at Sana’a University, Yemen (1995–1998) and King Abdulaziz University, Saudi Arabia (2003–2013). He obtained a BSc (distinction with Honors) in 1984 at Assiut University, a M.Sc. in 1988 at Zagazig University and a Ph.D. in 1992 in Structural Geology and Tectonics at Cairo University. His research focusses on the tectonic evolution of the Arabian–Nubian Shield via structural/microstructural, palaeostress and strain studies, aligned with geological mapping, geomorphology and remote sensing methodologies. His research output exceeds 80 research publications in national and international journals. Besides, he co-edited four Springer books. He is the president, and a founding member of the Arabian Geosciences Union, since 2012. He has been awarded the medal of the Egyptian Geological Society of Egypt in 2015 and the medal of the Arab Mining and Petroleum Association in 2016. His services to the profession include his roles as Associate Editor of the Egyptian Journal of Geology (1998–2002) and the Arabian Journal of Geosciences (2016–now); member of the Egyptian Universities Promotion Committee, Supreme Council for Universities, Egypt (2016–2019); Secretary of the National Committee for Geological Sciences, Academy of Scientific Research and Technology (2016–2019), and later nominated for President of this Academic Committee in 2020; and the IUGS-Representative for Egypt in Cape Town 2016 and Thailand 2017. Abdel-Rahman Fowler graduated with B.Sc. (Honors, 1st class) in geology from Sydney University, Australia, in 1977. After experience in the uranium and tin exploration and mining industry in NSW, Queensland and Northern Territory, he obtained a Ph.D. in structure and tectonics at the University of New South Wales (UNSW) in 1986. He has instructed in structural geology and other geoscience disciplines at UNSW and LaTrobe University, Victoria, and has served at the United Arab Emirates University since 1999. His main interests lie in structural geology and tectonics, but extend to stratigraphy/sedimentology and geochemistry. He has conducted research work in Australia, Egypt, Turkey and UAE, and has published more than 40 papers in international journals and has contributed to books dealing with the geology of the Arabian–Nubian Shield. He has received awards for research excellence at UAEU. Jean-Paul Liégeois spent his whole academic career at the Royal Museum of Central Africa (RMCA), Tervuren, Belgium. He is an isotope field geologist who has mainly focussed on the geochemistry of granitic magmatism and the geodynamics of the Pan-African Orogeny in Sahara. He holds the M.Sc. (ULg, 1979) from the University of Liège and the Ph.D. from the University of Brussels (ULB, 1987). His Ph.D. was devoted to the transition from calc-alkaline to alkaline granitic magmatism in the Pan-African Orogen in the Adrar des Iforas, Mali, for which he received the “Prix Lucien Cahen” 1987 of the Belgian Overseas Royal Academy of Science. He worked in the joined RMCA-ULB Belgian Centre for Geochronology, which he co-directed for 20 years. He mostly worked in the Sahara from the Atlantic Ocean to the Indian Ocean but mainly in the Tuareg Shield, Algeria, Mali and Niger, xxvii

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privileging working in partnership with colleagues and students from the concerned countries. He strongly contributed to decipher the meaning of the high-K calc-alkaline and alkaline granitoids and the terrane structure in the Saharan Pan-African belt and to establish the existence of the Saharan metacraton, as well as the concept of metacraton. He also worked in other parts of Africa especially in the West African craton (Morocco, Mali and Niger) but also in Egypt, Sudan, Somalia, Cameroon and Central Africa. He worked punctually in Europe (Belgium, France and Romania) in terranes having links with the NW-African geology. If his personal contribution was field, geochronology, geochemistry and geodynamics, he always privileged collaboration with specialists in other disciplines, especially structural geology, metamorphic petrology, geophysics and sedimentology, generating a multidisciplinary approach. He has published 120 research articles in international indexed and refereed journals (h = 44) and co-edited several books. He was Regional Editor and Associate Editor of the Journal of African Earth Sciences for 25 years and is member of the Editorial Board of several international journals. He is Officer of the Crown (Belgium) and the laureate of the R. Shackleton Award 2008 of the Geological Society of Africa for “Outstanding research on Precambrian Africa”. Alan Collins is a tectonic geologist who is interested in how the deep earth evolution of the planet has controlled and governed earth surface systems (atmosphere, hydrosphere and biosphere). He is based in The University of Adelaide, Australia, and has worked over the last 20 years on unravelling the Ediacaran-Cambrian amalgamation of Gondwana, working initially in Madagascar, and later in East Africa, Arabia, Ethiopia, Brazil, India, Australia and East Antarctica. Using tectono-stratigraphic techniques, he has helped unravel much of the East African Orogen to produce a model for the collision of Neoproterozoic India with the central Africa. Mohamed G. Abdelsalam is a geoscientist interested in research and education in structural geology, geophysical applications and geospatial information sciences. His research focusses on the tectonic evolution of rift systems and orogenic belts, and morpho-tectonics. His research has been supported by the Petroleum Industry, National Science Foundation (NSF) and National Aeronautics and Space Agency (NASA) with a total of *4.5 million dollars (*50% share). Results of his research are presented in 103 publications (83 peer-reviewed), 91 invited presentations and 211 abstracts with a total citation of 3070, h-index of 32 and i10 index of 53 (Scopus, March 2020). He taught over 80 classes in the geosciences and geospatial information sciences at UT Dallas (August 2000–July 2006), Missouri S&T (August 2006–May 2012) and Oklahoma State University (August 2012– December 2019). Besides, he has mentored 12 Ph.D. and 14 MS students to successfully complete their graduate degrees. Currently, he is supervising three Ph.D. and two MS students. He served as the Geosciences Graduate Advisor at UT Dallas between August 2005 and August 2006, the Geology and Geophysics Graduate Coordinator at Missouri S&T between September 2006 and May 2009, and he has been serving as the Geology Graduate Coordinator at Oklahoma State University since July 2013. Additionally, he has been active in the US and African geoscientific communities. He served as proposals evaluator and manuscripts reviewer for many national and international funding agencies and journals. He served as the President of the Geological Society of Africa between 2001 and 2004, and currently, he is serving as the Co-editor-in-chief of the Journal of African Earth Sciences and as a member of the editorial board of the journal of International Geology Review. He was elected a fellow of the Geological Society of America in May 2008 and a fellow of the Geological Society of Africa in November 2008. He received the Missouri S&T Faculty Excellence Award in December 2008. He also received Oklahoma State University Outstanding Graduate Coordinator Awavrd in April 2017.

About the Editors

About the Editors

xxix

Mohamed Abd El‐Wahed is a structural geologist at the Geology Department, Faculty of Science, Tanta University, Egypt. He has graduated 1988 with honor degree and got the M.Sc. degree 1995 from the same department. He got his Ph.D. degree in 1999 from Department of Structural Geology, Institute of Geological Sciences, Wroclaw University, Poland. He joined the Geology Department, Faculty of Science, Omar Al Mukhtar University, Libya, during the period 2005 to 2012. He is a judge member of the Egyptian Promotion Committee. He has worked in many field‐related sub‐disciplines of Earth Sciences including structural and microstructural analysis, geologic mapping, strain analysis, tectonic geomorphology, crustal deformation and remote sensing. He used all these fields to study key areas in the Egyptian Nubian Shield and to decipher their deformation history. He has co‐published 55 research articles and book chapters in national and international indexed and refereed journals and authored three Arabic books.

1

The Arabian–Nubian Shield, an Introduction: Historic Overview, Concepts, Interpretations, and Future Issues Peter R. Johnson

Abstract

The Arabian–Nubian Shield (ANS) is widely viewed as a crustal block of juvenile Neoproterozoic rocks in Northeast Africa and the western Arabian Peninsula. Exposed in mountainous terrain on either side of the Red Sea and Gulf of Aden or concealed beneath Phanerozoic cover, the rocks make up Earth’s largest block of juvenile Neoproterozoic crust. The ANS has been geologically investigated for many decades from prior to the development of plate tectonics to the application of modern concepts and methods such as microprobe geochronology, isotopic analyses, and geochemical and petrologic techniques that allow sophisticated examination of suprasubduction and subduction-modified magmatic processes. During this time, interpretations of the ANS expanded from a simple inference of a single subduction zone across the entire region to the recognition of multiple arcs of different ages and tectonic settings. The ANS formed over a period of about 450 million years, between *1000 Ma and 530 Ma, and its growth reflects complex crustal evolution as part of the Rodinia–Gondwana supercontinent cycle. The shield originated with the rifting of Rodinia and the onset of intraoceanic arc magmatism in the Mozambique Ocean in parts of the ocean proximal to the Saharan Metacraton and the Congo–Tanzania Craton. Development of the shield continued with episodic convergence and accretion of arcs, syntectonic intrusion and metamorphism, and deposition of volcanosedimentary successions in basins developed on the newly accreted arcs. Shield development terminated with the growth of stable continental crust following a process of cratonization involving crustal thickening and metamorphism of accreted arcs; mantle delamination; asthenospheric upwelling and crustal melting; emplacement of an exceptional suite of P. R. Johnson (&) 6016 SW Haines Street, Portland, OR 97219, USA e-mail: [email protected]

late calc-alkaline to alkaline and A-type granites reflecting late- to post-tectonic and post-collisional settings; and accretion of stable crust to western Gondwana blocks. The relationship of the ANS in the western part of the Arabian Peninsula to Tonian arc and younger Cryogenian–Ediacaran mainly sedimentary rocks exposed in the eastern part of the Arabian Peninsula is not certain. The rocks in the east have a shared history with typical ANS rocks in having originated in the Mozambique Ocean, but may represent deposits proximal to the Indian Craton rather than to the Saharan Metacraton and Congo– Tanzania Craton. In central Egypt and Sudan, the ANS is in sheared contact with the Saharan Metacraton. However a clear distinction between the ANS and Metacraton as separate geologic entities is complicated by the presence of Ediacaran granitoid inrusions in the eastern part of the Metacraton implying extensive Neoproterozoic rejuvenation of Archean-Paleoproterozoic crust in the Metacraton contemporary with final events in the ANS. In southern Kenya and northern Mozambique, the ANS has a sheared contact with the Mozambique Belt, but the original depositional relationship between the ANS and Mozambique Belt is debated. Archean–Paleozoic rocks, part of the north-trending Azania ribbon continent, are present in Ethiopia, Somali, and Yemen and probably extend into central Saudi Arabia. They are in sheared or possibly depositional contact with typical ANS rocks on the east and the west suggesting that Azania was a relatively narrow continental protrusion into the Mozambique Ocean. As generally used, the term “Arabian–Nubian Shield” refers exclusively to juvenile Neoproterozoic rocks, but the presence of extensively reworked Archean–Paleoproterozoic crust on the flanks of, and as enclaves within, the region of typical ANS juvenile rocks suggests that Neoproterozoic tectonic processes affected not only the Mozambique Ocean but surrounding regions. In other words, the “story” of the ANS encompasses more than merely the

© The Author(s), under exclusive license to Springer Nature Switzerland AG 2021 Z. Hamimi et al. (eds.), The Geology of the Arabian-Nubian Shield, Regional Geology Reviews, https://doi.org/10.1007/978-3-030-72995-0_1

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P. R. Johnson

development of juvenile rocks within an ocean basin. On this basis, it appears to this author that discussion of the geology of the region would be better served by reverting to the original geographic meaning of the term “shield”, in which sense the ANS would be viewed as comprising a crustal block in Northeast Africa and Arabia composed of juvenile Neoproterozoic rocks (as its main constituent) as well as Neoproterozoic rocks containing variable amount of older continental material and Archean–Paleoproterozoic crust present as enclaves within or as continental margins on the flanks of Neoproterozoic crust. Keywords










Juvenile crust Neoproterozoic Isotopes Interpretation problems Tectonics

1.1




Terranes




Introduction

The Arabian–Nubian Shield (ANS) is part of an accretionary orogen exposed on either side of the Red Sea and Gulf of Aden in Northeast Africa (Nubian Shield) and the western part of the Arabian Peninsula (Arabian Shield) (Fig. 1.1). It consists of Neoproterozoic rocks ranging in age from *980 to 535 Ma (robust single-zircon U–Pb dating) and Archean– Paleoproterozoic structural enclaves. The Neoproterozoic rocks formed during a supercontinent cycle (Nance et al. 2014) bracketed by the 900–720 Ma break-up of Rodinia (Li et al. 2008) and the *650–530 Ma multiphase assembly of Gondwana concurrent with East African Orogeny (Collins and Piskarevsky 2005; Meert and Liberman 2008) (Figs. 1.2 and 1.3). During this period, the Earth system was affected by remarkable changes in biology, chemistry, and structure (Stern 2008a). These included the rapid evolution of eukaryotic organisms (Lipps and Valentine 2004); at least three episodes of continental glaciation that may have extended to low latitudes (Fairchild and Kennedy 2007); the reappearance after a gap of *1000 million years of sedimentary banded-iron formations (Ilyin 2009; Bekker et al. 2010; Cox et al. 2013, 2015); significant C, Sr, and O isotopic evolution and perturbations (Jacobsen and Kaufamn 1999; Shields et al. 2019; Halverson et al. 2007); the unambiguous establishment of plate tectonic processes (Stern 2008b); and the creation of juvenile crustal blocks, of which the ANS is the largest (Patchett and Chase 2002). The impact of these changes in the ANS is evidenced by fossils found in Ediacaran basins (600–570 Ma) in the northern Arabian Shield (Fig. 1.4a) indicating multicellular life forms (Cloud et al. 1979; Binda and Ramsay 1980; Vickers-Rich et al. 2013; Cui et al. 2020); deposits of banded-iron formation in the Eastern Desert, Egypt and Midyan terrane, NW Arabian Shield, of

which the most significant (Deposit 3, Sawawin) is as much as 58 m thick and contains a resource of 96 Mt grading 42.5% Fe) (Collenette and Grainger 1994) (Fig. 1.4b); and poorly-sorted *750 Ma cobble-boulder conglomerates (diamictites) stratigraphically below the BIF formations that contain a significant amount of Neoarchean and Paleoproterozoic material as detrital zircons and large clasts and may be evidence of Cryogenian glaciation (Ali et al. 2010). The opportunity to trace the record of these changes in the rocks of the ANS is one of the reasons why geologists are interested in the shield. Another reason is that the ANS is notably different from other Neoproterozoic crustal blocks, because it is well exposed, lacks significant vegetative cover or deep weathering, is only moderately covered by eolian deposits and alluvium, and in some places relatively little altered by deformation or metamorphism, so that the primary depositional, intrusive, structural, mineralogical, and chemical characteristics of the ANS rocks are well preserved (Fig. 1.5). In addition, the ANS is the most richly endowed Neoproterozoic block on Earth in terms of variety and quality of metallic mineral resources. These features make the ANS a prime target for geologic survey, research, and exploration. The shield is a world-class natural laboratory for testing well-established geologic theory and techniques, but its complexity and range of unresolved problems challenge the geoscience community to continue field observations, expand existing geochemical, isotopic, and geochronologic datasets, and develop new stratigraphic, structural, and tectonic ideas. The aim of this chapter is to introduce the ANS to a wide readership and to provide a context for later chapters in this book. The chapter is not a formal description of the geology and tectonics of the ANS—that is provided by recent publications such as Fritz et al. (2013), Johnson et al (2011), Nehlig et al. (2002), and Fowler and Hamimi (2020) and other chapters herein—but an overview that presents historic background to some of the main concepts that impacted our understanding of shield development, reviews current issues and problems in interpretation of shield geology, and outlines fields for future study. Topics presented in the chapter include a brief history of exploration and survey in the ANS; an outline of how plate tectonics frame our understanding of the development of the shield; a description of the shield margins; a consideration of some of the internal divisions within the shield; comments on topics warranting further study; and a note on the meaning of the term Arabian–Nubian Shield.

1.2

ANS: Plate Tectonic Setting

The ANS came into being because of plate tectonics—the shifting and rearrangement of crustal blocks in the Earth’s crust over time—and is being modified and broken up because of current plate tectonics. The ANS rocks formed in

1

The Arabian–Nubian Shield, an Introduction …

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Fig. 1.1 Map of Northeast Africa and the Arabian Peninsula showing the distribution of juvenile Neoproterozoic crust and adjacent regions of Archean–Mesoproterozoic crust (after Fritz et al. 2013) and Cenozoic plate boundaries, Red Sea-Gulf of Aden spreading centers, and the East African Rift system. The locations of geologic features (arcs, faults, etc.) mentioned in the text are indicated where possible. CED = Central Eastern Desert, NED = North Eastern Desert; SED = South Eastern Desert; SES = South Ethiopian shield; WES = West Ethiopian shield

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P. R. Johnson

Fig. 1.2 Schematic illustration of tectonic events in the Arabian–Nubian Shield associated with the Rodinia–Gondwana supercontinent cycle (after Stern and Johnson 2010)

the Mozambique Ocean (Dalziel 1991), an ocean basin between the Saharan, Congo–Tanzania, and Indian cratons created by Rodinia break-up (Fig. 1.3a). They include volcanic and tonalite-trondhjemite-granodiorite (TTG)-type intrusive rocks that formed as oceanic island arcs within the Mozambique Ocean and converged and amalgamated to create incipient ANS continental crust, post-amalgamation volcanosedimentary basins deposited on this incipient crust, and plutons and batholiths of late- to post-tectonic granitoids that intruded the developing ANS crust during a period of crustal thickening and orogenesis associated with closure of the Mozambique Ocean (Johnson and Woldehaimanot 2003; Nehlig et al. 2002; Fritz et al. 2013; Johnson et al. 2011). This process, reflecting bulk E-W shortening and N-S extension or

tectonic escape, accreted the ANS rocks to the flanking cratonic blocks during East African Orogeny (Collins and Piskarevsky 2005) and culminated in the assembly of supercontinent Gondwana and a network of Pan-African orogenic belts in the supercontinent. The ANS rocks and strongly deformed and metamorphosed Neoproterozoic and reworked Archean–Mesoproterozoic rocks in the Mozambique Belt to the south (present-day coordinates) constitute the East African Orogen (EAO) (Stern 1994), which is the largest of the Pan-African belts in Gondwana and a type of axial “mega-suture” between the Congo–Tanzania Craton, Saharan Metacraton, Madagascar, and India (Fig. 1.3b). The ANS underlies parts of the African and Arabian Plates (Fig. 1.1), and the Nubian Shield is being further

1

The Arabian–Nubian Shield, an Introduction …

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Fig. 1.3 Schematic maps showing: a the likely location of the Mozambique Ocean between rifting blocks of Rodinia and the relative locations of sites of oceanic arc magmatism forming the eventual Arabian–Nubian Shield and juvenile crust in Oman, at about 750 Ma (after Li et al. 2008), and b Gondwana after assembly of its constituent cratonic blocks at about 550 Ma illustrating the axial location of the East African Orogen (after Grey et al. 2008)

subdivided by the East African Rift System so that the eastern portion of the Nubian Shield is in the nascent Somalia Plate. The correlation of basement structures in Egypt and Sudan with equivalent geologic features in Saudi Arabia is clear proof that the Arabian and Nubian Shields were contiguous prior to Red Sea rifting at *25 Ma (see Fig. 1.1 in Stern and Johnson 2019) although the degree of

proximity of the African and Arabian Plates prior to Red Sea rifting is strongly debated (Bosworth 2015). Correlation of basement structures across the Gulf of Aden is not as well established, but there is general consensus that pre-Neoproterozoic crust in northern Somalia and northeastern Ethiopia broadly correlates with pre-Neoproterozoic crust in Yemen and Oman. The gneissic Al-Mahfid terrane

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Fig. 1.4 a Specimen of algal remains of Harlaniella ingriana collected from the Ediacaran Dhaiqa formation in the northwestern Arabian Shield (Vickers-Rich et al. 2013; Cui et al. 2020). b Outcrop exposure of banded-iron formation (BIF) at Sawawin 3, Midyan terrane, northwestern Arabian Shield

and juvenile Al-Bayda and Al-Mukalla arcs in Yemen have counterparts in the Mora-Qabri Bahar and Mait complexes in northwestern Somalia (Whitehouse et al. 2001a) and pre-rift models for the Gulf of Aden place Socotra, currently at the northeastern tip of the Somalia plate, adjacent to the Mirbat region of southern Oman (Denèle et al. 2012).

1.3

ANS: Geologic Investigations

1.3.1 Historic Developments Historically, the rocks of the shield have preoccupied rulers, explorers, empire-builders, miners, craftsmen, sculptors, antiquarians, natural historians, and scientists for more than 6000 years. Not only is the ANS the largest block of juvenile Neoproterozoic crust on Earth, it is also one of the most mineralized blocks. Periodically, with greater to lesser degrees in intensity of activity reflecting regional geopolitical changes, the shield has been and is a source of gold, copper, lead, and zinc and has resources of titanium, niobium and uranium and indications of platinum-group elements (Klemm

and Klemm 2013; Klemm et al. 2001; El Aref et al. 2020; Küster 2009). Ancient gold explorers and miners typically sought quartz veins, iron- and copper-stained rocks on either side of the veins, and gold-bearing gravel, exploiting deposits that today are called orogenic gold and alluvial gold (Johnson et al. 2017) and leaving behind hundreds of locations identified by pits, stopes, piles of waste rock, grinding stones, smelting sites, trenches and pits in wadis (intermittent drainage channels), and ruined miners’ houses and workshops. Mahd adh Dhahab, an epithermal gold deposit in Saudi Arabia, was first mined about 3000 years ago, again between 1250 and 750 years ago, later from 1939 to 1944, and is currently a principal gold mine of Ma’aden, the Saudi Arabian Mining Company (Lowther 1994). Elsewhere, deposits of sulfide minerals (volcanic-massive sulfides) and gold-rich oxidized weathered zones above the sulfides provided ancient miners with copper, lead, gold, and silver and are major sites of modern mining, as in the Ariab mineral belt, Sudan (Bosc et al. 2012) and Bisha deposit, Eritrea (Barrie et al. 2007). Historically important copper deposits episodically worked for the past 6000 years are found at Timna, southern Israel. These deposits are younger than the ANS because they are in Cambrian sandstone and siltstone and lower Cretaceous quartz-arenite immediately above the regional Early Cambrian planation surface that bevels the shield (Fig. 1.6) (Beyth et al. 2013), but together with the innumerable working on the shield testify that the Near East was one of the earliest regions of metallurgical developments in the world. Other ANS commodities produced in historic times included fine-grained silica-rich volcanic rocks that were worked into arrow points and other tools and granite and basalt for grinding stones that sustained the processing of metallic ore and preparation of food. Red to purple volcanic rock from the Dokhan Volcanics formation (*605–595 Ma; Wilde and Yoseff 2000) at Gebel Dokhan in northern Egypt was prized as “Imperial porphyry” for the construction of monuments and statues in Ancient Rome. The same rock was shipped throughout the Roman Empire and is available today via the Internet from Chinese companies processing stone imported from Egypt. Massive rose-colored granite from Aswan (Monumental Granite, 606 ± 2 Ma; Finger et al. 2008) was favored by the Pharaohs for the creation of large-scale monuments so that Aswan granite remains one of the most conspicuous and instantly recognizable tangible features of Ancient Egypt. The same Monumental Granite from the same ancient quarries is commercially available today as part of the modern Egyptian granite industry that has a production of about 400,000 tonnes annually (El-Haggar 2007). Written records of geologic observations in the ANS exist as early as 1256 in a manuscript by Al-Qastalani describing a volcanic eruption close to Madinah on the northern margin of Harrat Rahat, the largest field of Cenozoic Red Sea basalt

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The Arabian–Nubian Shield, an Introduction …

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Fig. 1.5 Rock types in the Arabian–Nubian shield. a Pillow basalt, Bahah Group, Asir terrane, Arabian Shield. Possible back-arc basin MORB-type basalt. b Refolded isoclinal fold in marble in the hanging wall of the Bi’r Umq suture, Jabal Tharwah, west-central Arabian Shield. c Diorite-tonalite gneiss, Dhara pluton, representative of *850 Ma pre- to syntectonic TTG intrusions in the Bidah belt, Arabian Shield. d Massive, undeformed *650–600 Ma granite pluton in the Haml Batholith, Khida terrane, eastern Arabian Shield. e Cyclic sedimentation of conglomerate, sandstone, and siltstone in the Ediacaran Jibalah group (*570 Ma), Antaq Basin, eastern Arabian Shield

in Saudi Arabia. The document records how “a fire burst out in the direction of Al-Hijaz; it resembled a vast city with a turreted and battlemented fort, in which men appear to be drawing the flame about, as it were, while it roared, burned, and melted like a sea everything that came in its way. Presently a red and bluish stream, bursting from it, ran

close to Al-Madinah” (see translation of the original in Camp et al. 1979). During the nineteenth century, as geology emerged as a distinct branch of science from the general study of natural history, it became commonplace to include geologic references and maps in publications about travels and antiquities in the region. An account of a visit to the

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Fig. 1.6 Planation surface on the ANS. a View of the Ram unconformity (*530 Ma) west of Al ‘Ula, Saudi Arabia, where Cambro-Ordovician sandstone and siltstone (Siq and Sak Formations) overlie the vast regional planation surface developed across the shield following exhumation and cratonization at the end of the Neoproterozoic cratonization. The surface is present around the northern, western, and eastern margins of the shield. In the southeast, in the vicinity of the Saudi-Yemen border, the Lower Paleozoic deposits in contact with the shield are referred to as the Wajid Supergroup. A broadly equivalent exhumed planation surface is present across the Nubian Shield in Egypt (Embabi 2018). b View of the the Najd Pediplain, Hijaz Plateau south of Biljurshi, Saudi Arabia. The plateau is the result of uplift of the Najd Pediplain (Brown 1960) during Red Sea rifting. Uplift formed the Red Sea Escarpment, a prominent erosional scarp 30–75 km inland from the Red Sea coast, which rises to elevation of >3000 m which extends from At Ta’if (Saudi Arabia) to east of Aden (Yemen). In western Saudi Arabia, the Najd Pediplain is locally unconformably overlain by Cenozoic basalt, implying the pediplain is at least early Cenozoic. A similar surface is buried beneath Oligocene Trap Volcanics in Ethiopia (Coltori et al. 2007). Other planation surfaces developed during accretion of the ANS, expressed as unconformities at the base of volcancosedimentary basins that formed following periods of exhumation and erosion of arc assemblages between *780 and 560 Ma

island of Socotra (Wellsted 1835) describes the geology of the central, highest part of the island. “The lower part of the range is composed of limestone, feldspar and porphyry, through which granite spires protrude themselves…the line of junction between the granite and limestone was beautifully exposed…elevated 3000 feet above where we stood”

(p. 172), referring to the 840 and 780 Ma granitoids that underlie Mount Haggier and the unconformably overlying Cretaceous limestone (Beydoun and Bichan 1969; Denèle et al. 2012). Stanley (1875) included a map of Sinai in a book about travels in Egypt, Israel, and Jordan, depicting areas of gravel, sand, vegetation, limestone, sandstone, and

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The Arabian–Nubian Shield, an Introduction …

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Fig. 1.7 Nineteenth-century geologic maps of a Sinai (Stanley 1875) and b northwest Saudi Arabia (Doughty 1888)

granite (Fig. 1.7a), the “sandstone” areas corresponding to exposures of the Feiran-Solaf and Elat metamorphic complexes (Abu El-Enen and Whitehouse 2013; Eyal et al. 2019) and “granite” representing the Ediacaran granite that underlies much of Sinai. A few years later, Doughty (1888) produced a map of the northwestern part of the Arabian

Peninsula (Fig. 1.7b) showing areas of granite and metamorphosed volcanic rocks belonging to the Arabian Shield and sandstone and limestone of the present-day Cambro-Ordovician Siq and Saq Formations and Permian Khuff Group, respectively, that are part of the Phanerozoic sedimentary succession unconformable on the shield.

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During the twentieth century, geologic investigation of the ANS became strongly developed, based on field and laboratory studies conducted by hundreds of geoscientists employed by national geological surveys, universities, and oil, gas, and mineral companies, as well as by contingents of colleagues from universities and institutions worldwide. The history of geologic endeavor in Egypt is covered in detail by El-Sharkawi et al. (2020) in the recently published “The Geology of Egypt” (Hamimi et al. 2020). Geologic mapping in Sudan began in 1905, the Geological Survey of Ethiopia was set up as a department within the Ministry of Mines in 1968, and the Saudi Geological Survey (SGS) was established in 2000. The SGS is the successor to 5 decades of work by the Saudi Arabian Deputy Ministry for Mineral Resources, which was assisted by geoscientists of the United States Geological Survey, the Bureau de Recherches Géologiques et Minières, Riofinex Limited and others. In addition to acquiring much of the data of these predecessor organizations, SGS also inherited the archives and data of earlier geoscience organizations that worked in the Kingdom, including the Japanese and Pakistan Geological Surveys, Seltrust, Boliden, and the Saudi-Sudanese Red Sea Commission. Among the oldest archival data held by the SGS, are materials produced by SAMS, the Saudi Arabian Mining Syndicate, which worked in Saudi Arabia between the late 1930s and 1954, conducting gold mining, particularly at Mahd adh Dhahab, and exploration.

1.3.2 Mineral Resources and Exploration Intense modern exploration and mining are underway in the ANS in Sudan, Egypt, Saudi Arabia, Eritrea, and Ethiopia, for gold, copper, lead, zinc, cobalt, tin, tungsten, titanium, and other metals from deposits of volcanic-massive sulfides (VMS), orogenic gold, intrusion-related gold, epithermal gold, porphyry copper, and Nb–Ta–U–REE-rich granite (e.g., Bielein et al. 2015, 2020; Botros 2004; El Aref et al. 2020; Helmy et al. 2004; Johnson et al. 2017; Zoheir 2012; Zoheir et al. 2019; Barrie et al. 2007, 2016; Collenette and Grainger 1994; Küster 2009). Sudan is the third largest gold producer in Africa after South Africa and Ghana, and a large example of modern ANS gold mining is located in the Ariab mineral district, in the Red Sea Hills (Bosc et al. 2012). Eritrea produces copper, zinc, gold, and silver from Bisha mine, a world-class high-grade VMS deposit (primary sulfide reserve of *9.5 Mt containing 1.05% Cu, 6.16% Zn, 0.68 g/t Au, and 45 g/t Ag) (figures effective Dec. 31, 2016: Bisha Mining Share Company website; accessed May 27, 2020). Active mines in Egypt include Sukari (gold), Abu Dabab, El Nuweiba, and Umm Naggat (tantalum). The Saudi Arabian Shield rocks are host to six gold projects (Ma’aden:

P. R. Johnson

web site information downloaded July 6 2020), one active base metal mine (Jabal Sayid VMS copper–lead–zinc–gold), and contain known resources at other VSM deposits such as Al Masane. Metallic exploration targets in the ANS include VMS in Egypt (Bampton 2017); porphyry copper in Sudan (Bielerin et al. 2015, 2020), Eritrea (Bournas et al. 2015) and Saudi Arabia (AMEInfo, Web information downloaded July 6, 2020); orogenic gold deposits throughout the region; and Ta–Nb–Li–Sn deposits in Late Cryogenian–Ediacaran granites particularly in Egypt and Saudi Arabia (Küster 2009).

1.3.3 Geologic Mapping, Surveys, and Findings Over the course of the twentieth and twenty-first centuries, geologic mapping of the ANS became increasingly detailed. Lithostratigraphy became better established, geochronologic, geochemical, and isotopic analyses became more accurate and routinely applied, and deep geophysical probing of the shield crust became commonplace. As a result, our present day geologic understanding of the shield is reasonably sophisticated and internally consistent. At the present time, much of the ANS is covered by detailed geologic maps, extensively covered by aeromagnetic geophysical surveys, and is partly covered by gravity measurements. The geology of the Arabian Shield in Saudi Arabia, for example, is covered by 0.5°  0.5° quadrangle maps at a scale of 1:100,000 and 1.5°  1° 1:250,000-scale maps. Airborne magnetic data cover the entire Arabian Shield and ground-based gravity data cover parts of the shield, the Red Sea coastal plain, and eastern Saudi Arabia. A country-wide geologic map of Egypt was published as early as 1910 as six sheets at 1:1 million scale and a reduced-scale version at 1:2 million (J.W.J. 1911), but current geologic information is more accessible as 20 map sheets at 1:500,000 scale (Conoco 1987). A geologic map at 1:2 million scale of Sudan (including what is now South Sudan) was published in 1981 (G.M.R.D. 1981), and more detailed mapping and geophysical surveys have been done in Sudan and South Sudan in conjunction with oil and metallic mineral exploration. A foundational geologic map of Ethiopia (including what is now Eritrea) was compiled and described by Kazmin (1972), a 1:2 million-scale geologic map of Ethiopia and Somalia based on mapping in 1973 was published by Merla et al (1979), and a more up-to-date 1:2 million-scale geologic map of Ethiopia was presented by Tefera et al. (1996). Eritrea is covered by a 1:1 million-scale country-wide map (Department of Mines 2009), and 1:250,000-scale mapping is underway. Detailed mapping has been done for northeastern Somalia by Fantozzi and Mohamed (2002). A new geologic map of Yemen was recently published (Albaroot et al. 2016) superseding earlier mapping by Beydoun (1966).

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The Arabian–Nubian Shield, an Introduction …

Modern study of the ANS reveals a history of interacting geologic processes including • prolonged episodes of intraoceanic and continental–margin subduction and arc formation • episodic deformation, burial, and metamorphism • late-to post-tectonic and post-collisional magmatism resulting from varied situations of delamination melting of lower crustal and subduction modification of upper mantle • orogenic extension • complex interactions with older continental crust on the margins of, and within, the shield as a source of clastic debris and inherited zircons • retention of isotopic features of older crust imprinted on the juvenile oceanic rocks • overlap of deformational, metamorphic, and magmatic events at the contacts between the ANS and the flanking cratons. Overall, the surface rocks of the Arabian Shield include about 31% volcanic and sedimentary rocks arranged in volcanic arcs and volcanosedimentary post-amalgamation basins (Johnson 2003), 26% granodiorite and tonalite (mostly as TTG arc-related assemblages), and 17% granite (Fig. 1.8a). Vertically, a seismic refraction survey (Fig. 1.8b) indicates that the Arabian Shield has a layered crustal structure and is 40–45 km thick, typical of continental crust (Mooney et al. 1985, 1998). The proportions of crustal rock types are probably similar in the Nubian Shield although crustal thickness is slightly less in Egypt than in Arabia (*35 km depth to Moho; Kaban et al. 2018). In northern Sudan and southern Gulf of Aden, the crustal thickness varies from a likely 35 to 21–16 km beneath the island of Socotra, at the northeastern corner of the Somali Plate (Ahmed et al. 2014).

1.4

ANS: Nomenclature

The name of the Arabian–Nubian Shield stems from early use of the term “le bouclier Arabe”—the Arabian Shield— by Karpoff (1957, p. 659) describing the crystalline basement of western Saudi Arabia (Fig. 1.9). The region has been also referred to as “le socle antécambrien de l’Arabie”, “le socle ancien”, the “Arabo-Nubian” or “Arabian–Nubian massif”, and the “Red Sea fold belt”, but “Arabian–Nubian Shield” is the term most commonly used today. The formal definition of a geologic shield is a “large area of exposed basement rocks in a craton, commonly with a very gently convex surface, surrounded by sediment covered platforms” (Jackson 1997). The first use of the English term was in a translation by Sollas and Sollas (1904) of Das Antlitz der Erde (The Face of the Earth; Suess 1892). The term “craton”

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refers to large stable portions of the Earth’s crust that have been little deformed for prolonged periods of time, periods that are generally taken to be Precambrian (Jackson 1997). In a formal sense, the term shield refers to geography (an area) not geology (rocks of a particular age or origin). Thus, if formally applied, the Arabian–Nubian Shield would refer to any crystalline basement rock of any age and tectonic setting in Northeast Africa and the Arabian Peninsula exposed beneath the Phanerozoic sedimentary and volcanic rocks of the surrounding platform areas. In this sense, the Arabian–Nubian Shield would include Neoproterozoic juvenile arc assemblages, Neoproterozoic gneisses and migmatites, Cryogenian–Ediacaran late- to post-tectonic granites, and Cryogenian–Ediacaran volcanosedimentary deposits that form at least the western part of the belt of juvenile crust depicted in Fig. 1.1, as well as Archean– Paleoproterozoic granites, gneisses, and schists exposed on the margins of the juvenile rocks in the Bayuda Desert, Sudan, and eastern Ethiopia-northwestern Somalia, and as enclaves within the juvenile rocks in the Jabal Khida area of central Saudi Arabia and the Abas and Al-Mahfid areas of southern Yemen. However, the term “shield” as currently used in the expression “Arabian–Nubian Shield” has evolved from an exclusive geographic connotation to become a quasi-tectonostratigraphic term referring to a particular assemblage of Neoproterozoic, mainly juvenile, rocks that formed in a particular period of time, in a particular tectonic setting. As commented later, the present author is inclined to view this as unnecessarily restrictive and suggests reverting to the geographic sense of the term. The term “shield” retains its geographic sense as the norm in descriptions of geology elsewhere in the world, referring to crustal blocks containing units of differing ages and origins. The Baltic Shield, for example, comprises Precambrian rocks exposed in Scandinavia that formed between 3.5 and 1.5 Ga; they make up three separate provinces of different lithologies and geologic histories but are all encompassed by the term Baltic Shield, meaning that the Baltic Shield refers to an area of Precambrian rocks of varied origins not to a specific geotectonic assemblage. The Canadian Shield, likewise, includes rocks of varied ages and origins forming separate geologic provinces (e.g., the Superior Province) exposed over a wide geographic area in the Canadian portion of the Laurentian craton, the continental crust of North America (Bastedo and James-Abra 2006). The Brazilian Shield comprises three major pre-Neoproterozoic tectonic domains—the Amazon, São Francisco, and Rio de la Plata Cratons—as well as Neoproterozoic rocks (900–550 Ma) derived by reworking of older crust and to lesser extent by juvenile processes in belts surrounding the cratons (Hartmann and Delgado 2001). Traces of the formal (geographic) use of shield in Arabia and Northeast Africa are occasionally found, as for example,

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Fig. 1.8 a Structure of the lithosphere beneath the Arabian–Nubian Shield, based on geophysical data for the Eastern Desert Egypt, and pie diagrams illustrating the proportions of rocks types exposed at the surface of the shield (after Stern 2018 and Johnson and Kattan 2012). b Interpretative crustal section along the 1978 seismic refraction profile, across the southern part of the Arabian Shield (Mooney et al. 1985), based on P-wave velocity structure

in a report on geochronology and isotopes in southern and eastern Ethiopia (Teklay et al. 1998, p. 224) that “The zircon ages and Nd isotope data add further evidence to the concept

that the ANS consists of juvenile as well as ancient terranes which were swept together and amalgamated during the Pan-African Orogeny in Neoproterozoic times”. But in most

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The Arabian–Nubian Shield, an Introduction …

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Fig. 1.9 Map from the 1957 paper by Roman Karpoff describing the geology of the Arabian Shield observed by him during the course of making traverses across the Kingdom. The map shows the bulk of the shield as “Série de Madinah”, small exposures of “Série du W. Fatima” close to Jiddah, and overlying areas of “Basaltes récents”

current literature, the term “Arabian–Nubian Shield” preferentially refers to rocks of Neoproterozoic age exposed in Northeast Africa and the western Arabian Peninsula as a tectonostratigraphic entity composed of (1) volcanic and intrusive assemblages that formed in arcs and back-arcs in the Mozambique Ocean; (2) schist and gneiss derived from the volcanic and intrusive rocks; (3) volcanosedimentary successions deposited in basins of different origins on the accreted arcs; and (4) vast amounts of late- to post-tectonic granitic intrusions. An increasingly large dataset of single-zircon crystallization/rock-forming ages demonstrates

that these rocks range from *1000 Ma (Stern and Manton 1987; Eyal et al. 2014) to *525 Ma (Robinson et al. 2014). Sr, Nb, Hf, and O isotopic data demonstrate that the bulk of these rocks are juvenile, formed of material extracted from depleted mantle (Fig. 1.10), and U–Pb zircon dating indicates that most crystallization ages of the rocks are relatively close to their Nd-model ages, so much so that Stern (2002a, p. 115) floats the idea that the coincidence of the traditional limits of the Arabian–Nubian Shield and the area in Northeast Africa and Arabia containing Late Mesoproterozoic-Early Neoproterozoic Nd-model ages in the region amounts to a

14 Fig. 1.10 Age versus eNd and eHf diagrams showing the fields typical of juvenile Neoproterozoic rocks in the Arabian–Nubian Shield and fields typical of rocks from older continental crust or Neoproterozoic granites with older crustal components in Yemen and the Khida terrane, eastern Arabian Shield

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The Arabian–Nubian Shield, an Introduction …

“new way to define the Arabian–Nubian Shield”. Former arguments for the ANS being underlain to a variable degree by continental crust are now abandoned, but nevertheless, there is a significant body of isotopic and geochronologic data (see below) evidencing continental crust involvement in the formation of ANS Neoproterozoic volcanic and intrusive rocks as a component of magmatic melts from the mantle or by assimilation of sedimentary material (e.g., Hargrove et al. 2006b; Stern et al. 2010; Li et al. 2018).

1.5

ANS: External Boundaries

Assessment of the size and location of the Arabian–Nubian Shield depends on the definition of the shield. Delfour (1975) refers to the Arabian–Nubian Shield as underlying parts of Egypt, Sudan, Ethiopia, North Somalia, and the western part of the Arabian Peninsula. Johnson and Woldehaimanot (2003) effectively limit the shield to areas of essentially continuous basement exposures in Jordan, Saudi Arabia, Yemen, Egypt, Sudan, Eritrea, and northern Fig. 1.11 Side-by-side comparison of the sizes of the East African Orogen and Alpine-Himalayan Orogenic belt

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Ethiopia. Vail (1985) and Leonoir et al. (1994) include a similar area in the ANS but extend the southern limit to include juvenile crust in southern Ethiopia. Crystalline basement in Ethiopia is extensively covered by Cenozoic basalt of the Ethiopian Plateau, and windows of basement crust in western and southern Ethiopia are referred to as shields in their own right (Western Ethiopian Shield; Southern Ethiopian Shield) (Fig. 1.1). The formative paper on the East African Orogen (Stern 1994) infers that the ANS is a belt of juvenile crust extending from Sinai and southern Jordan as far south as Tanzania. A similar interpretation of regional geology was made by Fritz et al. (2013), in designating all juvenile crust between Sinai and Jordan in the north and Tanzania (including the shield areas of Ethiopia) in the south as ANS crust. In this view, the exposed ANS extends about 4000 km north–south and 1200 km east–west at its widest in North Africa and the Arabian Peninsula (allowing for virtual coast-to-coast closure of the Red Sea). The ANS and Mozambique Belt have a combined north– south length of about 7000 km, about half the length of the well-known Alpine-Himalayan orogenic belt (Fig. 1.11).

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1.5.1 ANS Northwestern Margin The northwestern margin of the ANS is an isotopic, geochronologic, and structural boundary broadly coincident with the Keraf suture (Abdelsalam et al. 1998), a north-trending zone of folding, shearing, and west-vergent ophiolite-decorated thrusts in the eastern Bayuda Desert, Sudan that separates Neoproterozoic juvenile rocks on the east from reworked Archean to Mesoproterozoic continental crust of the Saharan Metacraton on the west (Almond and Ahmed 1987; Abselsalam et al. 1998; Stern 1994). The suture is conventionally projected north across the basement rocks of northern Sudan toward Aswan, in Egypt, then farther north beneath Phanerozoic cover and recent alluvium toward the Nile Delta. Early papers (e.g., Greiling et al. 1994) placed the boundary east of the granitic basement rocks of Aswan, but evidence of a volcanic-arc signature for tonalitic gneiss at Aswan (Finger et al. 2008) suggests that the boundary may be farther west, as adopted in Fig. 1.1. Southward, the suture projects toward Sabaloka, in central Sudan, and then farther south toward the Nuba Mountains in southern Sudan. The Saharan Metacraton is a vast region of Archean and Mesoproterozoic migmatite gneiss and granite cratonic rocks in north-central Africa between the ANS and the Tuareg Shield of the West African Craton that were reworked during Neoproterozoic orogenic events and intruded by Neoproterozoic granites (Harms et al. 1990; Abdelsalam et al. 2002; Zhang et al. 2019a, b). Ediacaran granitoids (*628–580 Ma) emplaced in southern Egypt from the Libyan border to the Nile River (not shown on Fig. 1.1) constitute a transition between rejuvenated Archean–Paleoproterozoic crust of the Saharan Metacraton and the ANS (Zhang et al. 2019b). Metacraton basement rocks in the Bayuda and Nubian Deserts immediately west of the Keraf suture include Late Archean to Mesoproterozoic medium- to high-grade metasedimentary schists, paragneiss, and orthogneiss in the Delgo-Wadi Halfa area (Stern et al. 1994) and Paleoproterozoic to early Neoproterozoic quartzofeldspathic gneiss, mica schist, amphibolite in the Rahaba-Absol terrane (Küster et al. 2008) reworked during magmatism and metamorphism ending with the Bayudian Event at 920–900 Ma (Küster et al. 2008). The rocks have strongly negative Nd initial ratios (eNd = −21 to −7) and Archean to Mesoproterozoic Nd-model ages, evidencing that they belong to a pre-ANS crust (Stern et al. 1994). Along and east of the Keraf suture in the Abu Harik-Kurmut and Baileteb areas, and along the Atmur-Delgo suture, which projects west as a re-entrant into the Saharan Metacraton, are younger biotite gneiss, biotite-hornblende gneiss, amphibolite, serpentinite and pillow basalt, as well as lower grade chlorite schist, chlorite-actinolite schist, calc-silicate rocks, and marble (Harms et al. 1994; Ali et al. 2013a; Bailo et al. 2003; Evuk

P. R. Johnson

et al. 2014; Karmakar and Schenk 2015). They have juvenile Pb- and Sr-isotope signatures (Bailo et al. 2003; Evuk et al. 2014), positive Nd initial ratios (eNd = +4.2 to +6.8), and early Neoproterozoic model ages (1.28–0.75 Ga) (Stern et al. 1994; Harms et al. 1994; Shang et al. 2010) indicating derivation from juvenile mantle, for which reason these rocks are deemed to be part of the ANS. They include passive-margin epiclastic and carbonate and subductionrelated volcanosedimentary sequences and are interpreted as the westernmost deposits of the ANS oceanic basin in Sudan. Amphibolite-facies metamorphism at *670 Ma (Karmakar and Schenk 2015) was associated with *715–630 Ma north–south closure and collision of the Rahaba-Absol and Kurmut-Abu Harika terranes (Evuk et al. 2014); subsequent folding and shearing resulted from *650 to 600 Ma sinistral transpression; and final collision between East and West Gondwana occurred by *580 Ma (Abdelsalam et al. 1998). The Sabaloka basement complex is likely at or close to the Saharan Metacraton-ANS contact. Covering an area of about 3000 km (Almond 1980), the complex includes quartzofeldspathic biotite gneisses and minor metapelitic and metaigneous granulites, a younger set of high-K granite batholiths intruded about 590–600 Ma (Küster et al. 2008), and mid-Paleozoic ring complexes (Karmakar and Schenk 2015). The metapelitic granulite yields Archean–Paleoproterozoic detrital zircons and a 1.70 Ga Nd-model age (Kröner et al. 1987) suggesting that the sedimentary protolith of the granulite had a continental source (Karmakar and Shenk 2015). Overall however, the rocks of the complex are inferred to represent deposition at a passive margin on the Saharan Metacraton that ended about 780–720 Ma at a time of high-temperature crustal thickening followed by a high to ultrahigh-temperature decompression overprint at *602 Ma, likely due to the final collision of the Saharan Metacraton and the ANS in central Sudan (Karmakar and Shenk 2015). A continuation of the contact between the Saharan Metacraton and ANS is inferred in the Nuba Mountains, in southern Sudan. The mountains contain metavolcanosedimentary rocks; ophiolitic serpentinite, gabbro and pillow basalt; ortho- and paragneiss; and syntectonic and anorogenic intrusions. The mountains rise 400–900 m above Cretaceous sandstone and Nile valley alluvium in the surrounding lowlands and extend for about 150 km east–west and 65 km north–south. The Kabus suture (Fig. 1.1), an ophiolitic mélange in a north-trending shear zone about 10 km wide in the eastern part of the mountains, is the conventional ANS boundary. The Kabus suture contains serpentinite, talc schist, metagabbro, and amphibolite in east-vergent thrusts and separates high-grade gneisses to the west from low-grade volcanogenic oceanic assemblage of ANS character to the east (Abdelsalam and Dawoud 1991).

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The Arabian–Nubian Shield, an Introduction …

However, low-grade metamorphosed mafic intrusive and extrusive rocks, minor ultramafic rocks, and units of volcanosedimentary rocks in the Abu Zabed area, in the western part of the Nuba Mountain (Vail 1985), also have juvenile geochemical and geochronologic signatures. The rocks yield a *778 Ma Sm–Nd isochron age, a 814 Ma Nd-model age, and positive eNd(t) (+5.9) suggesting Neoproterozoic derivation from a depleted mantle source (Mohammed 2017). High-grade gneiss and synorogenic granitoids in the same area have Sm–Nd isochron ages of *980–975 Ma and a positive eNd value of +6.3, likewise suggesting a juvenile origin, and granite plutons in the area yield Ediacaran crystallization ages of 625–605 Ma and indications of 614–583 Ma high-grade metamorphism (metamorphic rims on zircon). It appears therefore that oceanic juvenile and metamorphic rocks of Neoproterozoic age are present in the western Nuba Mountains, implying that any ANS-Saharan Metacraton contact is farther west (Mohammed 2017).

1.5.2 ANS Southern Margin As described more fully elsewhere (Fritz et al. 2013) and in Chap. 27 in this book (Fritz and Hauzenberger, in press), the ANS tapers southward as it approaches the Mozambique Belt. In southern Kenyan, the ophiolite-decorated Sekerr-Athi shear zone (Berhe 1990; Frisch and Pohl 1986) juxtaposes the ANS rocks with the Archean Congo– Tanzania Craton and Mozambique Belt Eastern Granulite Complex (Fritz et al. 2009; Hepworth 1972). In the east, the Galana shear juxtaposes ANS rocks with Azania gneiss (Fig. 1.1). Farther south, the bounding shears become merely a few kilometers apart so that the ANS appears to pinch out, before extending as a composite shear zone toward correlative shears in Madagascar (Fritz et al. 2013; Fritz and Hauzenberger, in press).

1.5.3 ANS Eastern Margin The eastern margin of the ANS is not well established because extensive cover by Cenozoic basalt and unconsolidated alluvium and eolian deposits makes basement exposures discontinuous. High-grade gneiss and schist and younger granitic plutons in Somalia and eastern Ethiopia are inferred by Lenoir et al. (1994) to be a northeastern branch of the Mozambique belt, although more recently Collins and Windley (2002) correlated them with Archean–Paleoproterozoic crust extending north from central Madagascar as the Azania block (Collins and Piskarevsky 2005) linking northern Somalia with Madagascar rather than with the Mozambique Belt. Geochronologic and structural evidence suggests that Azania separated as a ribbon continent from the Congo–

17

Tanzania Craton sometime after deposition of the 1.7–1.5 Ga Itremo Group in Madagascar, leaving a back-arc basin between Azania and the Congo–Tanzania Craton, which effectively evolved as a re-entrant of the Mozambique Ocean in which juvenile Neoproterozoic rocks of southern Ethiopia, Kenya, and northern Mozambique were deposited (Collins and Piskarevsky 2005; Fritz and Hauzenberger, in press). Weak evidence of Azania in southern Ethiopia derives from an undeformed 568 ± 9 Ma Ediacaran granite that contains *2.5 Ga xenocrystic zircons, suggesting the granite melt incorporated *2.5 Ga crust or a sedimentary rock rich in Archean detrital zircon (Stern et al. 2012). Neoproterozoic tonalite, monzogranite and alkali-feldspar granite exposed in the Dire Dawa-Harar area of northeastern Ethiopia have d18Ozrn of 4.9–9.6‰, higher than mantle values, strongly negative eNd(t) values of −10.3 to −5.8, and Nd-model ages of 1.72–1.42 Ga, likewise suggesting a pre-Neoproterozoic continental crustal component (Yeshanew et al. 2016). Paleoprotereozic zircon xenocrysts (1.82 and 1.73 Ga) in Neoproterozoic gabbro (814 Ma), 1.82– 1.71 Ga xenocrysts in 808 Ma sillimanite-bearing paragneiss, and a 1.40 Ga xenocryst in 718 Ma orthogneiss are interpreted as evidence of a mid-Proterozoic crust at depth contributing to Neoproterozoic granitoid melts or providing a source of detritus for paragneiss protoliths in the the Qabri Bahar and Mora complexes a few hundred kilometers east of Harar in northwestern Somalia (Kröner and Sassi 1996). Farther east, the Somalia basement passes into the Mait complex, a greenstone belt composed of low-grade Neoproterozoic pillow lava, microgabbro, and greenschist-facies metasedimentary rocks (Utke et al. 1990; Lenoir et al. 1994) overlain by low-grade metasedimentary rocks of the Inda Ad Complex (Lenoir et al. 1994). Post-tectonic granodiorite dated at 626 ± 11 Ma provides a maximum deposition age for the Inda Ad Complex and a possible xenocrystic zircon of 987 Ma suggesting a contribution from significantly older crust (Lenoir et al. 1994). Deposition of the Ediacaran Inda Ad and Mait Complexes effectively defines the eastern margin of Azania in northern Somali and implies that a juvenile oceanic basin occurred east of Azania and that Azania was truly a ribbon continent.

1.5.4 ANS Margins in the Southwestern Arabian Peninsula A northward continuation of the western edge of Azania is projected into the heterogeneous basement of southern Yemen where gneissic terranes are interspersed with Neoproterozoic greenschist terranes (Windley et al. 1996; Whitehouse et al. 2001a). Old continental crust is evidenced in the Al-Mahfid terrane by late-Archean Nd-model ages and U–Pb zircon ages of 2.95 and 2.55 Ga. In the Abas terrane, a

18

single late-Archean zircon core, eNd(t) whole-rock values between −11 and +0.8, and Nd-model ages of 1.70–1.13 Ga in Neoproterozoic gneiss and granite are consistent with Neoproterozoic reworking of late-Archean crust (Yeshanew et al. 2015). Other rocks in these terranes include 625– 590 Ma post-tectonic granites that yield eNd(t) of −11 to +0.8 indicative of Neoproterozoic reworking and granite gneiss with *790–725 Ma protolith ages (Whitehouse et al. 2001a; Yeshanew et al. 2015). The island-arc Al Bayda terrane yields K-Ar and 40Ar–39Ar ages between 823 Ma and 614 Ma, which, although not robust, suggest a Neoproterozoic origin. Nevertheless, Paleoproterozoic Nd-model ages (2.53–1.99 Ga) (Windley et al. 1996), indicate the Al-Bayda island-arc assemblage also received significant ancient continental input. The northward extent of the Archean–Paleoproterozoic crust of Yemen is unknown because of cover by sedimentary rocks of the Arabian Platform, but Paleoproterozoic crust reappears in the Jabal Khida area at the eastern edge of the exposed Arabian Shield in Saudi Arabia (Fig. 1.1) (Stoeser and Stacey 1988; Stoeser et al. 2001; Whitehouse et al. 2001b). At Jabal Khida, U–Pb microprobe dating identifies a Paleoproterozoic age for the Muhayil granite (*1660 Ma) and 1.8–1.7 Ga zircon cores in the Fuwayliq granodiorite and Libab granite gneiss (Whitehouse et al. 2001b). The South Libab gneiss has even older 2.4–2.1 Ga zircon cores and 1.9–1.75 Ga rims suggesting a Paleoproterozoic metamorphic event. Outcrops of these old rocks are limited to an area of 68 wt%

Potassic/shoshonitic Domain

Highly fractionated CA & Alk

Nuba Mts postcollisional granite

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Fig. 2.8 Plot of major and trace elements of the late-orogenic granite samples in various discrimination diagrams. a #Fe* versus SiO2 of Frost et al. (2001). b The HFSEs versus FeO*/MgO, diagram of Whalen et al. (1987) that classify the samples as A-type granite. c The granitic classification scheme of Sylvester (1989). d NYTS diagram (normalized to Yenchichi–Telabit series; sliding normalization; Liégeois et al. 1998) samples from the Nuba plutons all plot in the potassic domain. The Nabati analysis is from Küster et al. (2008). Same symbols as in Fig. 2.7

plot in the potassic domain and are not related to an alkaline series (Fig. 2.8d). In the sliding normalization, each studied rock is normalized to the interpolated composition of the reference series that has the same SiO2 content as the sample, here the alkali-calcic Yenchichi-Telabit Series (Liégeois et al. 1998). This method amplifies differences in source compositions and in fractionation processes and allows comparison of rocks from basic to acid composition. Alkali feldspar granite and syenogranite samples have moderate to enriched ƩREE values (106 ppm to 657 ppm), a steep REE pattern ((La/Yb)N between 17 and 77) and a slightly negative Eu* anomaly (0.71–0.89), with sample Sk5 showing a more pronounced negative Eu* anomaly of 0.71 (Fig. 2.9a). This is strongly different from the REE pattern of the c. 778 Ma Arid subduction-related volcanic rocks (grey pattern; Ibinoof et al. 2016). In the primitive mantle-normalized multi-element spidergram (McDonough and Sun 1995), all of the Nuba granites show pronounced Ta and Nb negative anomalies, depletions in Sr, P2O5, Y and Ti and an enrichment in K2O and Rb (Fig. 2.9b), where the 500

samples cluster in the field of alkalic granite field with a slight overflow in the alkali-calcic field, reflecting a high total alkali content (Na2O + K2O = 9.1%–10.3%). By contrast, they are characterized by low to very low contents of TiO2 (0.16–0.44%), MgO (0.07–0.30%) and CaO (0.34– 1.32%). The two granitic types are slightly to moderately peraluminous, with A/CNK ratio (molar Al2O3/CaO + Na2O + K2O) ranging from 1.01 to 1.14 (Fig. 2.7d), which is due to quite high Al2O3 content for granites (14.2% to 16.0% + 1 sample at 18.7%), which is not characteristic of alkalic granites. The studied samples grouped in the field of ferroan granite on the #Fe* versus SiO2 diagram of Frost et al. (2001; Fig. 2.8a). Using the HFSE versus K2O/MgO diagram (Whalen et al. 1987), the samples split between the fields of fractionated and A-type granite (Fig. 2.8b). In the discrimination diagram of Sylvester (1989) for high-silica samples (SiO2 > 68%), the studied Nuba samples fall in the field common to the alkaline granites and to the highly fractionated calc-alkaline granites (Fig. 2.8c). In the NYTS diagram (normalized to Yenchichi-Telabit series; sliding normalization; Liégeois et al. 1998), the samples all

FeOt/MgO

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M. A. Ibinoof et al. 10

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Fig. 2.9 a Chondrite-normalized REE diagram for late-orogenic granite from the western Nuba Mountains. Normalizing values are from Sun and McDonough (1989). The grey shaded area represents samples from the metavolcanic rocks of the Nuba Mountains (Ibinoof et al. 2016); b Primitive mantle-normalized trace element diagram for late-orogenic granite from the western Nuba Mountains. Normalizing values are from McDonough and Sun (1995). The grey shaded area represents samples from the metavolcanic rocks of the Nuba Mountains (Ibinoof et al. 2016). c N–MORB-normalized trace element diagram for late-orogenic granite from the western Nuba Mountains. Normalizing values are from Sun and McDonough (1989). The grey shaded area represents samples from the metavolcanic rocks of the Nuba Mountains (Ibinoof et al. 2016)

2

The Boundary Between the Saharan Metacraton …

shaded area represents the metavolcanic rocks of Arid Unit, here also strongly different (c. 778 Ma; Ibinoof et al. 2016). Their REE pattern and MORB normalized spidergrams show a close resemblance to the Nabati pluton of the Bayuda desert of northern Sudan, which is related to the subduction– related continental margin rhyolites and rhyodacites (Küster et al. 2008). The presence of some Nuba samples in the NYTS “undifferentiated domain” (Fig. 2.8d), which includes the subduction-related granitoids is along the same lines.

2.7

Sr and Nd Isotopes

Four Nuba granitic samples (SK4, SK5, SK224 and SK275) were selected for Rb/Sr and Sm/Nd isotope studies. The two samples ZR1 and ZR2 selected for zircon separation and SHRIMP dating (see above), correspond to SK4 and SK5 samples.

2.7.1 Analytical Techniques Neodymium and strontium isotope analyses were conducted at the Department of Geological Sciences, University of Cape Town, South Africa. Sample powders were dissolved using HF–HNO3 acid mixture in closed Teflon beakers on hotplates at 140 °C for two days. Any minor undissolved material was removed by centrifuging the samples prior to sequential Sr and Nd separation chemistry (Pin et al. 1994; Míková and Denková 2007; Pin and Zalduegui 1997). Final Sr and Nd fractions were analysed for isotopic compositions using a Nu Instruments Plasma HR–MC–ICP–MS instrument coupled with a DSN-100 desolvating nebuliser. The external 2 sigma errors for the measured Sr and Nd isotope ratios are better than ± 0.000015. The BHVO-2 basaltic standard reference material yields values of 0.703530 ± 0.000009 for 86Sr/88Sr and 0.513005 ± 0.000009 for 134Nd/144Nd, comparing well with published data (Weis et al. 2006). Concentrations of Rb, Sr, Sm and Nd were measured using ICP–MS. The error on Rb/Sr and Sm/Nd is 15%, Mg * 3%, Y < 18 ppm; Sr/Y * 20–100, Stern 2002). The REE patterns have a gentle slope *10 with a few samples having a positive anomaly indicative of plagioclase cumulate (Fig. 5.6b). The spider plot displays a Nb–Ta anomaly consistent with a subduction-related environment for the igneous protolith of these amphibolites (Fig. 5.6c). Obeid and Azer (2015) described a sequence of basalt, basaltic andesite, and dacite as the lower part of Dokhan volcanics with typical adakitic affinity from the northeastern desert of Egypt. They concluded that the magma of this sequence was generated through partial melting of delaminated mafic lower crust interacting with overlying mantle-derived magma.

Fig. 5.5 Field photos of the Janub Metamorphic Suite. a Meta-conglomerates of the JMS. The arrows point to examples of the variably stretched granitic pebbles in the meta-conglomerates. b Contact between Ayn Al Hashim cordierite-biotite hornfels and the intruding Abu Jedda granite (AJ granite) from Yutum Suite; the sedimentary compositional layering is well preserved and dips to the NW at variable angles. c Agglomerates from the meta-pyroclastics of the JMS

5.2.2 Gabbroids and Granitoids Most of the rocks in the Aqaba Complex are felsic (granitoids) in composition with subordinate mafic to intermediate (gabbroids) varieties (Table 5.1). The suites encompass more

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Neoproterozoic Crustal Evolution …

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characterize the different suites. On the total alkali silica diagram of Le Maitre et al. (1989, Fig. 5.7a), and the high field strength classification system of Winchester and Floyd (1977, Fig. 5.7b), the intrusive suites of Aqaba Complex plot in the fields of gabbro, diorite, granodiorite, and granite. The majority have medium to high-K calc-alkaline affinity, with the exception of Duheila Hornblendic Suite (DHS) of calcic affinity (Fig. 5.7c). Furthermore, these intrusives possess a magnesian affinity except for Yutum, Darba, and Rumman suites which show both magnesian and ferroan affinity in the sense of Frost and Frost (2008; Fig. 5.7d).

Fig. 5.6 Abu Saqa Schist. a Field photo of Abu Saqa amphibole schist and conformable granitic gneiss within. b Plot of chondrites normalized REE content of Abu Saqa schist. Chondrite values are after Sun and McDonough (1989). c Plot of primitive mantle-normalized multi-elements spider diagram of Abu Saqa schist, using values by Sun and McDonough (1989)

than one lithological unit, commonly named after the type locality of the lithology, which can be either a Wadi or a mountain (Jabal in Arabic). These details are found in the bulletins and technical reports issued by the Natural Resources Authority (NRA) of Jordan. A representative geochemical database of 278 samples is used to construct diagrams to

5.2.2.1 Duheila Hornblendic Suite (DHS) The DHS occurs in the form of rafts and mega-xenoliths in the granitoids of the Rahma Suite (Fig. 5.8a) and comprises diorites with minor hornblendites, gabbros and diorite (Fig. 5.2a). Furthermore, they are found intruding the ABMS paragneiss in Wadi Rahma (Fig. 5.8b). Jarrar (1998, 2002) distinguished three varieties of the DHS in the field: (a) pegmatitic diorite, where amphibole occurs as conspicuous mega prismatic crystals up to 10 cm in length; (b) biotite-hornblende gabbro-diorite with prominent foliation; and (c) massive diopside-hornblende gabbro-diorite that can be easily recognized by the predominance of equidimensional blocky amphibole rhombs and the presence of diopside. The latter variety contains as well, anhedral interstitial pink orthoclase which encloses diopside and thus forms large and fresh orthoclase poikilocrystals. Sporadic outcrops of the so-called Thawr gabbro are found close to the outcrops of the Abu Saqa schist discussed above and are lumped together with the DHS. The hornblendites are characterized by prominent and abundant megacrysts of prismatic hornblende (Fig. 5.8c), sometimes up to 10 cm long. DHS major element composition and normative mineralogy suggest calc-alkaline and tholeiitic affinities. K–Ar dating of the DHS yielded an age of 610–615 Ma (Lenz et al. 1972 in Bender 1974). Meta-gabbros and meta-diorites of the Shahmon metabasites (north of Elat), are probable equivalents of the DHS and have been dated by single-zircon evaporation method at 650–640 Ma (crystallization age, Kröner et al. 1990). Rare earth and trace element patterns (Fig. 5.8d, e) of the DHS are similar to typical island arc lavas. The Duheila Hornblendic Suite is enriched in LILE relative to the HFSE and is moderately enriched in REE [(La/Lu)n = 5–11], traits typical of arc igneous rocks. Geochemical modeling suggests derivation by 10–15% melting of amphibole-bearing spinel lherzolite, metasomatized LILE-enriched mantle wedge over a subduction zone (Jarrar 2002). Subsequent geochemical evolution of the suite was controlled by fractional crystallization of mineral phases like amphibole, plagioclase, and magnetite.

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Fig. 5.7 Geochemical classification of the rocks of Aqaba and Araba complexes. a Classification diagram proposed by Le Bas et al. (1986) and modified by Le Maitre et al. (1989). Dividing line between alkaline and subalkaline series is that of Irvine and Baragar (1971). b Classification diagram using incompatible element ratios proposed by Winchester and Floyd (1977). c Classification diagram of Peccerillo and Taylor (1976) to distinguish between tholeiitic, calc-alkaline, and shoshonitic rock series. d Classification diagram of granitoids by (Frost and Frost 2008); F* is the Fe-number = FeOtot/(FeOtot + MgO). Geochemical data are compiled from (Ghanem and Jarrar 2013; Jarrar 2001, 2002; Jarrar et al. 2003, 2004, 2008, 2013b)

5.2.2.2 Rahma Foliated Suite (RFS) The granitoids of the Rahma “foliated” Suite show local gneissose fabric produced by rafts and mega-xenoliths of foliated metasediments assimilated by these granitoids. Hence, it is the description of this suite as foliated and gneissic by Ibrahim and McCourt (1995). These include the units of Turban granite to granodiorite, Taba monzogranite, Nab’a monzogranite, and Abu Radmar granodiorite. The major outcrops of RFS are aligned along the eastern shoulder of the Wadi Araba, in the Wadis from north to south: Umm Saiyala, Abu Barqa, Turban, and Rahma (Fig. 5.2a, b). These granitoids are holocrystalline hypidiomorphic, medium- to coarse-grained, porphyritic, occasionally foliated and/or rich in metasedimentary xenoliths, and biotite

streaks as remnants of assimilated metasediments. The phenocrysts are up to locally 5 cm long white or pinkish orthoclase in both Turban and Taba units. The white orthoclase phenocrysts are riddled with concentrically arranged albite grains, sometimes forming one-third of the phenocryst. Abu Radmar granodiorite differs from Turban granitoid in being very rich in schlieren. These granitoids are intruded into the Abu Barqa Metamorphic Suite and the Duheila Hornblendic Suite; therefore, it is common to encounter rafts, meso to mega-xenoliths of both suites. Rafts of amphibolites similar to the above discussed Abu Saqa schist are also common. Brook et al. (1990) obtained a Rb-Sr isochron on the Taba monzogranite with an age of 562 ± 72 Ma and an initial ratio of 0.7037 ± 0.0005. The

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Neoproterozoic Crustal Evolution …

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Fig. 5.8 Duheila Hornblendic Suite (DHS). a Field photo showing DHS intruded by the Rahma suite granitoids. b DHS intruding the ABMS gneiss. c Hornblende megacrysts in the hornblendite of DHS. d Plot of chondrites normalized REE content of DHS. Chondrite values are after Sun and McDonough (1989). e Plot of primitive mantle-normalized multi-elements spider diagram of DHS, using values by Sun and McDonough (1989)

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U–Pb conventional and LA–ICPMS U–Pb on zircons yielded the ages 610–615, and 617.9 ± 3.7 Ma for the Turban granites, and the Taba monzogranite, respectively (Jarrar et al. 1983, 2013a; Ben Sera 2018).

5.2.2.3 Darba Tonalitic Suite The Darba Tonalitic Suite comprises the Huwwar two-mica granite, Muhtadi quartz monzodiorite and the Wa’ra granodiorite outcropping at Wadi Huwwar, Wadi Muhtadi, and Wadi Wa’ra, respectively (Fig. 5.2a, b). The Huwwar two-mica granite is medium-grained, gray-colored granite with biotite and muscovite. Furthermore, this granite has a numerous gently dipping, up to half-meter thick, microcline-muscovite quartz pegmatite dikes. Roof-pendants of the staurolite-andalusite schist of the ABMS and mega-xenoliths of the Barraq granitic gneiss are abundant in the Huwwar two-mica granite. The Muhtadi unit is medium- to coarse-grained, black to whitish gray in color, rich in rectangular plagioclase phenocrysts and numerous intermediate-mafic rounded enclaves. The Wa’ra granodiorite differs from Muhtadi unit in being more porphyritic, having lower color index, and occasionally pinkish color. Mafic enclaves (Fig. 5.9) are common in Wa’ra granodiorite in particular and in the Aqaba granitoids in general. Rb–Sr dating of the Darba Suite yielded a crystallization age of 584 ± 18 Ma and an initial ratio of 0.7038 ± 0.0001 (Brook et al. 1990). LA– ICPMS U–Pb dating of zircons constrained the crystallization age of Wa’ra and Muhtadi units at 611.8 ± 4.9, 612.5 ± 3.6 Ma, respectively (Ben Sera 2018). SIMS U–Pb ages on zircons from the two-mica granite constrained the crystallization age of this granite at 612 ± 2 Ma (Jarrar et al. 2013a) 5.2.2.4 Urf Porphyritic Suite (UPS) The Urf Porphyritic Suite includes the porphyritic granitoid units of Filk, Abyad, Rubeiq, Marsad, Huneik, and the Mulghan granodiorite. They range from granite through monzogranite to granodiorite and typically form low, undulating ridges, which are highly weathered to a white– gray-colored hills. The granodioritic units have a higher color index that reflects the high biotite and amphibole content in the Marsad monzogranite, which also has mafic enclaves. The feldspar phenocrysts are pink and/or white and are composed of perthitic orthoclase and zoned plagioclase, respectively. Quite often, the contacts between the various units are gradational. Brook et al. (1990) constructed a good Rb-Sr isochron that yielded an age of 620 ± 14 Ma (MSWD = 0.60) and 87Sr/86Sr initial ratio of 0.7033 ± 0.0001. Ben Sera (2018) obtained a laser ablation U–Pb zircon crystallization age for the Mulghan unit of 613.7 ± 4.4 Ma.

G. H. Jarrar and H. Ghanem

5.2.2.5 Rumman Suite (RS) This suite includes four units: the Hubayra diorite, Ishaar monzogranite, Qara granite, and Sabil granodiorite, with the latter as the dominant unit. These units are typically weathered, green–gray-colored, generally forming low relief hills and are heavily diked. Hornblende is the characteristic mafic mineral of the suite along with biotite. The presence of abundant hornblende and ovoidal mafic enclaves is diagnostic for the Rumman and the Darba Suites. These enclaves are common in the granodiorite but are also present in the Ishaar monzogranite and Qara granite units, which might be taken as an evidence for a cogenetic relationship. The field relationships between the Sabil granodiorite and the Ishaar monzogranite, despite being poorly exposed, suggest that the Sabil granodiorite seems to be the older phase (Ibrahim and McCourt 1995). Brook et al. (1990) obtained a well-defined Rb–Sr isochron by combining the data points from three units together due to the limited spread of the data points of the individual units. This isochron yielded a crystallization age of the Rumman Suite at 583 ± 3 Ma with and 87Sr/86Sr initial ratio of 0.7035 ± 0.0001 (MSWD = 1.8). Conventional U–Pb dating of zircons from Sabil granodiorite gave an age of *615.8 ± 1.9 Ma (Moshtaha 2011). 5.2.2.6 Yutum Suite (YS) This suite was named after Wadi al Yutum (Fig. 5.2a), where the bulk of the two constituent units, the Imran monzogranite and Abu Jedda granite, crop out. The Humrat alkali feldspar granite and syenogranite, originally included in the Yutum suite by Ibrahim and McCourt (1995), are here excluded because they are distinctly younger based on new geochronologic data. Abu Jedda granite is named after its type locality of Jabal Abu Jedda 10 km north from Aqaba at the mouth of Wadi Al Yutum. It forms high rugged mountainous terrane with tafone (cavernous) weathering forms. It has a gradational contact to its twin unit, Imran monzogranite, and intrusive contacts to the granitoids of Rumman and Urf Suites. The Abu Jedda monzo-to syenogranite is a biotite two feldspar subsolvus granitoid. It is medium- to coarse-grained granite with white-colored, in hand specimen, and zoned in thin section plagioclase and pink-colored microcline microperthite with biotite as the sole ferromagnesian mineral. The Imran monzogranite differs from Abu Jedda granite in having some hornblende as an additional ferromagnesian mineral. Muscovite, apatite, and zircon are found as accessory minerals. Jarrar et al. (1983) obtained a conventional U–Pb zircon age of 608 Ma and a biotite-WR Rb-Sr isochron age of 589 Ma. Brook et al. (1990) obtained a combined Rb–Sr WR isochron for both units of the Yutum suite which yielded an age of 585 ± 3 Ma and an initial ratio of 0.7031 ± 0.0001. Ben Sera (2018) obtained a LA– ICPMS U–Pb crystallization ages on zircon for Abu Jedda

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Fig. 5.9 Mafic enclaves in Wa’ra granodiorite from Darba suite

granite and Imran monzogranite of 604.5 ± 4.6 Ma and 608.0 ± 5.4 Ma, respectively. The REE patterns of the Aqaba Complex granitoids are generally steep and lack or have a weak Eu anomaly, which implies a garnet-controlled source of magma (Fig. 5.10a). Moreover, the spider trace elements incompatibility plot displays a strong Nb–Ta anomaly in addition to the spike at Pb and negative anomalies at Ti and P (Fig. 5.10b). The Nb–Ta anomaly again suggests a subduction-controlled tectonic environment for the voluminous calc-alkaline batholith which makes up the bulk of the NBC of Jordan. This conclusion is supported by the exclusive plot of these granitoids in the field of volcanic arc granitoids of Pearce et al. (1984; Fig. 5.10c) and the extremely low Rb/Sr ratio of these granitoids (Fig. 5.10d). The granitoids of the Araba Complex, i.e.,

Humrat-Feinan-Mubarak are shown on both diagrams (Fig. 5.10c, d) for comparison. The Rb/Sr ratios of the Aqaba and Araba Complexes granitoids are plotted versus silica (Fig. 5.10d). The diagram shows the exponentially increasing trend of the Rb/Sr ratios with increasing silica and that most of the suites have ratios less than 1. The Yutum Suite which is the youngest of the Aqaba Complex together with the Humrat-Feinan-Mubarak Suite display ratios up to 30.

5.3

Araba Complex

The Araba Complex according to McCourt and Ibrahim (1990) is defined as all rock units confined between the regional unconformity marked by the Saramuj

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Fig. 5.10 Geochemical classification of the Aqaba and Araba complexes rocks. Symbols as in Fig. 5.7. a Plot of chondrites normalized REE content of the Aqaba Complex rocks. b Plot of primitive mantle-normalized multi-elements spider diagram of the Aqaba Complex rocks. c Geotectonic classification diagram of granitoids by Pearce et al. (1984) for samples from the Aqaba and Araba complexes; the shaded area represents the position of the volcanics of the Aheimir suite; d SiO2 wt% versus Rb/Sr plot for samples from the Aqaba and Araba complexes

Conglomerate Formation (Powell 1988) and the regional unconformity marked by the start of the Ram Group (Fig. 5.11a, b). These unconformities have been named as the Araba (*605 Ma) and Ram (*530 Ma) Unconformities, respectively (Powell et al. 2014, 2015). The Araba Unconformity separates the eroded the exhumed ABMS and its intruding Turban granite from Safi Group, where it is exposed at the southern flank of Wadi Abu Barqa (Fig. 5.2b). The unconformity coincides with the transition in tectonic style and magmatism from subductioncontrolled, medium to high-K calc-alkaline granitoids to extensional tectonic regime typified by bimodal mafic shoshonitic to alkaline/peralkaline felsic magmatic activity (Jarrar et al. 1992). During this transition, the earlier subduction-controlled magmatic rocks were exhumed and

eroded producing the unconformity from this erosional event.

5.3.1 Safi Group 5.3.1.1 Saramuj Conglomerate Formation (SCF) The type locality of this conglomerate is on the SE shore of the Dead Sea. The base of the SCF is not exposed at its type locality; but it is exposed at the southern flank of Wadi Abu Barqa (Figs. 5.2b and 5.12a). It unconformably overlies the white gray Turban granite and its host the ABMS. It occurs in the form of wedge-shaped outcrop, 2 (max.)  9 km (Fig. 5.3, Jarrar et al. 1991). Three lithofacies types can be distinguished:

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Fig. 5.11 Ram Unconformity. a Representative view of the Ram Unconformity showing the clear peneplain between the Cambrian sandstone on top and Al Humrat granite from the Araba Complex, the width of the outcrop is *2 km. Al Quwayra fault zone is separating the outcrop from the middle lower elevation outcrops; another parallel fault is placing Marsad granite from Al Urf suite next to the remarkable Ram Unconformity. b Closeup of the unconformity surface showing the deeply weathered Humrat granite overlain by the Cambrian basal conglomerates

(i) Massive clast-supported coarse conglomerates; most of the clasts range in size between 5 and 30 cm (Fig. 5.12b). Nevertheless, granite boulders up to 5  3 m have been reported in Wadi Said (Fig. 5.3) at the southernmost end of the wedge-shaped outcrop. The thickness of these units can reach up to 30 m. Clasts are dominantly granitoids, followed in frequency by mafic and felsic volcanics, and least common are the gneisses (Fig. 5.12b). Clasts range from well rounded to subangular, whereby rounding increases with increasing clast size. Furthermore, granitoids are well rounded while volcanic clasts tend to be subangular. Yaseen et al. (2013) reported SIMS U–Pb zircon ages showing that the SCF comprises pebbles that range in age between 734 Ma for a granitic gneiss, through 650 Ma for a granodiorite clast, down to 624–640 Ma for andesite and rhyodacite clasts. Based on the ages of zircons from the

matrix, the maximum age of deposition of the SCF was constrained to be younger than 615 Ma. (ii) Pebble-bearing sandy conglomerates: These are mostly less than 2 m thick beds of a sandy-granule matrix that rarely exceed 4 mm with floating pebbles. (iii) Medium- to coarse-grained lithic arkosic sandstone dominated by quartz, alkali feldspars, and lithic fragments. The fine matrix of this sandstone in addition to the matrix in the previous two types recrystallized to chlorite and epidote due to the burial of the conglomerate to a depth of *5–6 km. This depth was estimated using chlorite thermometry that gave temperatures between 250 to 300 °C and assuming a relatively high geothermal gradient at the time of 50º/ km (Jarrar et al. 1993; Ghanem 2009; Ghanem and Jarrar 2013). Lithologies 2 and 3 are characterized by planar and trough crossbedding (Fig. 5.12c).

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Fig. 5.12 Representative field images of the Saramuj Conglomerate Formation (SCF). a The SCF unconformably overlies the Abu Barqa Metamorphic Suite rocks (ABMS) on the southern flank of Wadi Abu Barqa. b The Saramuj conglomerates; the green color of mafic cobbles and the matrix reflect the greenschist burial metamorphism. Clasts like those dated by Yaseen et al. (2013) are indicated. c Crossbedded lithic arkosic sandstone overlain by the pebble-bearing sandy conglomerates of the SCF. d Saramuj meta-conglomerates affected by thermal metamorphism due to the emplacement of Qunai monzogabbro from the Araba Mafic Suite. The change in mineralogy due to the contact metamorphism is indicated by the recrystallized and darker matrix

The SCF was intruded by the Qunai monzogabbro (see Sect. 5.3.2) at about 595 Ma (Jarrar et al. 1993), which developed a contact metamorphic aureole of metaconglomerates (Fig. 5.12d), granofelses, and hornfelses. Index minerals such as biotite, tremolite-actinolite, hornblende, diopside, and hypersthene appeared in order of increasing grade toward the contact (Ghanem 2009; Ghanem and Jarrar 2013).

5.3.1.2 Haiyala Volcaniclastic Formation (HVF) This formation is exposed approximately 5.5 km north-northeast of Gharandal immediately to the south of Wadi Abu Barqa on the western, lower, slopes of Jabal al-Haiyala (Figs. 5.2b and 5.13) where the dip of the succession ranges between 15° and 25° to the southeast (McCourt and Ibrahim 1990; Jarrar et al. 1991). The HVF unconformably overlies about 30 meters of the Saramuj Conglomerate Formation. The HVF was originally mapped

by Bender (1974) as “slate graywacke series.” McCourt and Ibrahim (1990) re-named the series as the Haiyala Volcaniclastic Formation since it is dominated by claystones and siltstones intercalated with felsic volcaniclastics. The main sequence of 200–220 m thick HVF is comprised of steeply dipping volcaniclastic sediments and dominated by coarse-grained tuffs and fissile, finely laminated claystones and siltstones. These are gray–green to reddish-brown (Fig. 5.13), often mottled, and contain fragments of fine-grained brick-red “felsite” in a tuffaceous matrix. Thinly bedded sedimentary rhythms, up to 0.60 m thick, are characterized by a sharp lower boundary with graded granule to silt-grade pyroclastic fragments intercalated with claystone and siltstone laminae. The latter consists of quartz, feldspar, and muscovite, with kaolinite and illite being the most common clay minerals. Laminae are well developed in the volcanic ash layers, interbedded with fine-grained siliciclastics; laminae are often separated by sharp erosional

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contacts (Jarrar et al. 1991). The pyroclastics are associated with rhyolitic lava flows and often with ignimbrite layers up to 6 m thick. They include poorly sorted, fine- to coarse-grained ash tuffs and subordinate lapilli tuffs, both rich in lithic and crystal fragments, up to 60% in abundance.

5.3.2 The Araba Mafic Suite The mafic unit of this suite comprises the Mureihil diorites (Fig. 5.2b), Ghuweir Volcanic Suite (Fig. 5.2b), the Qunai monzogabbro (Fig. 5.3), and Ayn Al Hashim gabbro (Rosetta Gabbro of Jarrar et al. 2017). Mureihil diorite intrudes the Huwwar two-mica granite of the Darba Suite and the Barraq granitic gneiss of the ABMS. The diorite is holocrystalline, medium- to coarse-grained quartz diorites. It is characterized by the mineral assemblage: zoned plagioclases with andesine rims and labradorite cores, orthopyroxene, clinopyroxene, hornblende, biotite, quartz, orthoclase, titanomagnetite and accessories of ilmenite, titanite, zircon, and apatite. Quite often hornblende is rimming pyroxene forming well-developed corona texture. This diorite shows a genetic fractional relationship with a small stock of hornblende bearing aplite granite. The age of this diorite was constrained by conventional U–Pb zircon dating at 588 + 15/−6 Ma (Jarrar et al. 1983). The discordia was recalculated by forcing the lower intercept through zero and an age of 598 Ma was obtained (Jarrar et al. 2013a). The Ghuweir Volcanic Suite was investigated by Basta et al. (1982) who described thick masses of intermediate to

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mafic extrusive rocks with associated agglomerate and tuffs, cut by a few acidic dikes. Jarrar et al. (2008) obtained a Rb– Sr isochron for these volcanics which constrained, though not well, their age at 572 ± 48 Ma. Their contact with a small outcrop of Haiyala volcaniclastic formation seems to be intrusive (Jarrar et al. 2008). Nonetheless, a better age constraint is needed. The volcanics are comprised mainly of plagioclase, augite, and pseudomorphs after olivine in the basaltic samples, and titanomagnetite; chlorite and epidote are abundant secondary phases. Plagioclase is present as phenocrysts (up to 3 mm across) and as a principal constituent of the groundmass. Microgabbros display ophitic texture where augite encloses tabular plagioclase laths. Older studies correlated the Ghuweir Volcanics with the Dokhan Formation volcanics and Hulayfah and Shammar Groups of Saudi Arabia. However, recent geochronologic data suggest they are correlative to the Jibalah Group (Johnson et al. 2013; Johnson 2014 and references therein; Powell et al. 2015). The type locality of Qunai monzogabbro (Fig. 5.14a) is Wadi Qunai (Fig. 5.3) at the southeastern corner of the Dead Sea, where it intrudes the Saramuj Conglomerate and occurs in scattered outcrops with a total area of about 1 km2. Most of the intrusion is covered by alluvium. It is gray-colored, coarse-grained, with the andesine-labradorite plagioclase (Fig. 5.14a) showing labradorescence. It has a sharp vertical contact to the host conglomerate, which is transformed to mostly black meta-conglomerate with hornfelsic matrix (Fig. 5.12d) with two-pyroxene-hornblende-biotite groundmass. The monzogabbro itself consists of andesinelabradorite, titanaugite, intermediate olivine (Fo 48%;

Fig. 5.13 Haiyala Volcaniclastic Formation (HVF) as seen from Wadi Araba main road near Wadi Abu Barqa. The formation is separated by Abu Barqa Metamorphic Suite rocks (ABMS) and Turban granite from Rahma Suite by the Araba Unconformity. The unconformity surface is obscured by downfaulting of the HVF relative to the ABMS/Rahma granitoids. To the right, the HVF is overlain by the Aheimir volcanic rocks

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Gabbro by Jarrar et al. 2017); and phyric microgabbro. The outcrops are small and are not mappable at the scale of the map in Fig. 5.2a; however, a red star on the map indicates the approximate location of the small outcrop of the Rosetta Gabbro. The Rosetta Gabbro (Fig. 5.14c) forms a small (*750 m2) stock intruding the JMS and is in turn intruded by Abu Jedda granite, from which Jarrar et al. (2017) reported a crystallization age of 596 ± 5 Ma from five data points with low common Pb and discordance 1000 km wide, has also been termed the Transgondwanan Supermountains (Squire et al. 2006). The ANS represents the northern sector of the EAO. (IV) 600–570 Ma Continued Shortening, Escape Tectonics, Post-Amalgamation Basin Formation, and Orogenic Collapse: Collisional orogenesis involved considerable compression and shortening that continued during the first *50 million years of the Ediacaran Period (Veevers 2003). Deformation included strike-slip shear zones and tectonic collapse structures in the northern EAO (Egypt, Sudan, and northern Arabia), formation of N-trending upright tight folds and shear zones in the central EAO (Ethiopia, Eritrea, and southern Arabia), and formation and uplift of high-grade gneisses and granulites in the southern EAO (Abdelsalam and Stern 1996). The most intense collision occurred in the southern EAO, which had the thickest crust, highest mountains, and the deepest erosion. Accordingly, metamorphic grade of exposed ANS rocks decreases northward, compared to higher grade granulite facies further south (Stern 1994). By *570 Ma, ANS assembly had accreted to the Saharan Metacraton and evolution continued within the southern Paleotethys realm of northern Gondwana (Abdelsalam et al. 2003; Johnson et al. 2013). Greater Gondwana began to break up almost as soon as it formed at the end of Neoproterozoic time, shedding microcontinents into especially Asia all through Paleozoic and early Mesozoic time, with the core of Gondwana finally rupturing in Late Jurassic time. (V) 600–530 Ma Formation of the Afro-Arabian Peneplain: Extensive erosion of the EAO resulted in cutting of a widespread regional unconformity prior to Cambro-Ordovician time (Avigad et al. 2005), herein termed the Afro-Arabian Peneplain (AAP) (see Powell et al. 2015). This peneplain is recognized throughout northern Gondwana (Fig. 7.5D), extending from Morocco to eastern Arabia and Oman (Stern 1994; Meert and Van Der Voo 1997; Garfunkel 1999; Avigad et al. 2005; Squire et al. 2006; Avigad and Gvirtzman 2009; Al-Husseini 2014). Although the timing of the initial downcutting of the AAP is unknown, erosion would have begun as soon as regional uplift began. Some regional uplift could have initiated during the Sturtian glaciation, but more

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regional erosion of the ANS is likely to have spanned Marinoan and Ediacaran glaciations. Relief would have been enhanced by glacial eustasy, with isostatically adjusted sea level falls on the order of *500 m or more estimated for the Sturtian and Marinoan panglacial episodes (Hoffman et al. 2007; Liu and Peltier 2013). Glaciation of the developing EAO would have accelerated erosion, and we speculate here that it could have contributed to initial peneplain formation. As a singular erosion surface, the AAP principally corresponds with deeply eroded highlands of the EAO, but along the northern periphery of the EAO at least two separate regional unconformities are recognized: (1) the *585 Ma Sub-Jibalah Unconformity in N. Arabia (with equivalents in Iran, Jordan, and Oman) and (2) the a *530–520 Ma unconformity correlated in Jordan (known as the Ram Unconformity), Saudi Arabia (Sub-Siq Unconformity), Iran (Sub-Lalun Unconformity and Oman (Angudan Unconformity) (Al-Husseini 2011). We use the term AAP inclusively to represent the significant erosional episodes that affected the EAO prior to Cambro-Ordovician time and consider it equivalent to the globally recognized Great Unconformity. See Sect. 6.3.3 for further discussion. Neoproterozoic paleogeographic reconstructions (720– 550 Ma, Fig. 7.5) generally place the ANS (as inferred oceanic arcs) within the greater Mozambique Ocean between India and Congo cratons. For example, Collins and Pisarevsky (2005) place ophiolite-bearing strata of the Adola Belt (southern Ethiopia) outboard of the Congo Craton at 750 Ma. As the Adola Belt lies in the southern EAO, it is reasonable to infer a similar outboard location for the ANS. The earliest reliable regional palaeomagnetic data for the ANS derive from the time of Gondwana amalgamation (Fig. 7.5C– D); late Cryogenian (593 ± 15 Ma) Dokhan volcanics of Egypt (Davies et al. 1980; Wilde and Youssef 2000), interpreted to have a subtropical paleolatitude (20.6 ± 5.08°, Trindade and Macouin 2007; but see Nairn et al. 1987). Low subtropical paleolatitudes (9–13°) are also reported for Huqf Supergroup units in Oman between 722 and 544 Ma (Kempf et al. 2000; Kilner et al. 2005; Allen 2007), with Oman likely amalgamating with the ANS between 645 and 544 Ma (Rieu et al. 2007). To the extent that the ANS occupied a latitudinal range similar to the Congo- Sào Francisco and/or East-Sahara cratons during the early Cryogenian (c. 750 Ma), as implied by various paleogeographic reconstructions (e.g. Trindade and Macouin 2007; Li et al. 2013; Pisarevsky et al. 2008), and Oman was latitudinally adjacent, possibly along the margin of the Indian craton (Denèle et al. 2012; Johnson 2014; Whitehouse et al. 2016; Alessio et al. 2017), the ANS may have occupied low to intermediate paleolatitudes during much of

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Fig. 7.5 Global paleogeographic reconstructions showing the inferred position of the ANS during Neoproterozoic glacial intervals. A– C. Paleogeographic reconstructions (Li et al. 2013) indicate that the ANS (purple shaded ellipses) generally occupied tropical to subtropical latitudes during the Sturtian, Marinoan, and Ediacaran (Gaskiers) glaciations; see original reference for numbered localities. Rifting and break-up of Rodinia (c. 900–750 Ma) were associated with seafloor spreading, arc and back-arc basin formation, and terrane accretion in the Mozambique Ocean—the birthplace of juvenile Neoproterozoic crustal terranes that characterize the ANS. Accommodation space in most of the ANS likely inverted by 630 Ma in response to the closure of the Mozambique Ocean between converging elements of West (Sahara, Congo) and East Gondwana (Indian) and emergence of the EAO (aka Transgondwanan Supermountain, Squire et al. 2006, dashed ellipse in B). Craton/terrane name abbreviations in A–C: A Amazonia; Ae—Avalonia (east); Aw—Avalonia (west); B—Baltica; C—Congo; CAFB—Central Asian Fold Belt; EA—East Antarctica; ES—East Svalbard; G—Greenland; I—India; K—Kalahari; L—Laurentia; NA—Northern Australia; NC—North China; R— Rio Plata; S—Sahara; SA—Southern Australia; SC—South China; Sf—Sao Francisco; Si—Siberia; T—Tarim; WA—West Africa. D Paleogeographic and structural setting of the ANS at c. 550 Ma within the context of Greater Gondwana amalgamation and continued structural shortening of the EAO (modified from Meert and Lieberman 2008; Gray et al. 2008). Sedimentary records of extensional post-amalgamation basins, possibly formed in association with escape tectonics (

) along the Gondwanan margin (e.g., Jibalah Group basins), may preserve indications of Ediacaran

glaciation. The modern southern limit of Early Palaeozoic sandstone (dashed bold line) documents the continent-scale extent of the Afro-Arabian Peneplain (AAP) that developed following uplift and erosion of the EAO (Avigad et al. 2005)

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the Neoproterozoic (Fig. 7.5A–C; Miller et al. 2011). The closest modern analog to the ANS may be the Indo-Australian Archipelago, the islands and shallow submerged regions of which constitute portions of volcanic island arcs and volcanic arcs within a tropical climate.

7.4

Expected Manifestations of Neoproterozoic Glaciations in the ANS and EAO

Age constraints for Neoproterozoic glaciations have implications for the accumulation and preservation of Snowball Earth episodes within the ANS and EAO (Fig. 7.4). The long-duration Sturtian and Marinoan panglacial episodes are particularly likely to have impacted depositional systems within the evolving ANS and EAO. Because of the lack of direct paleolatitude evidence constraining a low latitude for the pre- *600 Ma ANS (Davies et al. 1980; Wilde and Youssef 2000), any Tonian (pre-Sturtian) glacial activity in the ANS could have been regional at higher latitudes similar to Phanerozoic icehouse episodes. As mentioned above, evidence of Tonian panglacial intervals is controversial (Rooney et al. 2015). Nonetheless, the possibility of regional glacial activity preceding the Sturtian panglacial episode cannot yet be entirely refuted. Tonian and Sturtian episodes (e.g., 800 * 660 Ma) would have occurred while ANS juvenile crust was being generated by magmatism associated with seafloor spreading, volcanic arcs, back-arc basins, and oceanic plateaus within and around the Mozambique Ocean (Fig. 7.4II). Such glacial episodes could also have accompanied or followed arc terrane accretion (*870–690 Ma), possibly including early phases of contractional tectonics associated with Mozambique Ocean closure prior to *630 Ma Gondwana fusion (Avigad et al. 2007). The recurrence of Neoproterozoic BIF, now considered to have been largely a Sturtian phenomenon (Cox et al. 2013), would have overlapped with this phase of ANS development. The physiography of continental shelves, oceanic plateaux, arcs, and accreted arc terranes within the Mozambique Ocean provided numerous basins with accommodation space capable of holding sediments that reflected glaciation. A wide range of water depths similar to those associated with modern back-arc basins and intraoceanic forearcs is probable, from shallow depths of a few hundred meters within the photic realm of continental shelves, oceanic plateaus, and island arcs, to abyssal depths (2500–5000 m). Sedimentation may also have involved marine carbonate deposition away from significant siliciclastic/volcaniclastic input—perhaps outboard of or along margins of magmatically dormant accreted arc terranes. We note that Neoproterozoic carbonate successions are particularly rare in the northern ANS but poorly studied candidates exist in the Keraf and Nakasib

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sutures of Sudan (Abdelsalam and Stern 1993; Abdelsalam et al. 1998) and in N. Ethiopia. Because of the value of marine C and Sr chemostratigraphy for Neoproterozoic correlation, low-grade marine carbonate successions in the ANS are of particular importance. Any sediments deposited in association with pre-Sturtian or Sturtian glacial episodes would have been deformed during later accretion and collision events (Fig. 7.4III–IV). Because of the southward increase in metamorphic grade within the ANS, optimal preservation of any glacigenic successions is most likely to occur within low grade greenschist the successions of middle (e.g., N. Ethiopia) and northern portions of the ANS. Preservation potential is lowest for Marinoan and Ediacaran glacial episodes because these occurred when the ANS began to rise out of the sea in response to collisions between various terrane fragments within the Mozambique Ocean, culminating in terminal collision between E. and W. Gondwana with full development of the EAO (Fig. 7.4III). Most of the ANS was likely above sea level by *630 Ma. As continental collision got underway, basins capable of preserving syn-glacial sediments moved away from the EAO to its flanks. By the end of the Neoproterozoic, the EAO may have been located at intermediate (*30–60 °S) latitudes (e.g., Dalziel 1997; Li et al. 2008) capable of hosting a thick continental ice sheet. In addition, the EAO was the site of the Transgondwanan Supermountains, which are thought to have been Himalayan in relief (Squire et al. 2006) and these were prone to glaciation. Continental glaciers are powerful agents of erosion, thus Marinoan and/or Ediacaran ice sheets could have contributed to erosion and downcutting of EAO hinterlands, as initial phases of AAP beveling (Fig. 7.4III– V). Associated glacial deposits could be preserved in deep grabens around the margins of the ANS, such as those preserving glacigenic successions of the Huqf Supergroup in Oman (Stewart 2016), or possibly also in continental shelf deposits on the northern flank of the end-Neoproterozoic supercontinent, such as may exist beneath Israel and N. Iran (Etemad-Saeed et al. 2016). Reworking of Marinoan and Ediacaran glacial deposits by alluvial processes may obfuscate definitive field evidence of primary glacial associations. Evidence that clasts were transported over great distances from terranes outside of the ANS may support an earlier glacial association, for example in the case of the Atud/Nuwaybah diamictite (Ali et al. 2010a). Minimum ages from detrital zircons recovered from clasts and matrix constrain when these glacial deposits were reworked.

7.5

Evidence for Glaciation in the ANS

Prospective Neoproterozoic glacigenic deposits in the ANS are reviewed below. These are presented in terms of four glacial episodes: Tonian (*780–755 Ma), Sturtian (720–

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660 Ma), Marinoan (*645–635 Ma), and Ediacaran (*585–550 Ma), of which the Sturtian and Marinoan episodes define most extreme global climate fluctuations of the Cryogenian Period. Outside the ANS, the closest Neoproterozoic glacial localities are known in NW Africa (Shields-Zhou et al. 2011), central Africa (Master and Wendorff 2011; Tait et al. 2011), and Oman (Allen et al. 2011a, b). Glacial deposits in NW and central Africa were hosted by different cratonic blocks outside of the Mozambique Ocean. The relationship of Oman to the ANS (and Mozambique Ocean) is poorly understood but Oman is thought to be related to buried crust in eastern Arabia (e.g., Johnson and Stewart 1995). Although Oman’s crystalline basement contains juvenile Neoproterozoic rocks similar to the ANS, contrasts in the timing of magmatism and sedimentation suggest Oman had a different geologic setting (Johnson et al. 2011) prior to Gondwana amalgamation during the latest Cryogenian and Ediacaran Periods (Rieu et al. 2007). Gondwana consolidation shortened the W-E distance (present coordinates) between the ANS northern sector of the EAO and some of these terranes within the northern peri-Gondwana margin. For brevity in the case of occurrences reported in Stern et al. (2006) for which there is no new information, this is simply stated.

7.5.1 Evidence for Tonian (c. 780–755 Ma) Glaciation 7.5.1.1 ~780 Ma Diamictite in Eastern Sudan (Meritri Group) and Central Arabian Shield (Mahd Group) Indications of possible Tonian glacial sedimentation in the ANS include *780 Ma diamictite deposits on the southern margin of the Nakasib-Bi’r Umq Suture Zone in the Meritri group of E. Sudan (Fig. 7.2, locality 1) and Mahd Group of the central Arabian Shield (Fig. 7.2, locality 2) (Stern et al.

N. R. Miller and R. J. Stern

2006, 2011). These were deposited in a passive margin setting *30 million years prior to *750 Ma collision between terranes (Hijaz-Gebeit, Jiddah-Haya) forming the suture zone. Both groups are dominated by volcanic and volcaniclastic rocks but contain basal polymict conglomerate units. Exposures of the *2 km thick Meritri group are located within the suture zone, and both the internal stratigraphy of the Meritri group and its contact relationships with bracketing units are unclear due to folding, faulting, and shearing. Oldest Meritri group deposits consist of polymict conglomerate bearing deformed clasts (  70 cm length) of granite, granodiorite, diorite, rhyolite, ignimbrite, and carbonate (Abdelsalam and Stern 1993). Mahd group outcrops occur further south of the suture zone, so stratigraphic relationships are less affected by structural complexities (Johnson et al. 2003; Stern et al. 2011). Oldest Mahd group deposits are matrix-supported diamictite (1– 5 m in thickness), with abundant angular to sub-angular clasts (  30 cm length) of granitic and felsic volcanic rocks (Fig. 7.6A–B) (Stern et al. 2011). Ages bracketing the Meritri group (790 ± 2 Ma to 779 ± 3 Ma; Stern and Absdelsalam 1998) and Mahd group (785 ± 6 Ma to 777 ± 5 Ma; Hargrove et al. 2006b) suggest that diamictite deposition was largely contemporaneous over a *10 myr interval. The Mahd group unconformably overlies diorite and tonalite of the *811 Ma Dhukhr batholith (Stoeser and Stacey 1988) indicating a substantial erosional interval at *810–780 Ma, which could be associated with glaciation (Johnson et al. 2003). The *790–780 depositional window for basal diamictite in the Meritri and Mahd groups coincides with deposition of the lower Tambien Group in the southern ANS (Sect. 6.2.1); for example, Tsedia Slate deposition was ongoing at 778.72 ± 0.24 Ma in Tsedia synclinorium and followed the Bitter Springs CIE (Swanson-Hysell et al. 2015). Additional mapping and study Meritri and Mahd group deposits are required to understand the origin of the unconformity at the base of the Mahd group

Fig. 7.6 Outcrop photographs of possible Tonian glacial diamictite at the base of the Mahd Group (from Stern et al. 2011). A Nonconformity of *770 Ma Mahd Group basal diamictite resting on truncated 806 Ma Dhukhar batholith, Saudi Arabia (Fig. 7.2, locality 2). B Mahd Group basal diamictite with dark matrix-supported angular granitic clasts

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and determine if they preserve evidence of glacial deposition. C and Sr isotope analysis of least altered carbonate units, if present, could aid in regional correlation.

of carbonate-rich facies in the S. ANS (Tambien Group), as next discussed.

7.5.1.2 ~750 Ma Ghamr and Amudan Volcanosedimentary Basins, Central Arabian Shield In the central Arabian Shield, Ghamr and Amudan basins (Fig. 7.2, localities 3–4) contain 400–600 m thick, weakly metamorphosed, volcano sedimentary successions that are discontinuously exposed along and southward of Bi’r Umq suture zone in west-central Arabian Shield. Named for the Ghamr group (Kemp et al. 1982) and Amudan formation (Ramsay 1986), these successions unconformably overlie truncated 816–775 Ma arc-related basement rocks (Calvez and Kemp 1982; Hargrove 2006a) and contain polymict conglomerate and matrix-supported pebble conglomerate units, with clasts sourced from underlying arc volcanic and plutonic basement (Johnson et al. 2013). Deposition may have occurred near volcanic centers in terrestrial (alluvial fan, drainage channels) and near-shore marine environments near a convergent margin, possibly in a retroforeland basin setting (Johnson et al. 2013). The basal contact for the Ghamr group is an angular unconformity where it overlies eroded arc rocks, and a nonconformity where it overlies plutonic rocks. Ages constraining Ghamr group deposition derive from a Rb–Sr whole-rock isochron age of 748 ± 22 Ma for subvolcanic rhyolite (Calvez and Kemp 1982) and U–Pb zircon SHRIMP condordia ages of 753 ± 6 Ma, 752 ± 4 Ma, and 746 ± 6 for dacite and andesite (n = 2) lavas, respectively (Hargrove 2006a). The Amudan basin succession is structurally conformable above arc rock basement but dating of the latter demonstrates a 15 myr erosional hiatus (Hargrove 2006a). Collectively *745–755 Ma Ghamr and Amudan deposition followed an extended erosional hiatus over older arc terranes. This time interval would have partially overlapped and followed *780–750 Ma collision and amalgamation of the Bi’r Umq suture zone (Johnson et al. 2013). The lengthy hiatus preceding Ghamr and Amudan deposition, along with basal diamictite units, could be associated with glacial erosion. No glacigenic deposits have been reported, but such evidence has not been sought. The association of Tonian arc-terrane collisional settings with intense erosional episodes (similarly applicable to the older Meritri and Mahd group deposits) could bear on greenhouse gas concentrations in the lead up to the Cryogenian, as juvenile crustal terranes located at low latitudes would have been particularly susceptible to chemical weathering and drawdown of pCO2 through the burial of carbon in carbonates and organic matter. The *775–750 Ma erosional interval preceding Ghamr/Admudan deposition corresponded with deposition

7.5.2 Evidence for Sturtian Glaciation (~717– 659 Ma) 7.5.2.1 Ethiopia/Eritrea—Tambien Group Evidence of Sturtian glaciation in the ANS occurs in the Tambien Group of northern Ethiopia (Tigrai province; Fig. 7.2, locality 5). The Tambien Group is a mixed siliciclastic and carbonate succession of Tonian and early Cryogenian age, preserved mainly in N–NE trending synclinoria and related thrust belts. The most-studied localities include the Shiraro region, Mai Kenetal, Tsedia, and Chemit synclinoria in western Tigrai, and Negash and Samre synclinoria in the eastern Tigrai (Fig. 7.7A–D). Prospectively equivalent upper Tambien Group strata of the Gulgula Group occur in western Eritrea (west-directed arrow above Fig. 7.7A) just west of the Shiraro area, as well as in the Adobha Abi terrane and Bizen Domain to the north in Eritrea (north-directed arrows above northern extent of Fig. 7.7). Carbonate transitioning upward into polymict diamictite attributed to the Sturtian glaciation is documented at the top of the Tambien Group in eastern Negash (Miller et al. 2003, 2009; Alene et al. 2006; Avigad et al. 2007; Swanson-Hysell et al. 2015) and Samre synclinoria (MacLennan et al. 2018; Park et al. 2019). No post-Sturtian cap carbonate succession has been discovered, making Tambien Group diamictites Ethiopia’s youngest known Neoproterozoic metasediments. Tambien Group exposures are unconformably overlain by Ordovician Enticho Sandstone, and this unconformity corresponds with the AAP (Fig. 7.4E). Tambien Group deposition followed magmatism and volcaniclastic deposition of the Tsaliet Group (*850– 755 Ma), corresponding with volcanic arc activity in the Tigrai sector of the Mozambique Ocean (Fig. 7.4II). Terminal collision between western and eastern fragments of Gondwana to form the EAO deformed the Tsaliet-Tambien Group supracrustal assemblage, producing tight upright folds (Fig. 7.4III–IV) that host the studied inliers. This collisional time interval overlapped with the long-duration Sturtian glaciation. Subsequent crustal thickening and melting of lower crustal rocks induced a phase of late- to post-orogenic igneous activity (so-called “Mereb” granitoids) that peaked around 620 Ma (Teklay et al. 2001; Bentor 1985; Stein 2003). Undeformed “Mereb” granitoids penetrate the entire Tsaliet-Tambien Group supracrustal sequence. The age of Tambien Group deposition is crudely constrained between deformed syn-tectonic granitoids and metavolcanics of the Tsaliet Group (*850–755 Ma) and the post-tectonic emplacement of undeformed “Mareb”

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b Fig. 7.7 Location of key Tambien Group exposures (unshaded units) within northern Ethiopia (Tigrai Province): A Shiraro area; B Mai Kenetal

Synclinorium; C Negash Synclinorium; D Samre area). Metavolcano–sedimentary block boundaries (Tadesse et al. 2000) occur only within the Tsaliet Group and are inferred to represent accreted arc terranes or slivers in a supra-subduction zone setting. Stars show geochronological localities for syn-tectonic intrusives (white), post-orogenic intrusives (black), and detrital zircons (white stars with black dots) discussed in the text. Prospectively equivalent upper Tambien Group strata of the Gulgula Group occur in western Eritrea (arrow) just west of the Shiraro area, as well as in the Adobha Abi terrane and Bizen Domain (arrows) to the north in Eritrea. M. Alem (Madahne Alem) marks the western limb locality in the Negash Synclinorium bearing 774.7 ± 4.8 Ma zircons in slate *16 m below lowest Tambien Group carbonate beds (Avigad et al. 2007). Negash Synclinorium is structurally bounded to the east by the Atsbi Horst (AH). Inset map (below) shows the location of an unnamed pebbly mudstone unit in northeastern Eritrea (Cecioni 1981) that could correlate to the Tambien Group. Figure modified from Miller et al. (2011)

granitoids (*660–580 Ma) (see Avigad et al. 2007 and Miller et al. 2011 for specific ages). Previous work—The determination of Sturtian age glacial deposits follows from the first investigations of prospective NSE glacigenic sediments within the ANS, in metasediments of the Bizen Domain (Fig. 7.2, locality 6) and Tambien Group in E. Eritrea and N. Ethiopia by Beyth et al. (2003) and Miller et al. (2003) (Fig. 7.7). Early mapping in this region by Verri (1909), Bibolini (1920, 1921, 1922), Cecioni, 1940-41 (in Cecioni 1981) and Beyth (1972), well before the concept of Snowball Earth had emerged (Kirschvink 1992), identified putative Proterozoic glacigenic sediments (variously termed: pebbly mudstones, conglomerate, arkosic sandstone, pebbly slate) in NE Eritrea, the Shiraro region, and core of Negash syncline in N. Ethiopia. From a small sampling of widespread carbonate units, some with negative d13C values, Beyth et al. (2003) suggested that the Tambien Group might preserve evidence of NSE episodes. From reconnaissance sampling, Miller et al. (2003) established rudimentary C and Sr isotope stratigraphies for the Tambien Group in Negash Syncline and concluded that the diamictite was likely glacial and attributable to the Sturtian glaciation. Prior to this study, the closest known Sturtian glacial locality was in the Huqf Supergroup of SE Oman, where glacigenic diamictite occurs in the basal Ghubrah Member of the Ghadir Manqil Formation (Tschopp 1967; Glennie et al. 1974). Brasier et al. (2000) determined a U–Pb zircon age of 723 + 16/−10 Ma for zircon recovered from tuffaceous wackes interbedded within the Ghubrah diamictite, providing one of the earliest age constraints for the onset of Sturtian glaciation. Resampling of this bed subsequently yielded zircons producing a more precise CA– TIMS U–Pb age (Bowring et al. 2007; 206Pb/238U date of 711.52 ± 0.20 Ma and 207Pb/206Pb date of 714.2 ± 0.6 Ma), indicating that glaciation began prior to *713 Ma in Oman. More systematic studies of the Tambien Group over the past 15 years have since confirmed its association with the Sturtian panglacial episode, extended understanding of its regional litho- and chemostratigraphy and depositional age, and now establish the Tambien Group as an exceptional archive of environmental change during the Tonian transition to Cryogenian Earth systems (Fig. 7.8). In Mai Kenetal

synclinorium, Alene et al. (2006) identified a negative d13C excursion in the lower carbonate unit (Assem limestone) of the Tambien Group, with likely equivalents in Tsedia and Chemit synclinoria to the east, and suggested it correlated with the non-glacial *800 Ma Bitter Springs negative CIE. Avigad et al. (2007) published SHRIMP U–Pb ages for magmatic and detrital zircons in units below, within, and above the Tambien Group that further constrained its depositional age. Miller et al. (2009, 2011) documented more detailed C and Sr isotope stratigraphies for Tambien Group exposures in the Shiraro Area, Mai Kenetal and Negash synclinoria, and the Samre area, proposed a regional chemostratigraphic correlation scheme, and suggested from field reconnaissance and satellite imagery that the Samre region might also host Sturtian diamictite above black limestone (Matheos Fm) similar to Negash. Subsequent investigations of the Samre region verified the occurrence of stratigraphic units equivalent to the Negash facies, including glacial diamictite (Bussert 2010; MacLennan et al. 2018; Park et al. 2019). Recent studies have also enhanced C- and Sr-isotope stratigraphies and added important absolute age constraints for deposition of Tambien Group units, including the diamictite (Swanson-Hysell et al. 2015; MacLennan et al. 2018; Park et al. 2019). Characteristics, age, and origin of glacigenic deposits— The transition from Matheos Formation black limestone to diamictite appears to be conformable in Negash and Samre inliers, involving increasing proportions of non-calcareous phyllitic slate that pass upward into pebbly slate without interbedded carbonate (Fig. 7.8). At Negash, the diamictite interval is *200 m thick above its transitional base. Diamictite clasts, initially sub-centimeter in scale, increase in abundance and size (up to 20 cm in diameter) upward (Fig. 7.8C–E). Clasts are matrix-supported and the finer matrix retains horizontal layering (1–3 cm), with lateral pinching and swelling. Syn-orogenic compression deformed the diamictite to varying extents, including elongated clasts with pressure shadows and foliation of pelitic intervals. Despite this, clasts that appear to preferentially deform underlying matrix laminae (possible dropstone textures) are common (Fig. 7.8F). Larger diamictite clasts have subrounded, elongate, and angular shapes (including some bullet-nosed clasts; Fig. 7.8E) and striated clast surfaces

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Fig. 7.8 Sturtian glacial diamictite of the Tambien Group exposed in the cores of Negash (A) and Samre (B) synclinoria. C–F Diamictite photos from Negash Synclinorium: C Steeply bedded diamictite with matrix-supported clasts (N. Miller for scale). D. Polymictic matrix-supported pebbles. E Outsized rounded cobble in fine-grained weakly-foliated pelitic matrix. F Probable dropstone (well-rounded volcaniclastic cobble)

have been reported (Miller et al. 2003; MacLennan et al. 2018). Clast lithologies are polymictic, including felsic volcanic rocks, granite, fine-grained limestone retaining primary depositional textures, and dolomite, low-grade

semipelitic sediments, and rare volcanic conglomerate consistent with the upper Tsaliet Group (Miller et al. 2011). Carbonate clasts in the diamictite show a wide range of d13C values consistent with derivation from older Tambien Group

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carbonate units, including negative d13C values similar to negative CIE intervals identified in western (e.g., Mai Kenetal, Tsedia, Chemit; Miller et al. 2006; Alene et al. 2006) and eastern (Negash, Samre; Swanson-Hysell et al. 2015; MacLennan et al. 2018; Park et al. 2019) inliers. SHRIMP U–Pb detrital zircon analysis of the Ordovician Enticho Sandstone, interpreted to reflect the proximal crustal composition of the southern ANS in Ordovician time, demonstrates two magmatic modes, an older mode at *800 Ma (major concentration between 820 and 760 Ma) consistent with contributions from Tsaliet-like crust generated in island arc settings, and a younger mode at 620 Ma (major concentration between 660 and 580 Ma), consistent with contributions from late- to post-tectonic “Mereb” granitoids formed in association with Gondwana collision and crustal thickening (Avigad et al. 2007). The *688 Ma midpoint of the Sturtian glacial interval (*717–659 Ma) is conspicuously centered within this *100 myr interval of reduced igneous activity. Subsidence and deposition of shallow marine carbonates and mudrocks of the Tambien Group therefore likely occurred in the magmatic lull of waning arc volcanism. Detrital zircon analysis of the Negash diamictite (and prospectively equivalent arkosic sandstone in the Shiraro region) found no zircons younger than *740 Ma, suggesting glacial deposition occurred between *740 Ma and the onset of “Mereb” magmatism (660–580 Ma; Avigad et al. 2007). In the Samre area, the culminating diamictite has recently been constrained to be younger than 719.7 ± 0.5 Ma, based on nearly identical U–Pb ID–TIMS dates from tuffaceous siltstones 74 and 84 m lower in the underlying transition (Marian Bohkakho Fm) from Matheos Formation black limestones (MacLennan et al. 2018). These ages, in addition to 87Sr/86Sr values of 0.7066 in uppermost carbonate intervals (Miller et al. 2003, 2009), strongly support that the Negash and Samre diamictite intervals are products of the *717–659 Ma Sturtian panglacial episode and potentially correlative with the Atud/Nuwaybah diamictites of Egypt and Saudi Arabia. The diamictites appear to have been derived from a proximal source, as neither clast inventory nor the detrital zircon spectra indicate significant contributions from preNeoproterozoic sources outside the ANS (Avigad et al. 2007). Clast lithologies consistent with derivation from underlying Tambien group carbonate and Tsaliet volcanics could mean the entire Tambien Group (*1500 m) was locally uplifted or tectonically inverted to generate the diamictite source terrane. Dramatic eustatic sea-level lowering on the order of >500 m related to expansion of the Sturtian cryosphere (Hoffman et al. 2007; Liu and Peltier 2013) would have enhanced erosion of shallow marine terranes. Regional lithostratigraphy and timing of Tambien Group deposition—There remain many unanswered questions about

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the timing of Tambien Group deposition and its regional lithostratigraphy prior to diamictite deposition. The most complete successions occur in Mai Kenetal and Negash synclinoria, but they have differing lithostratigraphies and it has proven difficult thus far to establish radiometric ages to definitely establish how they may correlate. The Mai Kenetal sequence involves successive alternations between slate and dark limestone, the lower limestone unit (Assem Limestone) of which is associated with a sharp basal contact and negative d13C compositions. The Negash succession begins with slate and increasing proportions of dolomite (Didikama Formation), below a sharp contact with overlying dark limestone (Matheos Formation) that in turn grades upward into glacial diamictite. Some regional correlation challenges follow from different lithostratigraphic frameworks used in compiling four map sheets (Mekele: Arkin et al. 1971; Adi Arkay: Hailu 1975; Adigrat: Garland 1980; and Axum: Tadesse 1999) that cover Tambien Group exposures in Tigrai. Regional facies differences are possible, if not likely, because Tambien Group deposition may have occurred while arc terranes were still amalgamating and/or during Mozambique ocean closure and relief differentiation from early compressional shortening. There are two broad schools of interpretation for regional correlation, related to how depositional units in Mai Kenetal synclinorium correlate to those exposed in Negash and Samre synclinoria. Several workers have established chemostratigraphic arguments, consistent with the suggestion of Beyth (1972) that the Mai Kenetal and Negash successions, both of which begin above Tsaliet Group metavolcanics, must somehow correlate laterally (Miller et al. 2003, 2009, 2011; Alene et al. 2006), although a significant unconformity may occur in the Negash succession below its contact with upper dark limestone unit (Garland 1980). The other interpretation is gleaned from the Adi Arkay map, which in one key area interprets the entire Mai Kenetal succession to underlie the Negash succession. The field relationships underlying this interpretation should be confirmed, as these map compilations were substantially made from aerial photography with ground-truthing in selective field traverses. 87 Sr/86Sr chronostratigraphy could help rectify lithostratigraphic relationships if all Tambien Group carbonate units were deposited in open marine basins of the Mozambique Ocean. The Tonian seawater evolution curve is still in an early stage of refinement (Fig. 7.1B), but the dominant trend over 850–700 Ma indicates increasing 87Sr/86Sr values from *0.7055 to 0.7068 by *770 Ma, followed by somewhat lower values 0.7067 before further decreasing near the onset of the Sturtian glaciation (Zhou et al. 2020). Sr isotope data in Miller et al. (2009) show increasingly radiogenic compositions up-section in the most-studied Negash and Mai Kenetal inliers. In Negash, values increase within the Didikama Fm, from *0.7055 ± near

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the base to *0.7062 near its top, before abruptly jumping to 0.7066 in overlying Matheos Fm black limestones. This jump, in addition to the abrupt lithologic change (dolomite to black limestone) may indicate an unconformity in the Negash sequence at this level (as previously indicated by Garland 1980). In the higher Matheos transition to diamictite, 87Sr/86Sr values decrease to values approaching 0.706. In Mai Kenetal, a nearly comparable range of 87Sr/86Sr values are documented. Values in the Assem Limestone (*0.7062) and Tsedia Slate (0.7063) are less radiogenic than the culminating Mai Kenetal Limestone (0.7067). If the Negash facies are younger than the Mai Kenetal facies, and the measured 87Sr/86Sr values are close to depositional values, then the Tonian seawater 87Sr/86Sr evolution curve requires a strong decrease of 0.0012 following Mai Kenetal Limestone deposition to reach values obtained in the lower Didikama Formation. In addition to similar 87Sr/86Sr compositions, the Mai Kenetal LS, and Matheos Formation limestones have comparable d13C values and uniquely enriched Sr concentrations. Although 87Sr/86Sr values could be differentially affected by dolomitization/alteration, they could indicate that Negash Didikama strata are older than the Mai Kenetal Limestone. The addition of absolute age constraints for Tambien Group units in key inliers, particularly from ash falls that can provide depositional ages, is critically needed and has recently begun using U–Pb ID–TIMS dates in recent publications. In Mai Kenetal synclinorium, deposition of Tsaliet Group volcanics was ongoing at 822.2 ± 1.3 Ma (Swanson-Hysell et al. 2015; Park et al. 2019). In eastwardly adjacent Tsedia syncline, tuff samples correlated within the Werri Slate and Tsedia Slate respectively indicate ongoing deposition at 815.29 ± 0.32 Ma and 778.72 ± 0.24 Ma (Swanson-Hysell et al. 2015). The lower limestone unit (Assem Limestone) at Mai Kenetal and its negative CIE is inferred to have been deposited between these ages (Swanson-Hysell et al. 2015). Consistent with this interpretation are U–Pb zircon and Re–Os dates from the Fifteenmile Group of northwest Canada, which constrain a maximum age for the onset of the Bitter Springs CIE of 811.51 ± 0.25 (Macdonald et al. 2010a) and 810.7 ± 6.3 Ma (Cohen et al. 2017), respectively. In Negash synclinorium, Tsaliet Group metavolcanic deposition was ongoing at 794.3 ± 0.6 Ma (Park et al. 2019) and detrital zircons in the overlying Amota Formation (considered within the Didikama Formation by Beyth 1972; Miller et al. 2003, 2009; Alene et al. 2006) indicate that overlying deposition was younger than 794.2 ± 0.7 Ma (SwansonHysell et al. 2015). The Tambien Group appears to contain at least three negative CIE events, which have been linked to the *800 Ma Bitter Springs event, the *737 Ma Islay anomaly, and the *720 Ma transition into the Sturtian diamictite

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(Fig. 7.1A). Interestingly, the Bitter Springs event has been linked to an episode of true polar wander (Maloof et al. 2006), suggesting a rapid reorientation of Earth’s magnetic field relative to its lithospheric shell. If the correlation is accurate, this perturbation followed deposition of the basal Tambien Group (Werii Slate) in western inlier exposures (Mai Kenetal, Chemit, Tsedia). In eastern inliers (Negash and prospectively Samre), basal Tambien Group deposition (Amota Formation, Negash E. Limb) is younger than 794.3 ± 0.6 Ma (Park et al. 2019) and deposition of lowest dolomite beds of the overlying Didikama Formation (Negash W. Limb) was younger than 774.7 ± 5.7 Ma (Miller et al. 2009). These dates bracket the 788.1 ± 0.2 Ma age of ongoing deposition of the Tsedia Slate in the Tsedia inlier (Swanson-Hysell et al. 2015), which followed recovery from the prospective Bitter Springs negative CIE in the Assem Limestone in Mai Kenetal synclinorium. The other two negative CIE events are currently only resolved in eastern inliers (Negash and Samre). A negative CIE prospectively correlated with the Islay anomaly, which precedes the Sturtian glaciation by >10 myr (Rooney et al. 2014) appears to be reasonable based on age constraints (U– Pb ID–TIMS data of 735.25 ± 0.25 Ma for tuff layers within basal Matheos Formation; MacLennan et al. 2018). However, the anomaly in Negash and Samre inliers begins within the uppermost Didikama Formation and recovers in overlying black limestone of the Matheos Formation. The abrupt lithologic change, previous reporting of a possible angular unconformity (Garland 1980), and abrupt jump in 87 Sr/86Sr (Miller et al. 2009, 2011) across this contact may indicate a considerable hiatus. The uppermost negative CIE occurs in the transition to diamictite deposition. This negative excursion from positive values following recovery from the Islay Anomaly is represented in other Sturtian glacigenic successions (e.g., Halverson et al. 2005; Rooney et al. 2011; Fairchild et al. 2018). Future efforts can greatly clarify Tambien Group regional lithostratigraphy by further constraining the age of units associated with negative CIEs, in least altered successions with minimal structural complications. Opportunities for further study—As a relatively continuous Tonian to Cryogenian succession, greenschist grade metasediments of the Tambien Group offer valuable opportunities to understand the lead up to extreme climate transitions of the Cryogenian. Although Tambien Group deposition occurred in a period of reduced magmatic activity, volcanic ashes occur sporadically throughout the sequence providing samples suitable for absolute age dating. The occurrence of shallow marine carbonate intervals with well-preserved primary depositional textures allows C and Sr isotopic determinations on least altered samples, whereas slate intervals may reveal the extent of chemical weathering.

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The onset of the Cryogenian may have been driven by consumption of atmospheric CO2 and reverse greenhouse cooling associated with weathering of large volumes of juvenile Neoproterozoic crust and/or flood basalts emplaced at low latitudes (Goddéris et al. 2003; Donnadieu et al. 2004). Carbonate and organic carbon in the Tambien Group may constitute a portion of this sequestered greenhouse carbon, and the associated siliciclastic intervals may represent weathering of volcanic arc terranes generated in the Mozambique basin. Preliminary work on thick slate units comprising the lower Tambien Group demonstrates high chemical weathering indices, consistent with Tsaliet arc accretion complexes undergoing protracted and intensive and silicate weathering prior to carbonate deposition (Sifeta et al. 2005; Miller et al. 2009). The contribution of weathering of ANS arc terranes to initiation of Cryogenian glaciation is a topic worthy of future studies. The paleolatitude of Tigrai during Tambien Group deposition is still poorly constrained. Extensive paleomagnetic testing of Matheos Formation limestone in the core of Negash syncline proved unsuccessful due to probable Quaternary remagnetization (Kidane et al. 2014). Further paleomagnetic testing of other inliers may reveal valid depositional paleolatitudes for the Tambien Group.

7.5.2.2 Banded Iron Formation of Egypt and Arabia BIFs are found in the Central Eastern Desert of Egypt and in the Silasia Formation of the Midian region in NW Saudi Arabia (Fig. 7.2, localities 7, 9; Fig. 7.10A–D). These distinctive sedimentary rocks occur as centimeter- to meter-scale interbeds in wackes distributed over a 200  100 km area (Fig. 7.9), possibly representing portions of an extended ANS BIF basin (Stern et al. 2006). They are metamorphosed into greenschist facies. BIF deposition occurred in a marine basin associated with arc/backarc basin volcanism and immature clastic sedimentation. Beds are composed of alternating iron- and silica-rich laminae (Fig. 7.10D), which may reflect seasonal changes in deposition of Fe vs. Si. Fe-rich layers are dominantly composed of primary fine-grained hematite “dust” and minor apatite, with abundant secondary magnetite. Rapid deposition is revealed by (1) major and trace element data indicating that ANS–BIF are very pure (8000; width >1000 km; Jacobs and Thomas 2004) and is estimated to have shed sediments in excess of 100 Mkm3 over a period of >100 Myr, as evidenced by common detrital zircon U–Pb age spectra in flanking terranes of India, Arabia, Africa, Australia, New Zealand, South America, and Antarctica (Squire et al. 2006). Widespread erosional unconformities throughout northern Africa and Arabia, separating Neoproterozoic basement from continentally derived Cambro-Ordovician quartz-rich sediments, testify to massive and sustained erosion of this vast collisional mountain range (Burke and Kraus 2000; Avigad et al. 2003). Fritz et al. (2013) argued that regions in the southern EAO experienced the greatest crustal thickening and uplift; the regions now occupied by Tanzania and Madagascar, in particular, experienced a Himalayan-type orogen with doubly thickened crust and correspondingly great uplift. Earliest orogenesis likely followed from the collision of northeastern Africa with the ANS at *650– 590 Ma (Boger and Miller 2004). By about 630 Ma, much of the ANS was probably above sea level. In orogenic hinterlands of the ANS, sediments that may have been deposited in association with the Marinoan panglacial episode would have been uplifted and increasingly eroded.

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Metamorphic mineral T and P relationships, in addition to 40 Ar/39Ar studies that constrain Neoproterozoic basement uplift (cooling) histories provide regional estimates of the extent of regional mountain range relief and beveling. For example, high-grade kyanite-bearing rocks in east Eritrea (Ghedem Domain) suggest formation at depths exceeding 40 km (Beyth et al. 1997; Ghebreab 1999), implying that the southern EAO had a crustal thickness of *70 km (Beyth et al. 2003). Structurally overlying low-grade metasediments (Bizen Domain), akin to the Tambien Group, underwent burial metamorphism at considerably shallower burial depths. In Ethiopia, greenschist metamorphism of Tsaliet Group arc volcanic rocks in the northern Tigre area likely occurred at depths of 6–8 km (e.g., Asrat et al. 2004), but as a result of subsequent bevelling there is no known record of post-Sturtian sediment accumulation in Ethiopia. In the NE Arabian Shield, Cole (1988) postulated epeirogenic uplift as the cause of major 615–585 Ma erosion. Consistent with this interpretation is 40Ar/39Ar uplift histories from micas that indicate rapid cooling at *600 Ma (Cosca et al. 1999). Most explanations for *600 Ma exhumation invoke tectonic unroofing (e.g., Al-Husseini 2000; Blasband et al. 2000), but we suggest Marinoan and younger Ediacaran glaciation could also account for much of this deep erosion. Unroofing farther south, in Sudan and southern Egypt, may have occurred *570 Ma, and thus could have intersected Ediacaran glaciation (Bailo et al. 2003). Marinoan or Ediacaran glacial episodes in the ANS would have involved continental glaciations, possibly as continental ice sheets. These powerful erosional agents could have rapidly lowered relief of growing EAO mountains, shedding sediments that ultimately were deposited far from the high mountains. Greatest erosion likely occurred in the southern EAO, where high-grade metamorphic terranes indicate relief was highest and unroofing has removed any post-Sturtian sedimentary record, but how and when this region was beveled is not understood. The extent of EAO uplift was not limited to juvenile terranes of the ANS, but also included bounding cratonic terranes of east and west Gondwana. The provenance of sediments shed across northern Gondwana was therefore not exclusively limited to uplifted ANS terranes (e.g., Dor et al. 2018). Any late Cryogenian or Ediacaran marine sediments in the ANS would have been deposited on the margins of consolidated Gondwana or in later-formed structural basins (e.g., Najd Fault Basins; Stern 1985, 1994; Johnson et al. 2011, 2013). In marginal marine settings, Marinoan and Ediacaran glacioeustasy likely played a role in controlling the composition and architecture of sediment packages. As carbonate rocks of this age are uncommon in the ANS, their chemostratigraphic records are of particular importance for comparison with well-calibrated marine basinal records.

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7.5.3.1 Post-amalgamation (Ediacaran) Basins Although definitive evidence of glacial deposition has not yet been reported, ANS depocenters capable of preserving evidence of Marinoan and/or Ediacaran glacigenic sedimentation include post-amalgamation basins of the northern Arabian Shield (Johnson 2003). Candidates capable of preserving evidence of the Marinoan glaciation include Furayh and Murdama basins in the Arabian Shield and buried basins in the Rub-al Khali (Stewart 2016) (Fig. 7.2, localities 14-16). Furayh basin contains 3300 m, is largely volcano-sedimentary but also includes carbonate units. Probable metazoan trace and body fossils are reported in two Jibalah Group basins (Dhaiqa and Jifn) (Vickers-Rich et al. 2013). Volcanic flows and pyroclastic beds demonstrate that the region was magmatically active during deposition. Jibalah Group regional correlation is not straightforward owing to high lithologic variability from basin to basin. Basal fill in many basins begins as polymict conglomerate (Fig. 7.11A) and limestone generally increases in abundance in the upper succession (Vickers-Rich et al. 2013). Matrix supported cobble and boulder diamictite intervals (Fig. 7.11H–J) are reported in at least one basin (Dhaiqa; Miller et al. 2008). Possible dropstones (Fig. 7.11B–C) have also been noted in Jifn Basin (Kusky and Matsah 2003) and in the Naghr Formation (Shagab Quadrangle, NW Arabia; presumed to be correlative with the Jibalah Group (Vickers-Rich et al. 2010). Jibalah Group intervals

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Fig. 7.11 Outcrop photographs of prospective glacigenic intervals in the Ediacaran Jiabalah Group, Arabian Shield. A Polymict basal conglomerate unconformably overlying Murdama rhyolite basement; photo from Kusky and Matsah (2003). B–C Possible dropstones in Jifn (Kusky and Matsah 2003 and Naghr basins (Vickers-Rich et al. 2010; photo from Cui et al. 2020). Photos D–J Dhaiqa basin. D Eastward view of well-bedded carbonate in the Dhaiqa Formation, showing underlying Mataar Formation siliciclastics, positions of metazoa and a diamictite interval, and the Siq Unconformity. E–F Outsized boulders in the ferruginous basal Mataar Formation (R. Stern for scale). G Closer view of Mataar polymictic conglomerate. I Stratigraphic transition from the Mataar Formation to initial carbonate beds of the Dhaiqa Formation. I–J Intra-Dhaiqa diamictite interval

associated with polymict conglomerate, matrix-supported diamictite, and dropstones are prime targets for follow-up studies of possible glacigenic sedimentation. Jibalah Group correlatives could be found in rift basins buried beneath the Rub Al-Khali (Fig. 7.2, locality 19; Stewart 2016) and similar buried basins beneath the Western Desert of Egypt. Jibalah Group sediments may also be correlative with Hammamat sediments of Egypt, Saramuj Conglomerate of Jordan, and Elat Conglomerate of Israel (Fig. 7.2, localities 20–22). An unresolved problem is whether Jibalah Group deposition directly followed basin subsidence from local sediment sources or if the deposition was more regionally sourced and only subsequently preserved in down-dropped

basins (Johnson 2003). The upward increase in carbonate units among basins suggests a marine transgression, but the prior isolated character of each basin could alternatively indicate a series of fault-controlled lakes (Johnson et al. 2003). Basin development could also have been diachronous along different Najd fault zones. The occurrence of metazoan trace fossils demonstrate that Ediacaran animals were present in some Jibalah Group basins, raising intrigue about their origins in marine vs. non-marine environments. Jibalah Group regional deposition is constrained to be younger than underlying shield rocks (  618 Ma; Nettle et al. 2013) and likely  605 ± 5 Ma;) and older than the overlying Lower Cambrian basal unconformity (*540–

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520 Ma) formed by peneplanation of the EAO. U–Pb zircon dating of volcanic intervals in several Jibalah Group basins confirm this depositional interval. Jibalah Group chronostratigraphy is currently best documented in three widely separated basins further detailed below: Jifn (aka Jif’n) basin in the NE Arabian Shield, Antaq basin in the E Arabian Shield, and Dhaiqa basin in the NW Arabian Shield (west of Al ‘Ula). In Jifn Basin (Fig. 7.2, locality 18c), deposition of a *3 km thick sequence of conglomerate, limestone and shale, is constrained by TIMS U–Pb zircon techniques between 625 ± 4 (Murdama Group basement) and 577 ± 5 Ma (felsite dyke cutting Jifn Formation) (Kusky and Matsah 2003). The Jifn Jibalah Group succession above basement consists of Lower Conglomerate (500 m), Umm al-Aisah Limestone (*340 m), and the clastic Jifn Formation (*1870 m), which begins with *290 m of Jifn Polymictic Conglomerate. Interestingly, Kusky and Matsah (2003) show an outcrop photo of a “possible dropstone”, but its position within the stratigraphic succession was not reported and its possible significance was not further explored. Metazoan (multi-cellular; tissue grade organism; string-of-beads like; Horodyskia-like) macrofossils were later discovered in lime-rich dark mudstones of the Jifn formation (Vickers-Rich et al. 2013). The stratigraphic position of this sample within the Jifn formation is also unclear, but the regional interval was described as consisting of brownish conglomerates with well-rounded pebbles, thin grey limestone beds, dark red calcareous sandstones, and repetitive grey clastic sediments (sandstone to shale). Halverson et al. (2013) obtained a LA-ICP-MS U-Pb zircon age of  589.5 ± 0.5 Ma for an ash bed in the lower part of the Umm al-Aisah Limestone and also documented d13C values of 5–8‰ for this unit, which if marine could be consistent with deposition before the onset of the Shuram negative d13C excursion (similar to Khufai Fm of the Nafun Group in Oman; Al-Husseini 2014). In Antaq basin (Fig. 7.2, locality 18a), the Jibalah Group consists of a *2500 m thick conformable succession of nearly undeformed and unmetamorphosed clastic and carbonate sediments, consisting from base to top of Rubtayn, Badayi, and Muraykhah formations (Hadley 1974; Nettle et al. 2013). The succession is constrained by CA–TIMS (Nettle et al. 2013) and SHRIMP (Kennedy et al. 2010a, b) U–Pb zircon dates of Antaq Basement to be younger than 618–597 Ma. Nettle et al. (2013) established LA–ICP–MS U–Pb zircon dates for ash beds within the overlying Jibalah group of  596 ± 17 Ma (Rubtayn Fm, above Polymictic Conglomerate); 579 ± 17 Ma (Muraykhah Fm, 255 m); and  604 ± 18 Ma (Muraykhah Fm, 325 m). Carbonate d13C values for three sections of the Muraykhah Fm mainly span from −4 to 0‰, with considerable stratigraphic scatter (Nettle et al. 2013). If Muraykhah deposition occurred in a

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marine environment, the modestly negative d13C values could conceivably intersect the Shuram negative d13C excursion, possibly during its recovery as demonstrated in the Buah Formation of the Nafun Group in Oman (Fike et al. 2006; Nettle et al. 2013; Al-Husseini 2014). The *1000 m thick Antaq Polymictic Conglomerate interval within the Rubtayn Fm is a prospective candidate for glacigenic deposition. It is described as blocky, with occasional cross-beds, predominantly consisting of sub-rounded white quartz clasts in a quartz-feldspathic matrix (Nettle 2009). The 596 ± 17 Ma date just above this unit (Muraykhah Fm, 255 m) could mean the conglomerate was deposited before this time; possibly during the *580 Ma Gaskiers glaciation. In Dhaiqa Basin (NW Arabian Shield, Fig. 7.2, locality 18f, Fig. 7.11D–J), near the confluence of Wadi al Jizl and Wadi Dayqah, the Jibalah Group begins above basement as the Mataar Formation (Davies 1985), a *150 m thick siliciclastic unit, consisting of poorly sorted arkose and conglomerate bearing subangular to subrounded clasts, including matrix-supported outsized boulders with up to m-scale diameters, which grades upward into arkose and lithic arenite (Fig. 7.11E–G). The Mataar Formation is gradationally and conformably overlain by the Dhaiqa Formation (Davies 1985), a 300–400 m thick carbonate succession, consisting of low-relief microbial and algal structures with evidence of subaerial exposure (desiccation cracks) and high energy deposition (carbonate intraclast rip-ups/tempesites?) in the lower portion. Miller et al. (2008) logged the stratigraphic section, established baseline Sr and C isotope stratigraphies, and documented the possible occurrence of metazoan trace fossils within the middle portion of the Dhaiqa Formation. Trace fossils have since been verified and probable Ediacaran body fossils discovered at higher levels in the Dhaiqa Formation (Vickers-Rich et al. 2010, 2013). Available U–Pb zircon ages support that Dhaiqa Jibalah Group deposition occurred between  609 and *530 Ma and was ongoing at 570–560 Ma in the upper portion of the Dhaiqa Formation. The undated Mataar Formation overlies granitoids having a SHRIMP U–Pb zircon age of 609 Ma (Kennedy et al. 2010a). No upper contact is preserved in the basin, but the Dhaiqa Formation certainly unconformably underlies the Cambrian Siq Sandstone below the *530 Ma Sub-Siq Unconformity, which correlates regionally with the Ram Unconformity in Jordan and Angudan Unconformity in Oman (Al-Husseini 2014; Powell et al. 2015). Two stratigraphic intervals within the Dhaiqa Formation provide further age constraints. A distinct 2–3 m-thick interval of glauconitic arenite overlain by poorly sorted polymict conglomerate (Intra-Dhaiqa Diamictite) with clast diameters up to 0.5 m (Fig. 7.11I–J), interrupts carbonate deposition in the upper Dhaiqa Formation (*180 m). Detrital zircons recovered from a fine-grained sandstone interval have

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SHRIMP U–Pb ages as young as 599 ± 4.8 Ma and 570 ± 4.6 Ma (measured from the core and rim of the same grain, respectively; Miller et al. 2008). Above the diamictite interval, in a different area of the basin, Vickers-Rich et al. (2010) determined detrital zircon LA–ICP–MS U–Pb ages between 837 ± 25 and 555 ± 15 Ma for a suspected re-deposited volcanic ash bed. An age of 560 ± 4 Ma obtained for youngest zircons (n = 17) was interpreted as the maximum depositional age for the Dhaiqa Formation at this level. Youngest detrital zircon ICP–MS U–Pb ages of 569 ± 3 Ma were obtained for a tuff at a level above the diamictite interval in yet another locality (Vickers-Rich et al. 2013). These dates demonstrate mid-to-upper Dhaiqa limestone deposition at 570–560 Ma. Metazoan trace fossils (Miller et al. 2008) and probable body fossils (Vickers-Rich et al. 2010, 2013) have been found in the Dhaiqa Formation, including Beltanelloides/ Nemiana, a Pteridinium-like impression, a Cyclomedusa/ Aspidella-like holdfast structure, rod-shaped fossils with oblique transverse marks similar to Harlaniella sp., and curvilinear conical tubes suggestive of frond-like taxa with a basal holdfast (e.g., Charnia). Lowest documented trace fossils occur *30 m below the level of the Intra-Dhaiqa Diamictite, indicating that they are older than *570– 560 Ma, if the dated volcanic units (prospective reworked ash and tuff) higher in the Dhaiqa sequence are eruptive (depositional) ages. On the basis of sedimentology (matrix-supported polymictic clasts and boulders), both the Mataar Formation and Intra-Dhaiqa Diamictite are candidates for glacigenic deposits. If so, the Intra-Dhaiqa Diamictite, constrained to be  570 Ma, must correspond to a glacial interval younger than the *580 Ma Gaskiers glaciation (Pu et al. 2016), whereas the Mataar Formation (  609–570 Ma) could correspond to the Gaskiers glaciation. Seawater C- and Sr-isotope evolution curves generated from least-altered marine carbonates indicate that Ediacaran oceans were characterized by large perturbations and systematic trends that provide a means for unambiguous regional correlation between marine sections. Sr isotope values of Dhaiqa formation carbonates (0.704–0.706), however, are well-below Ediacaran seawater values (Miller et al. 2008), indicating that the depositional setting may have been isolated (lake?) or that primary isotopic compositions were reset by diagenetic fluid exchange involving a predominant ensimatic Sr source—compatible with regional basement and Ediacaran volcanism, or hydrothermal alteration by juvenile fluids (Cui et al. 2020). C-isotopes in the Dhaiqa Formation, which are less susceptible to alteration, are predominantly positive (2.4 ± 2.3‰). If marine, the d13C compositions constrain deposition either before or after the long-lived Shuram negative d13C excursion. The Gaskiers glaciation is now thought to have slightly preceded the negative Shuram

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anomaly (574 ± 4.7 to 567.3 ± 3.0 Ma; Rooney et al. 2020). The available age control and positive d13C compositions support that Dhaiqa deposition likely followed the Shuram anomaly, and the Intra-Dhaiqa Diamictite interval is younger than the Gaskiers glaciation. The underlying Mataar Formation remains a candidate for a pre- *570 Ma glaciation. Establishing depositional ages for Mataar Formation and extending the d13C database to include transitional carbonate beds grading into the Dhaiqa formation, which might have overlapped with the negative Shuram CIE, could further constrain regional correlations. Until then, distinguishing the lacustrine, paralic, or marine depositional setting of the Dhaiqa Formation remains to be resolved. Ediacaran successions with similar lithologies and thicknesses, but lacking U–Pb zircon geochronology, occur in the Mashhad area near Al Ula, about 100 km ESE of Dhaiqa Basin, in addition to other basins further ESE. In the Jabal Rubtayn section, (Fig. 7.2, locality 18e) Hadley (1974) described the Mashhad area Ediacaran succession as consisting of Rubtayn (375 m), Badayi (120–150 m) and Muraykhah (330–370 m) formations. The succession is unconformably bracketed between the Shammar Rhyolite Group (*620–585 Ma, Al-Husseini 2011) and lower Cambrian Siq Sandstone. The Rubtayn Formation (375 m) consists of ascending boulder conglomerate, sandstone, red beds, and pebble conglomerate units (members), before being interrupted by m-scale amygdaloidal basalt flows of the Badayi Formation. The sedimentary succession resumes conformably in the Muraykhah Formation with deposition of basal conglomerate (similar to the upper Rubtayn Formation), followed by lower cherty (non-dolomitic) carbonate, red siltstone and mudstone, and an upper dolomitic (chert-poor) carbonate unit. Except for the volcanic interlude, the stratigraphic succession at Jabal Rubtayn is similar to the Mataar and Dhaiqa Formations in Dhaiqa Basin (Al-Husseini 2014). In the Sahl Al Matran section (Fig. 7.2, locality 18d, *70 km ESE of Mashhad; SE Sahl Al Matran Quadrangle) Hadley (1986) described the comparable sequence as three informal members: (1) Volcanic Conglomerate (700 m), (2) Polymictic Conglomerate (1500 m), and (3) Sandstone (1000 m). In Bir Sija Basin (Fig. 7.2, locality 18b, *590 km ESE of Mashhad; central Afif Quadrangle), the undated Jibalah Group also resides unconformably above the Shammar Rhyolite Group in eight units (Letalenet 1979): (1) Conglomerate with Shammar pebbles (500 m); (2) Andesite-Basalt (150 m); (3) Sandstone (200 m); (4) Polymictic Conglomerate (150 m); (5) Limestone (20 m), the “Bir Sija Limestone”; (6) Sandstone and Mudstone (700 m); (7) Sandstone and Limestone (300 m); and (8) Sandstone and Siltstone (m). Al-Husseini (2014) suggests that the Dhaiqa and Mashhad sequences begin with deposition of units comparable to Sija Basin unit 4 (Polymictic Conglomerate).

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Recognizing that Ediacaran glacioeustasy should have had a strong effect on the character and sequence of sediments (supersequences) deposited on Gondwana margins, Al-Husseini (2014) proposed that Ediacaran strata in the Nafun Group (Huqf Supergroup) of Oman may correlate to Jibalah Group deposits in the northern Arabian Shield (Fig. 7.12). The correlation assumes that a major glacial episode, perhaps Gaskiers, caused a widespread unconformity, which may be related to the Great Unconformity (see discussion in Sect. 6.3.3). Within the Huqf Supergroup, the Ediacaran Nafun Group of Oman (635–547 Ma) is divided into Lower (635–582 Ma, and Upper (582–547 Ma) Supersequences by an unconformity at the base of the Shuram Formation and the long-lived (e.g. *582–551 Ma, Bowring et al. 2007; *570–557 ± 3, Zhou et al. 2018; 574 ± 4.7–567.3 ± 3.0 Ma, Rooney et al. 2020) Shuram negative d13C excursion (Condon et al. 2005; Le Guerroué 2010; Gong et al. 2017; Shields et al. 2019). The age of this unconformity is poorly constrained and controversial (see Al-Husseini 2014 for details); it is broadly related to the Great Unconformity (see Sect. 6.3.3 for further discussion). Assuming constant sedimentation rates for the thickest known Nafun Group sequence (2308 m; offshore Masirah-1 Well), Al-Husseini (2014) estimated the unconformity to have formed close to *582 Ma, consistent with the known age of the Gaskiers glaciation in Newfoundland (Bowring et al. 2003, 2007). He suggested this putative glacial unconformity should be traceable to comparable “glacial” unconformities within the Jibalah Group and proposed that the Rubtayn Boulder Conglomerate (Jabal Rubtayn, Mashhad area), Mataar Polymict Conglomerate (Dhaiqa Basin), Antaq Polymict Conglomerate (Antaq Basin), and Jifn Polymict Conglomerate (Jifn Basin) are products of the same Ediacaran glaciation, and that limestone units overlying some of these units could be cap carbonates (Fig. 7.12B–G). In Jifn Basin, the proposed contact occurs within the Jifn Formation, separating Umm al-Aisah Limestone of the “lower Jibalah Supersequence” (605 ± 5–582 Ma) from Jifn Polymictic Conglomerate (  200 m thick) at the base of the Upper Jibalah Supersequence. In Antaq Basin, the contact is placed at the base of the Antaq Polymictic Conglomerate, based on the reasoning that immediately overlying age constraints (613–579 Ma) permit its deposition to have been synchronous with the *582 Ma Gaskiers glaciation and d13C values (−4 to 0‰) in the overlying Muraykhah Formation being compatible with Buah Formation (Upper Nafun Supersequence, Oman) deposition at the end of the negative Shuram anomaly. In Dhaiqa Basin, the unconformity is placed at the base of the Mataar Formation, constrained to be 570 Ma, consistent with overlying Dhaiqa Formation carbonate deposition, with predominantly positive d13C values, being equivalent to Buah Formation (Oman), following the Shuram negative

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d13C excursion (Miller et al. 2008). He suggests the composite Mataar-Dhaiqa succession is a second-order sequence, “upper Jibalah Supersequence’ that is correlative to the Upper Nafun Supersequence (582–547) in Oman. This intriguing correlation scheme needs additional work to assess its feasibility. Considerations for follow up studies include that the Masirah-1 well unconformity, with estimated age of 582 Ma (Al-Husseini 2014) has not been directly dated and could have a different age. In fact, the Gaskiers glaciation in Newfoundland is now constrained to a 340 kyr duration between 579.63 ± 0.15 and 579.88 ± 0.44 Ma; Pu et al. 2016), and it may have been one of several regional-scale mid-to-late Ediacaran glacial episodes on eight paleocontinents. Age constraints for many of these are limited, but sufficient to demonstrate multiple Ediacaran glacial episodes (e.g., Hebert et al. 2010; Vernhet et al. 2012; Etemad-Saeed et al. 2016; Linneman et al. 2018). There is some support for Ediacaran glaciation in northern peri-Gondwana prior to the appearance of Metazoa. Evidence of 605 to *560 Ma glaciation is reported in Morocco (Vernhet et al. 2012) and Late Ediacaran glaciation (*  560 Ma) is suspected elsewhere on the West African Craton (Bertrand-Sarfati et al. 1995; Caby and Fabre 1981; Deynoux et al. 2006). Linneman et al. (2018) recently documented *565 Ma glaciomarine diamictites along the periphery of the West African Craton (SW Iberia, Bohemia). Etemad-Saeed et al. (2016) reported possible  560 Ma glacial diamictite, overlying striated pavement, within the Kahar Formation of northern Iran, which they correlated within the upper Nafun Group Supersequence (Buah and Ara formations) of Oman. Prospective Ediacaran glacigenic units have also been documented in central (Mohsensi and Aftabi, 2015) and southern Iran (Aftabi 2001; Hassanlouei and Rajabzadeh 2019). These could all correlate within the Upper Jibalah Supersequence of Al-Husseini (2014). Evidence in support of Ediacaran glaciation in the northern ANS is currently limited to the occurrence of diamictites and polymict conglomerate units, some overlain by carbonate intervals (prospective cap carbonates), and rare documentation of possible dropstones within Ediacaran age Jibalah Group basins. Needed are more rigorous sedimentologic studies to confirm or reject glacigenic (post-glacigenic) associations with these units. The limited number and thickness of carbonate units, and possibility that Jibalah Group basins were non-marine or had limited marine connections during deposition, pose significant limitations for global chemostratigraphic correlations. The common stratigraphic occurrence of tuffs and ashes in Jibalah Group successions offers the potential for refined age control.

7.5.3.2 Gondwanan Margin Basins Other northern ANS sedimentary successions could represent glacial deposits or fluvial reworking of sediments from

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b Fig. 7.12 Prospective regional correlation of Al-Husseini (2014, 2015) for Ediacaran and early Cambrian units between Oman and Jibalah Group

basins of the Arabian Shield. A The correlation attributes the erosional unconformity at the base of the Shuram Formation (Masirah-1 well, offshore Oman) to the *581–580 Ma Gaskiers glacial interval and postulates that similar erosional unconformities should exist within Jibalah Group sedimentary successions in the Arabian Shield. B–G Proposed Gaskiers-related glacigenic sediments in Jibalah Group basins include: polymict conglomerate intervals in Dhaiqa Basin (Mataar Fm), Mashad region (Rubtayn Boulder member of the Rubtayn Formation), Jifn Basin (basal Jifn Fm), Antaq Basin (upper Rubtayn Fm), and Bir Sija Basins (Unit 4 polymict conglomerate). In this correlation, the proposed Upper Jibalah Group supersequence is equivalent to the Upper Nafun Group supersequence in Oman. Occurrence of possible dropstones in unspecified intervals of the Jifn Formation (Jifn Basin, Kusky, and Matsah 2003) and Nagr Formation (Vickers-Rich et al. 2010), in addition to the Intra-Dhaiqa diamictite (Miller et al. 2008) raise the possibility of post-Gaskiers glaciation in the ANS. Modified from Al-Husseini (2015)

Marinoan and/or Ediacaran glaciers. There seems to have been a major reorganization of drainage following the *530 Ma Rum Unconformity. After this time drainage was dominated by vast fluvial systems that transported and deposited huge sediment volumes in giant sedimentary fans that seem to have been mostly transported to the north (Dabbagh and Rogers 1983). Before this time, drainage was more local, including *600 Ma deposition of the Saramuj conglomerate of Jordan (Fig. 7.2, locality 21), which Jarrar et al. (1991) interpreted as deposited within a high velocity, braided stream/alluvial fan system, the Hammamat Group in NE Egypt (Fig. 7.2, locality 20; Jarrar et al. 1993; Wilde and Youssef 2002), and Jibalah Group sediments (Fig. 7.2, locality 18a–g). Pre-Rum Unconformity coarse sediments could be periglacial tillites of Marinoan or Ediacaran age, reworked by meltwater streams as glaciers receded, but deliberate studies are required to confirm or refute this suggestion. The >588 ± 10 Ma Elat Conglomerate of southern Israel (Fig. 7.2, locality 22) contains clasts up to 1.5 m and was deposited on a deeply dissected relief, suggesting that sea level was quite low (Be’eri-Shlevin 2008; Weissbrod and Sneh 2002). This timing is close to prospective glacial intervals in Morocco, SW Iberria, Bohemia, and N. Iran (Vernhet et al. 2012; Linneman et al. 2018; Etemad-Saeed et al. 2016), which would also have been situated along the northern Gondwana margin (present coordinates). The possibility that the Elat Conglomerate was deposited in a glacial or periglacial setting during an Ediacaran glacial episode, younger than the Gaskiers, is worthy of further study. The >2 km thick Zenifim Formation of Israel (Fig. 7.2, locality 23) is constrained to have been deposited between *550 and 600 Ma (Abdo et al. 2020). The Zenifim can be expected to record any glaciation impacting the area during this time but has not yet been studied for this purpose.

7.5.3.3 Role of Glaciation in the Formation of the Afro-Arabian Peneplain The Great Unconformity is a profound gap in Earth’s stratigraphic record often evident below the base of the Cambrian system. Keller et al. (2020) argue that it was caused by multiple episodes of global glaciation that removed a global average of 3–5 vertical kilometers of crust

and sediments. The age of the Great Unconformity in most places is difficult to constrain, but because tectonomagmatic activity and sedimentation continued throughout Neoproterozoic time, the age of Great Unconformity can be well-constrained in the ANS and environs (Fig. 7.2, localities 17 and 24—arbitrarily placed along the southern exposure limit of early Paleozoic sandstone; Fig. 7.13). The Jordan record, in particular, suggests that it formed as a result of multiple stages of downcutting over *100 million years. The Ediacaran Araba Complex of Jordan is bracketed by two major erosional unconformities: the basal Araba Unconformity overlies *605 Ma rocks and the overlying Ram unconformity, which is thought to have formed *530 Ma (Powell et al. 2015). An unconformity like the Araba unconformity is also recognized in southern Israel, where Garfunkel (1999) documented a “Main Erosion Phase” at *600 Ma, possibly involving 8–14 km of downcutting. Ediacaran glaciations may have contributed to downcutting of the Araba and Ram unconformities in Jordan. The compatible timing of the *580 Ma Sub-Shuram Unconformity, separating the Khufai and Shuram formations in subsurface Oman, with deposition of conglomeratic units in the Jibalah Group of Arabia (Al-Husseini 2014; as discussed in Sect. 6.3.1), may further support regional scale glacial erosion linked to the Gaskiers and possibly younger Ediacaran glaciations along the peri-Gondwana margin. Formation of the Great Unconformity in and around the ANS is likely related to alpine glaciation in the Transgondwanan Supermountain and set the stage for establishment of the peripheral superfan systems. Glacial activity was unlikely for final cutting as the presence of thick laterite immediately below the peneplain indicates a warm and humid climate (laterization could happen after glaciation). Erosion of this exceptional regional orogen would have been enhanced by the lack of terrestrial vegetation and also increased water-rock interaction (chemical weathering) in lower latitudes approaching the paleoequator (Squire et al. 2006). Orogen development would have transgressed both the Marinoan panglacial episode as well as regional scale Ediacaran glaciation, and its weathering intensity could have influenced carbon burial and greenhouse gas concentrations, perhaps even triggering glacial episodes. These continental glaciations would have greatly enhanced erosion and moved

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Fig. 7.13 Expressions of the Great Unconformity (Afro-Arabian Peneplain) in the Arabian-Nubian Shield. A Southern Jordan (Wadi Ram). Ram Unconformity separating peneplained Neoproterozoic basement (Aqaba Complex granitoids) from overlying lower Cambrian sandstone (Ram Group). Photo from Powell et al. (2015). B Northern Arabia (25 km west of Al ‘Ula) Siq Unconformity separating beveled Neoproterozoic granitic basement from overlying Cambrian-Ordovician sandstone. Horizontal field of view is *130 m. c Northern Ethiopia (Adigrat region)—AAP separating Neoproterozoic basement from overlying Cambro-Ordovician sandstone

glacial till to lower elevations and ultimately marine margins. Alpine glacial environments are intrinsically transient, and preservation of Marinoan and Ediacaran glacial deposits may only have been possible in distal margin settings or within down-dropped basins. Some deposits in lowland settings may have developed secondarily from reworking of earlier glacigenic sediments. How the basal Cambrian

peneplain formed has not been well studied, and the possible role of glaciation in its early cutting should be considered further. The final episode in the latest Ediacaran time after final cutting of the peneplain was the establishment of a hot climate. This is shown by evaporites in eastern Arabia and Oman (Fig. 7.12A), where ash beds in evaporites of the Ara

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Formation give U–Pb zircon ages of 541, 542, and 547 Ma (Bowring et al. 2007). This warm climate is also shown by a thick soil horizon at the base of the Great Unconformity in Israel (Sandler et al. 2012).

7.6

Conclusions

In many ways, the sedimentary and erosional history of the ANS has been ignored in international efforts to reconstruct the causes and events of Neoproterozoic glaciations. In spite of this, the diverse sedimentary environments and contemporary igneous activity combine to make the ANS a potentially excellent place to independently constrain Neoproterozoic climate evolution. There is some evidence for Tonian glaciation in the ANS, including *770 Ma diamictite deposits on the southern margin of the Bi’r Umq-Nakasib Suture Zone in Sudan (Meritri Group) and Arabia (Mahd Group), but studies focused on these rocks are needed in order to confirm or refute this possibility. The *717–659 Ma Sturtian glacial episode occurred when the ANS was mostly marine, so evidence for this episode is expected to be preserved in ANS rocks. There is increasing evidence for Sturtian glaciation in the form of diamictites in Egypt, Saudi Arabia, and Ethiopia and Banded Iron Formations of Egypt and Saudi Arabia. We have made good progress in constraining their age and understanding their significance, but further research is needed. There is an especially important opportunity in Sudan, where polymictic conglomerates of the Meritri Group in the Nakasib Suture Zone need to be studied to determine if these are glacigenic and if they are pre-Sturtian or Sturtian deposits. Evidence for Marinoan and Ediacaran glaciations is less well preserved in the ANS because this was a time that the ANS was rising out of the sea as a result of continent-continent collision and formation of a huge mountain range in the EAO to the south, so erosion dominated over deposition. Consequently, evidence for Marinoan and Ediacaran glacial episodes is less clear. There was certainly glaciation in the high Transgondwanan Supermountains during the Marinoan panglacial episode and other times in the Ediacaran but when and where any of these glaciers reached the ANS is unclear. It is likely that much of the Great Unconformity was cut by multiple episodes of Marinoan and Ediacaran glaciation, including the *605 Ma Araba Unconformity and *530 Ma Rum Unconformity. Such evidence might be preserved in Ediacaran basins in the Arabian Shield, but such evidence must be sought for intentionally. Acknowledgements We thank Peter Johnson and Yasser Abd El-Rahman for thoughtful reviews that improved the manuscript.

References Abd El-Rahman Y, Gutzmer J, Li X-H, Siefert T, Li C-F, Ling X-X, Li J (2020) Not all Neoproterozoic iron formations are glaciogenic: sturtian-aged non-Rapitan exhalative iron formations from the Arabian-Nubian Shield. Miner Deposita 55:577–596 Abdelsalam MG, Stern RJ (1993) Tectonic evolution of the Nakasib suture, Red Sea Hills, Sudan: evidence for a late Precambrian Wilson Cycle. J Geol Soc London 150:393–404 Abdelsalam MG, Stern RJ (1996) Sutures and shear zones in the Arabian-Nubian shield. J Afr Earth Sc 23:289–310 Abdelsalam MG, Stern RJ, Copeland P, El Faki EM, El Hur B, Ibrahim FM (1998) The Neoproterozoic Keraf Suture in NE Sudan: Sinistral transpression along the eastern margin of West Gondwana. J. Geology 106:133–147 Abdelsalam MG, Abdel-Rahman EM, El-Faki EM, Al-Hur B, El-Bashier FRM, Stern RJ, Thurmond AL (2003) Neoproterozoic deformation in the northeastern part of the Saharan Metacraton, northern Sudan. Precambr Res 123:203–221 Abdo A, Avigad D, Gerdes A, Morag N, Vainer S (2020) Cadomian (ca. 550 Ma) magmatic and thermal imprint on the North Arabian-Nubian Shield (south and central Israel): new age and isotopic constraints. Precambr Res 346:105804 Aftabi A (2001) Introduction of Bandar Abbas iron ore as the newest type of Rapitan banded iron formation. J Min Metall 60–70:53–66 (in Persian with English abstract) Alene M, Jenkin GRT, Leng MJ, Darbyshire DP (2006) The Tambien Group, Ethiopia: an early Cryogenian (ca. 800–735 Ma) Neoproterozoic sequence in the Arabian-Nubian Shield. Precambr Res 149:79–89 Alessio B, Blades M, Murray G, Thorpe B, Collins AS, Kelsey D, Foden JD, Payne J, Al-Khirbash S, Jourdan F (2017) Origin and tectonic evolution of the NE basement of Oman: a window into the Neoproterozoic accretionary growth of India? Geol Mag 155:1150– 1174 Al-Husseini MI (2000) Origin of the Arabian Plate structures: Amar collision and Najd rift. GeoArabia 5:527–542 Al-Husseini MI (2011) Ediacaran-Cambrian middle east geologic time scale 2014—late ediacaran to early Cambrian (Infracambrian) Jibalah Group of Saudi Arabia. GeoArabia 16:69–90 Al-Husseini MI (2014) Ediacaran-Cambrian middle east geologic time scale 2014—proposed correlation of Oman’s Abu Mahara Supergroup and Saudi Arabia’s Jibalah Group. GeoArabia 19:107–132 Al-Husseini MI (2015) Middle east geologic time scale 2015 Ediacaran and Cambrian Periods. J Middle East Petrol Geosci 20:511–528 Ali KA, Stern RJ, Manton WI, Kimura J-I, Khamees HA (2009) Geochemistry, Nd isotopes and U–Pb SHRIMP zircon dating of Neoproterozoic volcanic rocks from the Central Eastern Desert of Egypt: new insights into the *750 Ma crust-forming event. Precambr Res 171:1–22 Ali KA, Stern RJ, Manton WI, Johnson PR, Mukherjee SK (2010a) Neoproterozoic diamictite in the Eastern Desert of Egypt and Northern Saudi Arabia: evidence of *750 Ma glaciation in the Arabian-Nubian Shield. Int J Earth Sci 90:705–726 Ali KA, Stern RJ, Manton WI, Kimura J-I, Whitehouse M, Mukherjee SK, Johnson PR, Griffin WR (2010b) Geochemical, U–Pb zircon and Nd isotopic investigations of the Ghawjah Metavolcanic rocks, Northwestern Saudi Arabia. Lithos 120:379–392 Allen PA (2007) The Huqf Supergroup of Oman: basin development and context for Neoproterozoic glaciation. Earth-Sci Rev 84:139– 185 Allen PA, Rieu R, Etienne JL, Matter A, Cozzi A (2011a) The Ayn formation of the Mirbat Group, Dhofar, Oman. In: Arnaud E et al. (eds) The geological record of Neoproterozoic glaciations. Geological Society of London Memoir 36, 239–249

7

Evolution of the Arabian Nubian Shield and Snowball Earth

Allen PA, Leather J, Brasier MD, Rieu R, Mccarron M, Le Guerroué E, Etienne JL, Cozzi A (2011b) The Abu Mahara Group (Ghubrah and Fiq formations), Jabal Akhdar, Oman. In: Arnaud E et al. (eds) The geological record of Neoproterozoic glaciations. Geological Society of London Memoir 36, 251–261 Antcliffe J (2013) Questioning the evidence of organic compounds called sponge biomarkers. Paleontology 56:917–925 Arkin Y, Beyth M, Dow DB, Levitte D, Haile T, Hailu T (1971) Geological map of Mekele Sheet area ND37-11 Tigre province. Imperial Ethiopian Government, Ministry of Mines. Geological Survey of Ethiopia. Scale 1:250,000 Asmerom Y, Jacobsen SB, Knoll AH, Butterfield NJ, Swett K (1991) Strontium isotopic variations of Neoproterozoic seawater: implications for crustal evolution. Geochim Cosmochim Acta 55:2883– 2894 Asrat A, Barbey P, Ludden JN, Reisberg L, Gleizes G, Ayalew D (2004) Petrology and isotope geochemistry of the Pan-African Negash Pluton, northern Ethiopia: mafic-felsic magma interactions during the construction of shallow-level calc-alkaline plutons. J Petrol 45:1147–1179 Avigad D, Gvirtzman Z (2009) Late Neoproterozoic Arabian—rise and fall of the northern Nubian Shield: the role of lithospheric mantle delamination and subsequent thermal subsidence. Tectonophysics 477:217–228 Avigad D, Kolodner K, McWilliams M, Persing H, Weissbrod T (2003) Origin of northern Gondwana Cambrian sandstone revealed by detrital zircon SHRIMP dating. Geology 31:227–230 Avigad D, Sandler A, Kolodner K, Stern RJ, McWilliams M, Miller N, Beyth M (2005) Mass Cambro production of Ordovician quartz-rich sandstone as a consequence of chemical weathering of Pan-African terranes: environmental implications. Earth Planet Sci Lett 240:818–826 Avigad D, Stern RJ, Beyth M, Miller N, McWilliams M (2007) Detrital zircon U–Pb geochronology of Cryogenian diamictites and lower Palaeozoic sandstone in Ethiopia (Tigrai): age constraints on Neoproterozoic glaciation and crustal evolution of the southern Arabian-Nubian Shield. Precambr Res 154:88–106 Bailo E, Schandelmeier H, Franz G, Sun C-H, Stern RJ (2003) Plutonic and metamorphic rocks from the Keraf Suture (NE Sudan): a glimpse of the tectonic evolution of the NE margin of W. Gondwana during Neoproterozoic time. Precambr Res 123:67–80 Bartley JK, Semikhatov MA, Kaufman AJ, Knoll AH, Pope MC, Jacobsen SB, (2001) Global events across the Mesoproterozoic-Neoproterozoic boundary: C and Sr isotopic evidence from Siberia. Precambr Res 111:165–202 Be’eri-Shlevin Y (2008) The origin and evolution of neoproterzoic magmatism in the northern Arabian-Nubian-Shield (Sinai Peninsula, Egypt, and southern Israel): evidence from the stable and radiogenic isotope record. Thesis (Ph. D.)-Ben-Gurion University of the Negev, Department of Geological and Environmental Sciences, 2008, Beer-Sheva Bentor YK (1985) The crustal evolution of the Arabo-Nubian massif with special reference to the Sinai peninsula. Precambr Res 28:1–74 Bertrand-Sarfati J, Moussine-Pouchkine A, Amard B, Ait Kaci Ahmed A (1995) First Ediacaran fauna found in Western Africa and evidence for an early Cambrian glaciation. Geology 23:133– 136 Beyth M (1972) The geology of central western Tigre, Ethiopia. PhD thesis, Bonn, University of Bonn Beyth M, Stern RJ, Matthews A (1997) Significance of highgrade metasediments from the Neoproterozoic basement of Eritrea. Precambr Res 86:45–58 Beyth M, Avigad D, Wetzel HU, Matthews A, Berhe SM (2003) Crustal exhumation and indications for snowball Earth in the East

187 African Orogen: North Ethiopia and East Eritrea. Precambr Res 123:187–201 Bibolini A (1920) Risultali preliminary delle osservazioni faite nel Nord-est della Colona Eritrea. Asmara Bibolini A (1921) Sui conglomerati di Rore Babla e dei Monti Haggar in Colonia Eritrea. Bollettino della Società Geologica Italiana 40:169–176 Bibolini A (1922) Contributions a l’étude de la géologi de l’Afrique orientale Italienne. 13th International Geological Congress (1922, Brussels, Belgium). Title Comptes rendus de la XIIIe session, en Belgique, parts 1–3:797–814 Blasband B, White S, Brooijmans P, de Brooder H, Viser W (2000) Late Proterozoic extensional collapse in Arabian-Nubian Shield. J Geol Soc London 157:615–628 Boger SD, Miller JM (2004) Terminal suturing of Gondwana and the onset of the Ross-Delamerian Orogeny: the cause and effect of an Early Cambrian reconfiguration of plate motions. Earth Planet Sci Lett 219:35–48 Bold U, Smith EF, Rooney AD, Bowring SA, Buchwaldt R, Dudás FÖ, Ramezani J, Crowley JL, Schrag DP, Macdonald FA (2016) Neoproterozoic stratigraphy of the Zavkhan terrane of Mongolia: the backbone for Cryogenian and early Ediacaran chemostratigraphic records. Am J Sci 316:1–63 Bonavia FF, Chorowicz J (1992) Northward expulsion of the Pan-African of northeast Africa guided by a reentrant zone of the Tanzania Craton. Geology 20:1023–1026 Bowring SA, Myrow PM, Landing E, Ramezani J (2003) Geochronological constraints on terminal Neoproterozoic events and the rise of metazoans. Geophys Res Abstr 5(13219):219 Bowring SA, Grotzinger JP, Condon DJ, Ramezani J, Newall MJ, Allen PA (2007) Geochronologic constraints on the chronostratigraphic framework of the Neoproterozoic Huqf Supergroup, Sultanate of Oman. Am J Sci 307:1097–1145 Brasier MD, Shields G, Kuleshov VN, Zhegallo EA (1996) Integrated chemo-and biostratigraphic calibration of early animal evolution: Neoproterozoic-Early Cambrian of southwest Mongolia. Geol Mag 133:445–485 Brasier M, McCarron G, Tucker T, Leather J, Allen P, Shields G (2000) New U–Pb zircon dates for the Neoproterozoic Ghubrah glaciation and for the top of the Huqf Supergroup, Oman. Geology 28:175– 178 Burke K, Kraus JU (2000) Deposition of immense Cambro-Ordovician sandstone bodies, now exposed mainly in N. Africa and Arabia, during the aftermath of the final assembly of Gondwana. Geol Soc Am Abstr Program 32:249 Burke K, Sengör C (1986) Tectonic escape in the evolution of continental crust. In: Reflection seismology—the continental crust. Am Geophys Union, Geodynamic series 14:41–53 Burns SJ, Matter A (1993) Carbon isotopic record of the latest Proterozoic from Oman. Eclogae Geol Helv 86:595–607 Bussert R (2010) Exhumed erosional landforms of the Late Palaeozoic glaciation in northern Ethiopia: indicators of ice-flow direction, palaeolandscape and regional ice dynamics. Gondwana Res 18:356–369 Butterfield NJ (2015) The neoproterozoic. Curr Biol 25:R859–R863 Caby R, Fabre J (1981) Late proterozoic to early palaeozoic diamictites, tillites and associated glacigenic sediments in the Série Pourprée of western Hoggar, Algeria. In: Harland WB, Hambrey MJ (eds) Earth’s pre-pleistocene glacial record, pp 140–145. Cambridge University Press Calvez JY, Kemp J (1982) Geochronological investigations in the Mahd Adh Dhahab Quadrangle, Central Arabian Shield. Deputy Ministry for Mineral Resources: Jiddah, Saudi Arabia, BRGMTR-02-5, pp 1–41

188 Canfield DE (2005) The early history of atmospheric oxygen: homage to Robert M. Garrels. Annu Rev Earth Planet Sci 33:1–36 Canfield DE, Poulton SW, Narbonne GM (2007) Late-neoproterozoic deep-ocean oxygenation and the rise of animal life. Science 315:92– 95 Canfield DE, Poulton SW, Knoll AH, Narbonne GM, Ross G, Goldberg T, Strauss H (2008) Ferruginous Conditions dominated later neoproterozoic deep-water chemistry. Science 321:949–952 Canfield DE, Knoll AH, Poulton SW, Narbonne GM, Dunning GR (2020) Carbon isotopes in clastic rocks and the Neoproterozoic carbon cycle. Am J Sci 320:97–124 Cawood PA (2005) Terra Australis Orogen: rodinia breakup and development of the Pacific and Iapetus margins of Gondwana during the Neoproterozoic and Paleozoic. Earth Sci Rev 69:249–279 Cecioni G (1981) Precambrian pebbly mudstones in Eritrea, northeastern Ethiopia. In: Hambrey MJ, Harland WB (eds) Earth’s prepleistocene glacial record. Cambridge University Press, A24, 150 Cohen KM, Finney SC, Gibbard PL, Fan J-X (2013) updated 2015. The ICS international chronostratigraphic chart. Episodes 36:199– 204 Cohen PA, Strauss JV, Rooney AD, Sharma M, Tosca N (2017) Controlled hydroxyapatite biomineralization in an *810 millionyear-old unicellular eukaryote. Sci Adv 3: Cole JC (1988) Geology of the Aban Al Ahmar Quadrangle, Sheet 25F, Kingdom of Saudi Arabia (explanatory notes). Deputy Ministry for Mineral Resources Map GM-105A, C Collins AS, Pisarevsky SA (2005) Amalgamating eastern Gondwana: the evolution of the circum-Indian Orogens. Earth-Sci Rev 71:229– 270 Collins AS, Windley BF (2002) The tectonic evolution of central and northern Madagascar and its place in the final assembly of Gondwana. J Geol 110:325–339 Condon D, Zhu M, Bowring S, Wang W, Yang A, Jin Y (2005) U-Pb ages from the Neoproterozoic Doushantuo Formation, China. Science 308:95–98 Cosca MA, Shimron A, Caby R (1999) Late Precambrian metamorphism and cooling in the Arabian-Nubian Shield: petrology and 40 Ar/39Ar geochronology of metamorphic rocks of the Elat area (southern Israel). Precambr Res 98:107–127 Cox GM, Lewis CJ, Collins AS, Halverson GP, Jourdan F, Foden J, Nettle D, Kattan F (2012) Ediacaran terrane accretion within the Arabian-Nubian Shield. Gondwana Res 21:341–352 Cox GM, Halverson GP, Minarik WG, Le Heron DP, Macdonald FA, Bellefroid EJ, Strauss JV (2013) Neoproterozoic iron formation: an evaluation of its temporal, environmental and tectonic significance. Chem Geol 362:232–249 Cox GM, Halverson GP, Stevenson RK, Vokaty M, Poirier A, Kunzmann M, Li Z-X, Denyszyn SW, Strauss JV, Macdonald FA (2016) Continental flood basalt weathering as a trigger for Neoproterozoic Snowball Earth. Earth Planet Sci Lett 446:89–99 Cox GM, Foden J, Collins AS (2018) Late Neoproterozoic adakitic magmatism of the eastern Arabian Nubian Shield. Geosci Front 10:1981–1992 Cui H, Kaufman AJ, Zou H, Kattan FH, Trusler P, Smith J, Ivantsov A, Rich TH, Al Qubsani A, Yazedi A, Liu X-M, Johnson P, Goderis S, Claeys P, Vickers-Rich P (2020) Primary or secondary? A dichotomy of the strontium isotope anomalies in the Ediacaran carbonates of Saudi Arabia. Precambr Res 343: Dabbagh ME, Rogers JJW (1983) Depositional environments and tectonic significance of the Wajid Sandstone of southern Saudi Arabia. J Afr Earth Sci 1:47–57 Dalziel IWD (1997) Neoproterozoic-paleozoic geography and tectonics: review, hypothesis, environmental speculation. Geol Soc Am Bull 109:16–42

N. R. Miller and R. J. Stern Davies FB (1985) Explanatory notes to the geologic map of the Al Wajh Quadrangle, Kingdom of Saudi Arabia. Geoscience Map GM-83, scale 1:250,000, sheet 26B. Deputy Ministry for Mineral Resources, Ministry of Petroleum and Mineral Resources, Kingdom of Saudi Arabia, p 27 Davies J, Nairn AEM, Ressetar R (1980) The palaeomagnetism of certain late Precambrian and early Palaeozoic rocks from the Red Sea Hills, eastern desert, Egypt. J Geophys Res 85:3699–3710 De Souza Filho CR, Drury SA (1998) A Neoproterozoic supra-subduction terrane in northern Eritrea, NE Africa. J Geol Soc London 155:551–566 Delfour J (1970) Le Groupe de J’Balah, une nouvelle unite du Bouclier Arabe. Bureau de Recherche Geologique et Minieres Bulletin 4:19–32 Denèle Y, Leroy S, Pelleter E, Pik R, Talbot JY, Kahanbari K (2012) The Cryogenian juvenile Arc formation and successive high-K calc-alkaline plutons intrusion of Socotra Island Yemen. Arab J Geosci 5:903–924 Deynoux M, Affaton P, Trompette R, Villeneuve M (2006) Pan-African tectonic evolution and glacial events registered in Neoproterozoic to Cambrian cratonic and foreland basins of West Africa. J Afr Earth Sc 46:397–426 Dohrmann M, Wörheide G (2019) Dating early animal evolution using phylogenomic data. Nat Sci Reports 7:3599 Donnadieu Y, Goddéris Y, Ramstein G, Nédélec A, Meert J (2004) A ‘snowball Earth’ climate triggered by continental break-up through changes in runoff. Nature 428:303–306 Dor YB, Harlavan Y, Avigad D (2018) Provenance of the great Cambrian sandstone succession of northern Gondwana unraveled by strontium, neodymium and lead isotopes of feldspars and clays. Sedimentology 65:2595–2620 Drury SA, Berhe SM (1993) Accretion tectonics in northern Eritrea revealed by remotely sensed imagery. Geol Mag 130:177–190 El-Shazly AK, Khalil KI (2016) Metamorphic and geochronologic constraints on the tectonic evolution of the Central Eastern Desert of Egypt. Precambr Res 283:144–168 El-Shazly AK, Khalil KI, Helba HA (2019) Geochemistry of banded iron formations and their host rocks from the Central Eastern Desert of Egypt: a working genetic model and tectonic implications. Precambr Res 325:192–216 Ernst RE, Wingate MTD, Buchan KL, Li ZX (2008) Global record of 1600–700 Ma Large Igneous Provinces (LIPs): implications for the reconstruction of the proposed Nuna (Columbia) and Rodinia supercontinents. Precambr Res 160:59–178 Etemad-Saeed N, Hosseini-Barzi M, Adabi M, Miller NR, Sadeghi A, Stockli Houshmandzadeh DF (2016) Evidence for *560 Ma Ediacaran glaciation in the Kahar Formation, Central Alborz Mountains, northern Iran. Gondwana Res 31:164–183 Evans D (2000) Stratigraphic, geochronological, and paleomagneticconstraints upon the Neoproterozoic climatic paradox. Am J Sci 300:347–433 Evans D, Raub TD (2011) Neoproterozoic glacial palaeolatitudes: a global update. In: Arnaud E, Halverson GP, Shields-Shou G (eds) The geological record of Neoproterozoic glaciations. Geological Society of London, London, pp 93–112 Eyal M, Be’eri-Shlevin Y, Eyal Y, Whitehouse MJ, Litvinovsky B (2014) Three successive proterozoic island arcs in the northern arabian–nubian shield: evidence from SIMS U–Pb dating of zircon. Gondwana Res 25:338–357 Fairchild IJ, Spiro B, Herrington PM (2000) Controls on Sr and C isotope compositions of Neoproterozoic Sr-rich limestones of East Greenland and North China. In: Grotzinger JP, James NP (eds) Carbonate sedimentation and diagenesis in an evolving precambrian world, SEPM Special Publication vol 67, pp 297–313

7

Evolution of the Arabian Nubian Shield and Snowball Earth

Fairchild IJ, Spencer T, Ali D, Anderson R, Anderton R, Boomer I, Dove D, Evans J, Hambrey M, Howe J, Sawaki Y, Wang Z, Shields G, Zhou Y, Skelton A, Tucker M (2018) Tonian-Cryogenian boundary sections of Argyll, Scotland. Precambr Res 319:37–64 Feulner G (2012) The faint young Sun problem. Rev Geophys 50:364– 370 Fike DA, Grotzinger JP, Pratt LM, Summons RE (2006) Oxidation of the Ediacaran Ocean. Nature 444:744–747 Fitzsimons ICW (2000) Grenville-age basement provinces in east Antarctica: evidence for three separate collisional orogens. Geology 28:879–882 Fowler A, Hassen I, Hassan M (2015) Tectonic evolution and setting of the Sa’al Complex, southern Sinai, Egypt: a Proterozoic continental back-arc rift model. J Afr Earth Sci 104:103–131 Fritz H, Abdelsalam M, Ali KA, Bingen B, Collins AS, Fowler AR, Ghebreab W, Hauzenberger CA, Johnson PR, Kusky TM, Macey P, Muhongo S, Stern RJ, Viola G (2013) Orogen styles in the East African Orogen: a review of the Neoproterozoic to Cambrian tectonic evolution. J Afr Earth Sci 86:65–106 Gad S, Kusky T (2007) ASTER spectral ratioing for lithological mapping in the Arabian-Nubian shield, the Neoproterozoic Wadi Kid area, Sinai, Egypt. Gondwana Res 11:326–335 Garfunkel Z (1999) History and paleogeography during the Pan-African orogen to stable platform transition: reappraisal of the evidence from the Elat area and the northern Arabian-Nubian Shield. Israel J Earth Sci 48:135–157 Garland CR (1980) Geology of the Adigrat area. Ministry of Mines Memoir No. 1, p 51. Addis Ababa 1:250,000 map Gernon TM, Hincs TK, Tyrrell T, Rohling EJ, Palmer MR (2016) Snowball Earth ocean chemistry driven by extensive ridge volcanism during Rodinia breakup. Nat Geosci 9:242–248 Ghebreab W (1999) Tectono-metamorphic history of Neoproterozoic rocks in eastern Eritrea. Precambr Res 98:83–105 Glennie KW, Boeuf MGA, Hughes-Clarke MW, Moody-Stuart M, Pilaar WFH, Reinhardt BM (1974) Geology of the Oman mountains. Verhandelingen van het Koninklijk Nederlands geologisch mijnbouwkundig Genootschap 31:423 Goddéris Y, Donnadieu Y, Nédélec A, Dupré B, Dessert C, Grard A, Ramstein G, François LM (2003) The Sturtian “snowball” glaciation: fire and ice. Earth Planet Sci Lett 211:1–12 Gong Z, Kodama KP, Li YX (2017) Rock magnetic cyclostratigraphy of the Doushantuo Formation, South China and its implications for the duration of the Shuram carbon isotope excursion. Precambr Res 289:62–74 Gostin VA, McKirdy DM, Webster LJ, Williams GE (2010) Ediacaran ice-rafting and coeval asteroid impact, South Australia: insights into the terminal Proterozoic environment. Aust J Earth Sci 57:859–869 Gray DR, Foster DA, Meert JG, Goscombe BD, Armstrong R, Truow RAJ, Passchier CW (2008) A Damaran perspective on the assembly of southwestern Gondwana. Geol Soc London, Special Publications 294:257–278 Grotzinger JP, Fike DA, Fischer WW (2011) Enigmatic origin of the largest-known carbon isotope excursion in Earth’s history. Nat Geosci 4:285–292 Guilbaud R, Poulton SW, Butterfield NJ, Zhu M, Shields-Zhou GA (2015) A global transition to ferruginous conditions in the early Neoproterozoic oceans. Nat Geosci 8:466–470 Hadley DG (1974) The taphrogeosynclinal Jubaylah Group in the Mashhad area, northwestern Hijaz. Saudi Arabian Directorate General of Mineral Resources Bulletin 10, p 18 Hadley DG (1986) Explanatory notes to the geologic map of the Sahl Al Matran Quadrangle, Kingdom of Saudi Arabia. Geoscience Map GM-86 C, scale 1:250,000, sheet 26C. Deputy Ministry for Mineral Resources, Ministry of Petroleum and Mineral Resources, Kingdom of Saudi Arabia. p 24

189 Hailu T (1975) Geological map of Adi Arkay. Adis Ababa, Geological Survey of Ethiopia Technical Report, scale 1:250,000 Halverson GP (2006) A Neoproterozoic chronology. In: Xiao S, Kaufman AJ (eds) Neoproterozoic geobiology and paleobiology: dordrecht. Springer, Netherlands, pp 231–271 Halverson GP, Hoffman PF, Schrage DP, Maloof AC, Rice AHN (2005) Toward a Neoproterozoic composite carbon-isotope record. Geol Soc Am Bull 117:1181–1207 Halverson GP, Dudás FÖ, Maloof AC, Bowring SA (2007a) Evolution of the 87Sr/86Sr composition of Neoproterozoic seawater. Palaeogeogr Palaeoclimatol Palaeoecol 256:103–129 Halverson GP, Maloof AC, Schrag DP, Dudás FÖ, Hurtgen M (2007b) Stratigraphy and geochemistry of a ca 800 Ma negative carbon isotope interval in northeastern Svalbard. Chem Geol 237:5–27 Halverson GP, Cox GM, Théou-Hubert L, Schmitz M, Hagadorn JW, Johnson P, Sansjofre P, Kunzmann M, Schumann D (2013) A multi-proxy geochemical record from a late Neoproterozoic volcano-sedimentary basin, eastern Arabian Shield. McGill University, Canada, unpublished poster Halverson GP, Porter SM, Gibson TM (2018) Dating the late Proterozoic stratigraphic record. Emerg Top Life Sci 2:137–147 Hargrove US (2006a) Crustal evolution of the Neoproterozoic Bi’r Umq suture zone, Kingdom of Saudi Arabia. Geochronological, Isotopic, and Geochemical Constraints. Ph.D. Thesis, University of Texas, Dallas, TX, USA Hargrove US, Stern RJ, Kimura J-I, Manton WI, Johnson PR (2006) How juvenile if the Arabian-Nubian Shield? Evidence from Nd isotopes and pre-Neoproterozoic inherited zircon in the Bi’r Umq suture zone, Saudi Arabia. Earth Planet Sci Lett 252:308– 326 Harland WB (1964) Evidence of late Precambrian glaciation and its significance. In: Nairn AEM (ed) Problems in palaeoclimatology. Interscience, London, pp 119–149 Hassanlouei BT, Rajabzadeh MA (2019) Iron ore deposits associated with Hormuz evaporitic series in Hormuz and Pohl salt diapirs, Hormuzgan province, southern Iran. J Asian Earth Sci 172:30–55 Hebert CL, Kaufman AJ, Penniston-Dorland SC, Martin AJ (2010) Radiometric and stratigraphic constraints on terminal Ediacaran (post-Gaskiers) glaciation and metazoan evolution. Precambr Res 182:402–412 Hedge CE (1984) Precambrian geochronology of part of northwestern Saudi Arabia, Kingdom of Saudi Arabia. US Geological Survey Open File Report 83–381, p 12 Helmy HM, Morad AE, Abdel Rahman HB (2021) Um Zariq formation, southeast sinai, Egypt: a new record of the Sturtian Snowball Earth event in the Arabian Nubian Shield. J Afr Earth Sc 173: Hill AC, Arouri K, Gorjan P, Walter MR (2000) Geochemistry of marine and nonmarine environments of a Neoproterozoic cratonic carbonate/evaporite: the Bitter Springs Formation, Central Australia. In: Grotzinger JP, James NP (eds) Carbonate sedimentation and diagenesis in the evolving precambrian world. SEPM, 327–344 (Spec. Pub. 67) Hoffman PF (1999) The break-up of Rodinia, birth of Gondwana, true polar wander and the snowball earth. J Afr Earth Sci 29:17–33 Hoffman PF (2013) The great oxidation and a Siderian snowball Earth: MIF-S based correlation of Paleoproterozoic glacial epochs. Chem Geol 362:143–156 Hoffman PF, Schrag DP (2002) The snowball Earth hypothesis: testing the limits of global change. Terra Nova 14:129–155 Hoffman PF, Kaufman AJ, Halverson GP, Schrag DP (1998) A Neoproterozoic snowball earth. Science 281:1342–1346 Hoffman PF, Halverson GP, Domack EW, Husson JM, Higgins JA, Schrag DP (2007) Are basal Ediacaran (635 Ma) post-glacial “cap dolostones” diachronous? Earth Planet. Sci Lett 258:114–131

190 Hoffman PF, Abbot DS, Ashkenazy Y, Benn DI, Brocks JJ, Cohen PA, Cox GM, Creveling JR, Donnadieu Y, Erwin DH, Fairchild IJ, Ferreira D, Goodman JC, Halverson GP, Jansen MF, Le Hir G, Love GD, Macdonald FA, Maloof AC, Partin CA, Ramstein G, Rose BEJ, Sadler PM, Tziperman E, Voigt A, Warren SG (2017) Snowball earth climate dynamics and cryogenian geologygeobiology. Sci Adv 3:e1600983 Horton F (2015) Did phosphorus derived from the weathering of large igneous provinces fertilize the Neoproterozoic ocean? Geochem Geophys Geosystems 1:1723–1738 Husson JM, Maloof AC, Schoene B, Chen CY, Higgins JA (2015) Stratigraphic expression of Earth’s deepest d13C excursion in the Wonoka Formation of South Australia. Am J Sci 315:1–45 Hyde WT, Crowley TJ, Baum SK, Peltier WR (2000) Neoproterozoic ‘Snowball Earth’ simulations with a coupled climate/icesheet model. Nature 405:425–429 Jacobs J, Thomas RJ (2004) Himalayan-type indenter-escape tectonics model for the southern part of the late Neoproterozoic-early Paleozoic East African-Antarctic orogen. Geology 32:721–724 Jacobs J, Bauer W, Fanning CM (2003) Late Neoproterozoic/Early Palaeozoic events in central Dronning Maud Land and significance for the southern extension of the East African Orogen into East Antarctica. Precambr Res 126:27–53 Jarrar GH, Wachendorf H, Zellmer D (1991) The Saramuj Conglomerate: evolution of a Pan-African molasse sequence from southwest Jordan. N Jb Geol Palaontol Mh 6:335–356 Jarrar GH, Wachendorf H, Zachmann D (1993) A Pan-African alkaline pluton intruding the Saramuj Conglomerate, southwest Jordan. Geol Rundschau 82:121–135 Javoy M, Pineau F, Delorme H (1986) Carbon and nitrogen isotopes in the mantle. Chem Geol 57:41–62 Johnson PR (1995) Proterozoic geology of Western Saudi Arabia— North-Central Sheet: explanatory notes on precambrian stratigraphic relations. Saudi Arabian Deputy Ministry for Mineral Resources, Jiddah, Saudi Arabia, USGS-OF-95-5, pp 1–44 Johnson PR (2003) Post-amalgamation basins of the NE Arabian shield and implications for Neoproterozoic III tectonism in the northern East African Orogen. Precambr Res 123:321–337 Johnson PR (2014) An expanding Arabian-Nubian Shield geochronologic and isotopic dataset: defining limits and confirming the tectonic setting of a Neoproterozoic accretionary orogen. Open Geol J 8:3–33 Johnson P, Kattan F (2008) Lithostratigraphic revision in the Arabian Shield: the impacts of geochronology and tectonic analysis. Arab J Sci Eng 33:3–16 Johnson PR, Kattan FH (2012) The geology of the Saudi Arabian Shield. Saudi Geological Survey, Jiddah, Saudi Arabia, pp 1–479 Johnson PR, Stewart ICF (1995) Magnetically inferred basement structure in central Saudi Arabia. Tectonophysics 245:37–52 Johnson PR, Woldehaimanot B (2003) Development of the Arabian-Nubian Shield: perspectives on accretion and deformation in the northern East African Orogen and the assembly of Gondwana. In: Yoshida M, Windley BF, Dasgupta S (eds) Proterozoic East Gondwana: supercontinent assembly and breakup, Geological Society, London, Spec. Pub. 206, pp 289–325 Johnson PR, Abdelsalam MG, Stern RJ (2003) The Bi’r Umq-Nakasib Suture Zone in the Arabian-Nubian Shield: a key to understanding crustal growth in the East African Orogen. Gondwana Res 6:523– 530 Johnson PR, Andresen A, Collins AS, Fowler AR, Fritz H, Ghebreab W, Kusky T, Stern RJ (2011) Late cryogenian-ediacaran history of the Arabian-Nubian Shield: a review of depositional, plutonic, structural, and tectonic events in the closing stages of the northern East African Orogen. J Afr Earth Sci 61:167–232

N. R. Miller and R. J. Stern Johnson PR, Halverson GP, Kusky TM, Stern RJ, Pease V (2013) Volcanosedimentary Basins in the Arabian-Nubian Shield: markers of repeated exhumation and denudation in a neoproterozoic accretionary orogen. Geosciences 3:389–445 Jones DS, Maloof AC, Hurtgen MT, Rainbird RH, Schrag DP (2010) Regional and global chemostratigraphic correlation of the early Neoproterozoic Shaler Supergroup, Victoria Island, Northwestern Canada. Precambr Res 181:43–63 Kaufman AJ, Jacobsen SB, Knoll AH (1993) TheVendian record of Sr and C isotopic variations in seawater: implications for tectonics and paleoclimate. Earth Planet Sci Lett 84:27–41 Kaufman AJ, Knoll AH, Narbonne GM (1997) Isotopes, ice ages, and terminal Proterozoic earth history. Proc Natl Acad Sci USA 94:6600–6605 Keller CB, Husson JM, Mitchell RN, Bottke WF, Gernon TM, Boehnke P, Bell EA, Swanson-Hysell NL, Peters SE (2020) Neoproterozoic glacial origin of the Great Unconformity. Proc Nat Academy Sci 116:1136–1145 Kemp J (1981) Geologic map of the Wadi Al Ays Quadrangle, Sheet 25C, Kingdom of Saudi Arabia, Saudi Arabian Deputy Ministry for Mineral Resources: Jiddah, Saudi Arabia, GM 53, pp 1–39 Kemp J, Gros Y, Prian JP (1982) Geologic Map of the Mahd adh Dhahab Quadrangle, Sheet 23E, Kingdom of Saudi Arabia, Saudi Arabian Deputy Ministry for Mineral Resources: Jiddah, Saudi Arabia, GM 64, pp 1–39 Kempf O, Kellerhals P, Lowrie W, Matter A (2000) Palaeomagnetic directions in late Precambrian glaciomarine sediments of the Mirbat Sandstone Formation, Oman. Earth Planet Sci Lett 175:181–190 Kennedy A, Johnson PR, Kattan FH (2004) SHRIMP geochronology in the Northern Arabian Shield Part I: data acquisition; Saudi geological survey: Jiddah, Saudi Arabia, SGS-OF-2004-11, pp 1–28 Kennedy A, Johnson PR, Kattan FH (2005) SHRIMP geochronology in the Northern Arabian Shield Part II: data acquisition 2004, Saudi Geological Survey: Jiddah, Saudi Arabia, SGS-OF-2005–10, pp 1–44 Kennedy A, Kozdroj W, Kattan FH, Ziolkowska-Kozdroj M, Johnson PR (2010a) SHRIMP Geochronology in the Arabian Shield (Midyan Terrane, Afif Terrane, Ad Dawadi Terrane) and Nubian Shield (Central Eastern Desert Terrane) Part IV: data acquisition 2008; Saudi Geological Survey: Jiddah, Saudi Arabia; SGS-OF2010-10, pp 1–101 Kennedy A, Kozdroj W, Kadi K, Ziolkowska-Kozdroj M, Johnson PR (2010b) SHRIMP geochronology of the Arabian Shield (Midyan Terrane, Afif Terrane) and Nubian Shield (Central Eastern Desert Terrane), Part V: data acquisition 2009; Saudi Geological Survey, Jiddah, Saudi Arabia, SGS-OF-2010-11, p 80 Kidane T, Bachtadse V, Alene M (2014) Quaternary remagnetization of the Neoproterozoic limestone of Negash Synclinorium (Arabian– Nubian Shield, northern Ethiopia): With implications of no paleomagnetic testing for the proposed Snowball Earth events. Phys Earth Planet Inter 235:1–12 Kilner B, Conall MN, Brasier M (2005) Low-latitude glaciation in the Neoproterozoic of Oman. Geology 33:413–416 Kirschvink JL (1992) Late Proterozoic low-latitude global glaciation: the snowball Earth. In: Schopf JW, Klein C (eds) The Proterozoic biosphere—a multidisciplinary study. Cambridge University Press, Cambridge, pp 51–52 Kiyokawa S, Suzuki T, El-Dokouny HA, Dawoud M, Abuelhasan AA (2020) Stratigraphy, petrology, and geochemistry of a Neoproterozoic banded iron sequence in the El-Dabbah Group, central Eastern Desert, Egypt. J Afr Earth Sc 168: Klaebe RM, Smith MP, Fairchild IJ, Fleming EJ, Kennedy MJ (2018) Facies-dependent d13C variation and diagenetic overprinting at the onset of the Sturtian glaciation in north-east Greenland. Precambr Res 319:96–113

7

Evolution of the Arabian Nubian Shield and Snowball Earth

Knoll AH, Walter MR, Narbonne GM, Christie-Blick N (2004) The Ediacaran Period: a new addition to the geologic time scale. Lethaia 39:13–30 Knoll AH, Walter MR, Narbonne GM, Christie-Blick N (2006) A new period the geologic time scale. Science 205:621–622 Kröner A (2001) The mozambique belt of East Africa and madagascar: significance of zircon and Nd model ages for Rodinia and Gondwana supercontinent formation and dispersal. S Afr J Geol 104:151–166 Kröner A, Muhongo S, Hegner E, Wingate MTD (2003) Single zircon geochronology and Nd isotopic systematics of Proterozoic highgrade rocks from the Mozambique belt of southern Tanzania (Masasi area): implications for Gondwana assembly. J Geol Soc London 160:645–757 Kusky TM, Matsah MI (2003) Neoproterozoic dextral faulting on the Najd fault system, Saudi Arabia, preceded sinistral faulting and escape tectonics related to closure of the Mozambique Ocean. Geol Soc Spec Publ 206:327–361 Küster D, Liégeois J-P, Matukov D, Sergeev S, Lucassen F (2008) Zircon geochronology and Sr, Nd, Pb isotope geochemistry of granitoids from Bayuda Desert and Sabaloka (Sudan): evidence for a Bayudian event (920–900 Ma) preceding the Pan-African orogenic cycle (860–590 Ma) at the eastern boundary of the Saharan Metacraton. Precambr Res 164:16–39 Kuznetsov AB, Semikhatov MA, Maslov AV, Gorokhov IM, Prasolov EM, Krupenin MT, Kislova IV (2006) New data on Sr-and C-isotopic chemostratigraphy of the Upper Riphean type section (Southern Urals). Stratigr Geol Correl 14:602–628 Le Guerroué E (2010) Duration and synchroneity of the largest negative carbon isotope excursion on Earth: the Shuram/Wonoka anomaly. CR Geosci 342:204–214 Letalenet J (1979) Explanatory notes to the geologic map of the ‘Afif Quadrangle, Kingdom of Saudi Arabia. Geoscience Map GM-47C, scale 1:250,000, sheet 23F. Deputy Ministry for Mineral Resources, Ministry of Petroleum and Mineral Resources, Kingdom of Saudi Arabia, p 20 Li ZX, Bogdanova SV, Collins AS, Davidson A et al (2008) Assembly configuration and break-up history of Rodinia: a synthesis. Precambr Res 160:179–210 Li Z-X, Evans DAD, Halverson GP (2013) Neoproterozoic glaciations in a revised global palaeogeography from the breakup of Rodinia to the assembly of Gondwanaland. Sed Geol 294:219–232 Li X-H, Abd El-Rahman Y, Abu Anbar M, Li J, Ling X-X, Wu L-G, Masoud AE (2018) Old continental crust underlying juvenile oceanic arc: evidence from northern Arabian-Nubian Shield, Egypt. Geophys Res Lett 45:3001–3008 Linneman U, Pidal AP, Hofmann M, Drost K, Quesada C, Gerdes A, Marko L, Gärtner A, Zieger J, Ulrich J, Krause R, Vickers-Rich P, Horak J (2018) A *565 Ma old glaciation in the Ediacaran of peri-Gondwanan West Africa. Int J Earth Sci https://doi.org/10. 1007/s00531-017-1520-7 Liu Y, Peltier WR (2013) Sea level variations during snowball Earth formation: 1. A preliminary analysis. J Geophys Res 118:4410– 4434 Love GD, Grosjean E, Stalvies C, Fike DA, Grotzinger JP, Bradley AS, Kelly AE, Bhatia M, Bowring SA, Condon DJ, Summons RE (2009) Fossil steroids record the appearance of demospongiae during the cryogenian. Nature 457:718–722 Macdonald FA, Schmitz MD, Crowley JL, Roots CF, Jones DS, Maloof AC, Strauss JV, Cohen PA, Johnston DT, Schrag DP (2010a) Calibrating the cryogenian. Science 327:1241–1243 Macdonald FA, Strauss JV, Rose CV, Dudas F, Schrag DP (2010b) Stratigraphy of the port Nolloth Group of Namibia and South Africa and implications for the age of neoproterozoic iron formations. Am J Sci 310:862–888

191 MacLennan SA, Park Y, Swanson-Hysell NL, Maloof AC, Schoene B, Gebreslassie M, Antilla E, Tesema T, Alene M, Haileab B (2018) The arc of the Snowball: U–Pb dates constrain the Islay anomaly and the initiation of the Sturtian glaciation. Geology 46:539–542 Maloof AC, Schrag DP, Crowley JL, Bowring SA (2005) An expanded record of early Cambrian carbon cycling from the Anti-Atlas Margin, Morocco. Can J Earth Sci 42:2195–2216 Maloof AC, Halverson GP, Kirschvink JL, Schrag DP, Weiss BP, Hoffman PF (2006) Combined paleomagnetic, isotopic, and stratigraphic evidence for true polar wander from the Neoproterozoic Akademikerbreen Group, Svalbard, Norway. GSA Bulletin 118:1024–1099 Master S, Wendorff M (2011) Neoproterozoic glaciogenic diamictites of the Katanga Supergroup, Central Africa. In: Arnaud E et al (eds) The geological record of neoproterozoic glaciations. Geol Soc London Memoir 36:173–183 McCaffrey MA, Moldowan JM, Lipton PA, Summons RE, Peters KE, Jeganathan A, Watt DS (1994) Paleoenvironmental implications of novel C30 steranes in Precambrian to Cenozoic age petroleum and bitumen. Geochimica et Cosmochimica Acta 58:529–532 Meert J (2003) A synopsis of events related to the assembly of eastern Gondwana. Tectonophysics 362:1–40 Meert JG, Lieberman BS (2008) The Neoproterozoic assembly of Gondwana and its relationship to the Ediacaran-Cambrian radiation. Gondwana Res 14:5–21 Meert JG, Van Der Voo R (1997) The assembly of Gondwana 800–550 Ma. J Geodyn 23:223–235 Miller NR, Alene M, Sacchi R, Stern R, Conti A, Kröner A, Zuppi G (2003) Significance of the Tambien Group (Tigre, N. Ethiopia) for snowball Earth events in the Arabian-Nubian Shield. Precambr Res 121:263–283 Miller N, Schilman B, Avigad D, Stern RJ, Beyth M (2006) Neoproterozoic Snowball Earth—the northern Ethiopia record. Abstract presented in the “Snowball Earth 2006” conference, Monte Verita, Switzerland, pp 74–75 Miller N, Johnson P, Stern R (2008) Marine versus non-marine environments for the Jibalah Group, NW Arabian Shield: a sedimentological and geochemical survey and report of possible Metazoa in the Dhaiqa Formation. Arab J Sci Eng 22:55–77 Miller NR, Stern RJ, Avigad D, Beyth M, Schilman B (2009) Neoproterozoic carbonate-slate sequences of the Tambien Group, N. Ethiopia (I): pre-‘Sturtian’ chemostratigraphy and regional correlation. Precambr Res 170:129–156 Miller NR, Avigad D, Stern RJ, Beyth M (2011) The Tambien Group, northern Ethiopia (Tigre). In: Arnaud E et al (eds) The geological record of Neoproterozoic glaciations. Geol Soc London Memoir 36:263–276 Moghazi A-KM, Ali KA, Wilde SA, Zhou Q, Andersen T, Andresen A, El-Enen MMA, Stern RJ (2012) Geochemistry, geochronology, and Sr–Nd isotopes of the Late Neoproterozoic Wadi Kid volcano-sedimentary rocks, Southern Sinai, Egypt: implications for tectonic setting and crustal evolution. Lithos 154:147–165 Morag N, Avigad D, Gerdes A, Belousova E, Harlavan Y (2011a) Crustal evolution and recycling in the northern Arabian-Nubian Shield: new perspectives from zircon Lu–Hf and U–Pb systematics. Precambr Res 186:101–116 Morag N, Avigad D, Gerdes A, Belousova E, Harlavan Y (2011b) Detrital zircon Hf isotopic composition indicates long-distance transport of North Gondwana Cambrian-Ordovician sandstones. Geology 39:955–958 Nairn AEM, Perry TA, Ressetar R, Rogers S (1987) A palaeomagnetic study of the Dokhan volcanic formation and younger granites, eastern desert of Egypt. J Afr Earth Sc 6:353–365 Nettle D (2009) A sequence stratigraphic, geochronological and chemostratigraphic investigation of the Ediacaran Antaq Basin,

192 Eastern Arabian Shield, Saudi Arabia. Unpublished Honours Thesis, Geology and Geophysics, University of Adelaide, Australia; Saudi Geological Survey Technical Report SGSTR-2010-5, p 83 Nettle D, Halverson GP, Cox GM, Collins AS, Schmitz M, Gehling J, Johnson PR, Kadi K (2013) A middle-late Ediacaran volcano-sedimentary record from the Eastern Arabian-Nubian Shield. Terra Nova 26:120–129 Newman MJ, Rood RT (1977) Implications of solar evolution for the Earth’s early atmosphere. Science 198:1035–1037 Nicholson PG, Janjou DLA, Fanning CM, Heaman LM, Grotzinger JP (2008) Deposition, age, and Pan-Arabian correlation of late Neoproterozoic outcrops in Saudi Arabia (abstract), 8th Middle East Geoscience Conference and exhibition, GEO 2008, Manama, Bahrain AAPG Search and Discovery Magazine, article 90077 Och LM, Shields-Zhou GA (2012) The Neoproterozoic oxygenation event: environmental perturbations and biogeochemical cycling. Earth Sci Rev 110:26–57 Pallister JS, Stacey JS, Fischer LB, Premo WR (1988) Precambrian ophiolites of Arabia: geologic settings, U–Pb geochronology, Pb-isotope characteristics, and implications for continental accretion. Precambr Res 38:1–54 Park U, Swanson-Hysell NL, MacLennan SA, Maloof AC, Gebreslassie M, Tremblay MM, Schoene B, Alene M, Anttila ESC, Tesema T, Haileab B (2019) The lead-up to the Sturtian Snowball Earth: neoprotoerozoic chemostratigraphy time-calibrated by the Tambien Group of Ethiopia. GSA Bulletin 132:1119–1149 Patchett PJ, Chase CG (2002) Role of transform continental margins in major crustal growth episodes. Geology 30:39–42 Pellaton C (1979) Geologic map of the Yanbu’al Bahr Quadrangle, Sheet 24C, Kingdom of Saudi Arabia. Saudi Arabian Directorate General of Mineral Resources, Jiddah, Saudi Arabia Pisarevsky SA, Murphy JB, Cawood PA, Collins AS (2008) Late Neoproterozoic and early Cambrian palaeogeography: models and problems. Geol Soc, London 294:9–31 Special Publications Pisarevsky SA, McCausland PJ, Hodych JP, O’Brien SJ, Tait JA, Murphy JB, Colpron M (2011) Paleomagnetic study of the late Neoproterozoic Bull Arm and Crown Hill formations (Musgravetown Group) of eastern Newfoundland: Implications for Avalonia and West Gondwana paleogeography. Can J Earth Sci 49:308–327 Planavsky NJ, Reinhard CT, Wang X, Thomson D, McGoldrick P, Rainbird RH, Johnson T, Fischer WW, Lyons TW (2014) Low mid-Proterozoic atmospheric oxygen levels and the delayed rise of animals. Science 346:635–638 Pollard D, Kasting JF (2005) Snowball Earth: a thin-ice solution with flowing sea glaciers. J Geophys Res 110:C07010 Poulton SW, Canfield DE (2011) Ferruginous conditions: a dominant feature of the ocean through Earth’s history. Elements 7:107–112 Powell JH, Abed A, Jarrar GH (2015) Ediacaran araba complex of jordan. GeoArabia 20:99–156 Pu JP, Bowring SA, Ramezani J, Myrow P, Raub TD, Landing E, Mills A, Hodgin E, Macdonald FA (2016) Dodging snowballs: geochronology of the Gaskiers glaciation and the first appearance of the Ediacaran biota. Geology 44:955–958 Ramsay CR (1986) Geologic Map of the Rabigh Quadrangle, Sheet 22D, Kingdom of Saudi Arabia. Saudi Arabian Deputy Ministry for Mineral Resources: Jiddah, Saudi Arabia, GM 84:1–49 Rieu R, Allen PA, Cozzi A, Kosler J, Bussy F (2007) A composite stratigraphy for the Neoproterozoic Huqf Supergroup of Oman: integrating new litho-, chemo- and chronolstratigraphic data of the Mirbat area, southern Oman. J Geol Soc, London 164:997–1009 Rooney AD, Chew DM, Selbey D (2011) Re–Os geochronology of the Neoproterozoic-Cambrian Dalradian Supergroup of Scotland and Ireland: Implications for Neoproterozoic stratigraphy, glaciations and Re–Os systematics. Precambr Res 185:202–214

N. R. Miller and R. J. Stern Rooney AD, Macdonald FA, Strauss JV, Dudás FÖ, Hallmann C, Selby D (2014) Re-Os geochronology and coupled Os-Sr isotope constraints on the Sturtian snowball Earth. Proc Natl Acad Sci 111:51–56 Rooney AD, Strauss JV, Brandon AD, Macdonald FA (2015) A Cryogenian chronology: two long-lasting, synchronous Neoproterozoic snowball Earth glaciations. Geology 43:459–462 Rooney AD, Cantine MD, Bergmann KD, Gómez-Pérez I, Al Baloushi B, Boag TH, Busch JF, Sperling EA, Strauss JV (2020) Calibrating the coevolution of Ediacaran life and environment. Proc Nat Acad Sci 117:16824–16830 Rothman DH, Hayes JM, Summons RE (2003) Dynamics of the Neoproterozoic carbon cycle. Proc Natl Acad Sci 100:8124–8129 Runnegar B (2000) Loophole for snowball earth. Nature 405:403–404 Sahoo SK, Planavsky NJ, Kendall B, Wang X, Shi X, Scott C, Anbar AD, Lyons TW, Jiang G (2012) Ocean oxygenation in the wake of the Marinoan glaciation. Nature 489:546–549 Sandler A, Teutsch N, Avigad D (2012) Sub-Cambrian pedogenesis recorded in weathering profiles of the Arabian-Nubian Shield. Sedimentology 59:1305–1320 Sawaki Y, Kawai T, Shibuya T, Tahata M, Omori S, Tsuyoshi K, Yoshida N, Hirata T, Ohno T, Windley BF, Shigenori M (2010) 87 Sr/86Sr chemostratigraphy of Neoproterozoic Dalradian carbonates below the Port Askaig Glaciogenic Formation, Scotland. Precambr Res 179:150–164 Schrag DP, Higgins JA, Macdonald FA, Johnston DT (2013) Authigenic carbonate and the history of the global carbon cycle. Science 339:540–543 Shang CK, Morteani G, Satir M, Taubald H (2010) Neoproterozoic continental growth prior to Gondwana assembly: constraints from zircon-titanite geochronology, geochemistry and petrography of ring complex granitoids, Sudan. Lithos 118:61–81 Shields GA, Mills BJW, Zhu M, Raub RD, Daines SJ, Lenton TM (2019) Unique Neoproterozoic carbon isotope excursions sustained by coupled evaporite dissolution and pyrite burial. Nat Geosci 12:823–827 Shields-Zhou GA, Denoux M, Och L (2011) The record of Neoproterozoic glaciation in the Taoudéni Basin, NW Africa. In: Arnaud E et al (eds) The geological record of Neoproterozoic glaciations. Geol Soc London Memoir 36:163–171 Shields-Zhou GA, Porter S, Halverson GP (2016) A new rock-based definition for the Cryogenian Period (circa 720–635 Ma). Episodes 39:3–8 Sifeta K, Roser BP, Kimura JI (2005) Geochemistry, provenance, and tectonic setting of Neoproterozoic metavolcanic and metasedimentary units, Werri area, Northern Ethiopia. J Afr Earth Sc 41:212–223 Spence GH, Le Heron DP, Fairchild IJ (2016) Sedimentological perspectives on climatic, atmospheric and environmental change in the Neoproterozoic Era. Sedimentology 63:253–306 Squire RJ, Campbell IH, Allen CM, Wilson CJ (2006) Did the Transgondwanan Supermountain trigger the explosive radiation of animals on Earth? Earth Planet. Sci Lett 250:116–133 Stein M (2003) Tracing the plume material in the Arabian-Nubian Shield. Precambr Res 123:223–234 Stern RJ (1985) The najd fault system, Saudi Arabia and Egypt: a late Precambrian rift-related transform system? Tectonics 4:497–511 Stern RJ (1994) Arc-assembly and continental collision in the Neoproterozoic African Orogen: implications for the consolidation of Gondwanaland. Annu Rev Earth Planet Sci 22:319–351 Stern RJ (2002) Crustal evolution in the East African Orogen: a neodymium isotopic perspective. J Afr Earth Sci 34:109–117 Stern RJ (2018) The evolution of plate tectonics. Philosophical Transactions A, 376, 20170406. https://doi.org/10.1098/rsta.2017. 0406

7

Evolution of the Arabian Nubian Shield and Snowball Earth

Stern RJ, Abdelsalam MG (1998) Formation of juvenile continental crust in the Arabian-Nubian Shield: evidence from granitic rocks of the Nakasib suture. NE Sudan Geol Rundschau 87:150–160 Stern RJ, Johnson P (2010) Continental lithosphere of the Arabian Plate: a geologic, petrologic, and geophysical synthesis. Earth Sci Rev 101:29–67 Stern RJ, Miller NR (2018) Did the transition to plate tectonics cause Neoproterozoic snowball Earth? Terra Nova 30:87–94 Stern RJ, Miller NR (2019) Neoproterozoic glaciation—snowball earth hypothesis. Encyclopedia of geology, 2nd ed. https://doi.org/10. 1016/B978-0-12-409548-9.12107-4 Stern RJ, Avigad D, Miller NR, Beyth M (2006) Geological society of Africa presidential review: evidence for the snowball earth hypothesis in the Arabian-Nubian Shield and the East African Orogen. J Afr Earth Sc 44:1–20 Stern RJ, Johnson PR, Ali KA, Mukherjee S (2011) Evidence for early and mid-cryogenian glaciation in the northern Arabian-Nubian Shield (Egypt, Sudan, and western Arabia). In: The geological record of neoproterozoic glaciation. In: Arnaud E, Halverson GP, Shield-Zhou G (eds) Geol Soc London, UK, 2011; Memoirs 36, 277–284 Stewart SA (2016) Structural geology of the Rub’ Al-Khali Basin, Saudi Arabia. Tectonics 35:2417–2438 Stoeser DB, Stacey JS (1988) Evolution, U–Pb geochronology, and isotope geology of the Pan-African Nabitah orogenic belt of the Saudi Arabian Shield. In: El-Gaby S, Greiling RO (eds) The Pan-African Belt of NE Africa and adjacent areas, Friedr. Vieweg & Sohn, Braunschweig, pp 227–289 Sultan M, Arvidson RE, Duncan I, Stern RJ, El Kaliouby M (1988) Extension of the Najd Shear System from Saudi Arabia to the Central Eastern Desert of Egypt based on Integrated Field and Landsat Observations. Tectonics 7:1291–1306 Swanson-Hysell NL, Rose CV, Calmet CC, Halverson GP, Hurtgen MT, Maloof AC (2010) Cryogenian glaciation and the onset of carbon-isotope decoupling. Science 328:608–611 Swanson-Hysell NL, Maloof AC, Condon DJ, Jenkin GR, Alene M, Tremblay MM, Tesema T, Rooney AD, Haileab B (2015) Stratigraphy and geochronology of the Tambien Group, Ethiopia: evidence for globally synchronous carbon isotope change in the Neoproterozoic. Geology 43:323–326 Tadesse T (1997) The geology of axum area (ND 37-6). Ethiopian Institute of Geological Surveys, Addis Ababa (Memoir No. 9) Tadesse T (1999) Axum sheet geological map. Geological Survey of Ethiopia, Addis Ababa, Ethiopia. 1:250,000 map Tadesse T, Hoshino M, Suzuki K, Iizumi S (2000) Nd, Rb–Sr and Th– U–Pb zircon ages of syn- and post-tectonic granitoids from the Axum area of northern Ethiopia. J Afr Earth Sci 30:313–327 Tait J, Delpomdor F, Préat A, Tack A, Straathof G, Nkula VK (2011) Neoproterozoic sequences of the West Congo and Lindi/Ubangi Supergroups in the Congo Craton, Central Africa. In: Arnaud E et al (eds) The geological record of Neoproterozoic glaciations. Geol Soc London Memoir 36:185–193 Teklay M, Kröner A, Mezger K (2001) Geochemistry, geochronology and isotope geology of Nakfa intrusive rocks, northern Eritrea: products of a tectonically thickened Neoproterozoic arc crust. J Afr Earth Sci 33:283–301 Tostevin R, Clarkson MO, Gangl S, Shields GA, Wood RA, Bowyer F, Penny AM, Stirling CH (2019) Uranium isotope evidence for an expansion of anoxia in terminal Ediacaran oceans. Earth Planet Sci Lett 506:104–112 Trindade RIF, Macouin M (2007) Palaeolatitude of glacial deposits and palaeogeography of Neoproterozoic ice ages. CR Geosci 339:200–211 Tschopp RH (1967) The general geology of Oman. In: Proceedings of the 7th world petroleum congress 2, pp 231–242

193 Tsige L, Abdelsalam MG (2005) Neoproterozoic-early Paleozoic gravitational tectonic collapse in the southern part of the Arabian-Nubian Shield: the Bulbul Belt of southern Ethiopia. Precambr Res 38:31–297 Tziperman E, Halevy I, Johnston DT, Knoll AH, Schrag DP (2011) Biologically induced initiation of Neoproterozoic snowball-Earth events. Proc Natl Acad Sci USA 108:15091–15096 Vail JR (1983) Pan-African crustal accretion in northeast Africa. J Afr Earth Sci 1:285–294 Veevers JJ (2003) Pan-African is Pan-Gondwanaland: oblique convergence drives rotation during 650–500 Ma assembly. Geology 31:501–504 Vernhet R, Youbi N, Chellai EH, Villeneuve M El, Archi A (2012) The Bou-Azzer glaciation: evidence for an Ediacaran glaciation on the West African Craton (Anti-Atlas, Morocco). Precambr Res 196– 197:106–112 Verri P (1909) Contributo allo studio geografio Della Colonia Eritrea. Bollettino della Societa’ Geografica Italiana, 10, pp 251–320. Carta Geologica 1:1,500,000 Vickers-Rich P, Kozdroj W, Kattan FH, Leonov M, Ivantsov A, Johnson PR, Linnemann U, Hofmann M, Al Garni SM, Al Qubsani A, Shamari A, Al Baraki A, Al Kaff MH, Ziolkowska-Kozdroj M, Rich TH, Trusler P, Rich B (2010) Reconnaissance for an Ediacaran fauna, Kingdom of Saudi Arabia. Saudi Geological Survey, Technical Report, SGS-TR-2010-8, p 42, 74 figs., 1 table, 1 pl Vickers-Rich P, Ivantsov A, Kattan FH, Johnson PR, Al Qubsani A, Kashghari W, Leonov M, Rich R, Linnemann U, Hofmann M, Trusler P, Smith J, Yazidi A, Rich B, Al Garni SM, Shamari A, Al Barakati A, Al Kaff MH (2013) In search of the Kingdom’s Ediacarans: the first genuine Metazoans (macroscopic body and trace fossils) from the Neoproterozoic Jibalah Group (Vendian/Ediacaran) on the Arabian Shield. Saudi Geological Survey Technical Report SGS-TR-2013-5, p 21., 19 figs., 1 table Weissbrod T, Sneh A (2002) Sedimentology and paleogeography of the Late Precambrian–Early Cambrian arkosic and conglomeratic facies in the northern margins of the Arabo-Nubian Shield Bulletin. Geol Survey Israel 87:44 Whitehouse MJ, Pease V, Al-Khirbash S (2016) Neoproterozoic crustal growth at the margin of the East Gondwana continent—age and isotopic constraints from the easternmost inliers of Oman. Int Geol Rev 58:2046–2064 Wilde SA, Youssef K (2000) Significance of SHRIMP dating of the imperial porphyry and associated Dokhan volcanics, Gebel Dokhan, northeastern desert, Egypt. J Afr Earth Sci 31:403–413 Wilde SA, Youssef K (2002) A re-evaluation of the origin and setting of the late Precambrian Hammamat Group based on SHRIMP U–Pb dating of detrital zircons from Gebel Umm Tawat, North Eastern Desert, Egypt. J Geol Soc London 159:595–604 Williams GE (2008) Proterozoic (pre-Ediacaran) glaciation and the highobliquity, low-latitude ice, strong seasonality (HOLIST) hypothesis: principles and tests. Earth Sci Rev 87:61–93 Williams GE, Schmidt PW, Young GM (2016) Strongly seasonal Proterozoic glacial climate in low palaeolatitudes: radically different climate system on the pre-Ediacaran Earth. Geosci Front 7:555–571 Wu C, Yang T, Shields GA, Bian X, Bao B, Ye H, Li W (2020) Termination of Cryogenian ironstone deposition by deep ocean euxinia. Geochem Perspect Lett 15:1–5 Yin Z, Zhu M, Davidson EH, Bottjer DJ, Zhao F, Tafforeau P (2015) Sponge grade body fossil with cellular resolution dating 60 Myr before the Cambrian. Proc Nat Acad Sci 141:E1453–E1460 Zhou MT, Luo T, Huff WD, Yang Z, Zhou G, Gan T, Yang H, Zhange D (2018) Timing the termination of the Doushantuo negative carbon isotope excursion: evidence from U–Pb ages from

194 the Dengying and Liuchapo formations, South China. Sci Bull 63:1431–1438 Zhou Y, Pogge von Strandmann PAE, Zhu M, Ling H, Manning C, Li D, He T, Shields GA (2020) Reconstructing Tonian seawater 87 Sr/86Sr using calcite microspar. Geology 48:462–467

N. R. Miller and R. J. Stern Zimmer M, Kröner A, Jochum KP, Reischmann R, Todt W (1995) The Gabal Gerf complex: a Precambrian N-MORB ophiolite in the Nubian Shield, NE Africa. Chem Geol 123:29–51

8

The Arabian–Nubian Shield Within the Neoproterozoic Plate Tectonic Circuit Alan S. Collins, Morgan L. Blades, and Andrew S. Merdith

Abstract

8.1

The limited shortening and crustal thickening in the Arabian–Nubian Shield (ANS) during central Gondwana amalgamation have preserved the pre-continental collision accretionary orogen in a manner that allows us to reconstruct past plate tectonic kinematics. We interpret this to show remarkable changes in plate tectonic motions at ca. 720 Ma that reflect a major change in subduction of the Mozambique Ocean. This is represented in the ANS by the transition from earlier (presently oriented) NE–SW trending suture zones (Bi’r Umq–Nakasib and Yanbu–Sol Hamed) to younger approximately NNW–SSE trending sutures (Keraf and Nabitah). This plate reorganization event is seen elsewhere in the planet at this time and is interpreted to represent the beginning of Neoproterozoic India’s southward progression from Tonian high latitudes to more equatorial locations as it advanced orthogonally towards the Neoproterozoic African continents on one side and obliquely against Western Australia on the other. The geology of the ANS provides vital constraints in the endeavour of reconstructing the plate tectonic circuit of the globe in deep time. Keywords

Neoproterozoic reconstructions








Plate tectonics Full-plate Arabian-Nubian Shield

A. S. Collins (&)
M. L. Blades Tectonics and Earth Systems (TES), Department of Earth Sciences, The University of Adelaide, Adelaide, SA 5005, Australia e-mail: [email protected] A. S. Merdith UnivLyon, Université Lyon 1, Ens de Lyon, CNRS, UMR 5276 LGL-TPE, Villeurbanne, F-69622, France

Introduction

It is uncontroversial to point out that earth surface systems such as the hydrosphere, atmosphere and biosphere are largely controlled by element transfer from the deep earth through plate tectonic-driven processes such as mountain building (and their subsequence weathering) and volcanicity. For the later Phanerozoic, full-plate tectonic models of the evolving globe have been essential in understanding the feedbacks and thresholds that govern how our planet modulates its climate, hydrosphere and atmosphere circulation, and ecology (Ramstein et al. 2019; Zhang et al. 2018). However, in deep time, especially before the oldest existing ocean crust, it is very hard to model the perturbations on the earth surface that govern ocean and atmosphere circulation. To get towards these holistic earth system models, we first need well-constrained full-plate tectonic models of the earth (e.g. Domeier 2018; Merdith et al. 2017a, 2021; Muller et al. 2019; Young et al. 2019). Unlike palaeomagnetic-solutiongoverned continent reconstructions (e.g. Li et al. 2008), the main constraints on full-plate models are geological—the evidence of past plate margins incorporated in the rock record. The Arabian–Nubian Shield (ANS) preserves the evidence for plate interactions between some of the major plates in the Neoproterozoic over at least 400 million years of earth history. Here, we discuss how interpreting the geology of the ANS is critical in reconstructing the globe through this key time of earth system development. The Arabian–Nubian Shield (ANS) is one of the most extensive preserved areas of newly formed Proterozoic crust on Earth (Stern 2002), and it likely forms the continuation of earlier accretion systems hidden beneath the Sahara (Blades et al. 2021; de Wit and Linol 2015; Şengör et al. 2020) that likely decorated a kernel of Archaean/Palaeoproterozoic craton (Abdelsalam et al. 2002; Liégeois et al. 2013; Sobh et al. 2020). Its preservation is exceptional and is largely due to the region lying peripheral to the main Neoproterozoic India/Congo continent–continent collision that formed the

© The Author(s), under exclusive license to Springer Nature Switzerland AG 2021 Z. Hamimi et al. (eds.), The Geology of the Arabian-Nubian Shield, Regional Geology Reviews, https://doi.org/10.1007/978-3-030-72995-0_8

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central part of Gondwana (Collins and Pisarevsky 2005; Merdith et al. 2021). Consequently, it retains much of the tectonic architecture of the pre-Gondwana forming accretional orogens that mantled both sides of the Mozambique Ocean during the late Mesoproterozoic and Neoproterozoic (at least from ca. 1080–600 Ma). To the south, the ANS passes into the Mozambique Belt (Fritz et al. 2013)—a region of considerably greater crustal thickening. This region is interpreted to be the main locus of continent–continent collision where the earlier accretionary history is harder to reconstruct due to the grade of metamorphism of the exhumed rocks. Together, the ANS and the Mozambique Belt form the East African Orogen (Fritz et al. 2013; Johnson et al. 2011; Stern 1994).

8.2

A Full-Plate Tectonic Reconstruction

There are two broad categories of models that can be constructed to describe Earth’s tectonic or palaeogeographic history. The first category are ‘continental drift’ type models that reconstruct the motion of continents drifting across the Earth’s surface. The second type are ‘full-plate models that, in addition to tracking the motion of continents, trace the evolution of plate boundaries and by implication, the evolution of tectonic plates themselves (Gurnis et al. 2012). Full-plate models are more encompassing than continental drift models. However, they are also much harder to modify iteratively. Full-plate models have a notable benefit over continental drift models as they allow the quantification of various plate tectonic controls on the earth system. These include: the plate tectonic forces causing the global distribution of continents; the assessment of plate kinematics; the rates of subduction and mid-ocean ridge formation; the length of plate margins through time; and the bathymetry of the ocean basins (e.g. Collins et al. 2021; Merdith et al. 2019; Williams et al. 2021). Neoproterozoic full-plate models were introduced by Merdith et al. (2017a) and have recently been extensively updated and combined with Phanerozoic reconstructions to produce a continuous 1 Ga to present full-plate reconstruction (Merdith et al. 2021). Here, we discuss the ANS within this model and emphasize how the geology of the ANS is vital in constraining these models.

8.3

The ANS Within the Northern ‘East African Orogen’

The ANS is laced with suture zones that represent collisions between different terranes as subduction zones consumed the intervening oceanic crust (Fig. 8.1). A dramatic feature of the region is that pre-715 Ma sutures are aligned approximately 90° from post-715 Ma sutures (Johnson et al. 2011;

Robinson et al. 2015). Collins et al. (2021) recently suggested that this suture zone strike change reflects a major change in plate convergence direction, and we use this as the start of a higher-order reconstruction of this region in a full-plate context.

8.4

The Mozambique Ocean, Azania and Afif–Abas

The Mozambique Ocean closed as Neoproterozoic India converged on the African parts of Gondwana (Kalahari, Congo, Sahara) to form central Gondwana (Meert 2003; Schmitt et al. 2018). The East African Orogen resulted from the collision between these major continents and amalgams of smaller terranes, during the Neoproterozoic to early Cambrian. Sandwiched within the EAO lies a broad band of Archaean to Palaeoproterozoic crust that was identified by Collins and Windley (2002) as a microcontinent (subsequently named ‘Azania’), whose remains are found in southern India, central Madagascar, Somalia, eastern Ethiopia and Arabia (Fig. 8.1). In Yemen, the Al-Mafid Terrane is correlated with Azania (Collins and Windley 2002) and this is separated from a second pre-Neoproterozoic terrane called the Abas Terrane by a Neoproterozoic arc terrane (the Al Bayda terrane). Because of this, Collins and Windley (2002) suggested that a second microcontinent existed that they called Afif–Abas due to the continuation of the Abas terrane into Saudi Arabia as the Afif terrane. Azania and Afif–Abas are interpreted to have collided with the eastern margins of the Congo craton and Saharan Metacraton by approximately 630 Ma to form the East African Orogeny sensu stricto (Stern 2002). A younger orogeny (ca. 570–520 Ma) was interpreted to represent the final collision between India and the amalgamated Africa/Arabia and called the Malagasy orogeny by Collins and Pisarevsky (2005).

8.5

The Eastern Margin of the EAO (NW India to Oman)

The easternmost margin of the northern East African Orogen is the boundary between the Mesoproterozoic terranes of India and the Stenian–Tonian crust that extends west from the Delhi–Aravalli Orogen. This has been interpreted to be the eastern margin of the northern East African Orogen. During the Stenian and Tonian, progressive arc accretion of volcanic arc rocks onto the NW margin of Neoproterozoic India occurred, extending into the basement rocks of Pakistan and the inliers of Oman (Alessio et al. 2018; Blades et al. 2019a). This was later covered by an extensive Cryogenian–Ediacaran passive margin succession, with

8

The Arabian–Nubian Shield Within the Neoproterozoic Plate …

197

Fig. 8.1 Map of present northern Indian Ocean region with distribution of juvenile Stenian–Ediacaran crust, pre-Stenian exposed crust, pre-Stenian exposed crust reworked thermally and structurally during the Neoproterozoic and Proterozoic sedimentary basins in NE Africa, Arabia and the Indian subcontinent. The late Tonian plate reconfiguration is represented by the notable change from pre-715 Ma, approximately NE–SW sutures to post-715 Ma, approximately NNW–SSE striking sutures. MB Mozambique Belt, NED Northern Eastern Desert, CED Central Eastern Desert, SED South Eastern Desert

comparable sequences continuing into the Cambrian (Cozzi et al. 2012).

8.6

The Arabian–Nubian Shield (ANS)

The ANS is dominated by low-grade volcano-sedimentary sequences and associated plutonic and ophiolitic remnants. The tectonic history of the ANS is complicated and preserves a complex mix of terranes, accreted arcs that record subduction polarity reversals that are reviewed and summarized in a number of papers (Johnson et al. 2011; Robinson et al. 2015). There are no reliable palaeomagnetic data available to constrain these blocks, so we have constrained their positions by their relation to each other and through plate kinematic constraints. The oldest terrane in the ANS is the late Mesoproterozoic Sa’al Metamorphic complex (1.03–1.02 Ga) in Sinai, marking the initiation of magmatism in the northernmost ANS (Fig. 8.2, Be’eri-Shlevin et al. 2012; Eyal et al. 2014). The location of this Stenian terrane was uncertain, but coeval subduction magmatism occurred within the region of the Saharan Metacraton (see below).

The Tonian to Cryogenian history of the ANS is marked by the formation of oceanic volcanic arcs and arcs built on Azanian (or Afif–Abas) crust that amalgamated to form a larger intra-Mozambique ocean terrane separate from both Neoproterozoic India and African Gondwanan continents. Several terranes in the ANS are correlated as equivalents, separated by the opening of the Red Sea, from south to north; these are the Asir and Tokar/Barka terranes, the Haya and Jiddah terranes, the Hijaz and Gabgaba/Gebeit terranes, and the Eastern Desert and Midyan terranes (Johnson et al. 2011) (Fig. 8.2). It is unclear whether the combined Asir-Tokar/Barka terrane and Haya–Jiddah terranes were ever on separate plates as, in Saudi Arabia, no clear suture is seen between them. In SE Sudan and Eritrea, the Barka suture does appear as the site of ocean closure, so these may form a complex middle Tonian amalgam. The older, Tonian to earliest Cryogenian, amalgamation history of the ANS is marked by approximately NE–SW oriented sutures (present orientation; Fig. 8.2) between juvenile Neoproterozoic ocean-arc terranes. The oldest of these sutures is between the Jiddah–Haya and Gabgaba/ Gebeit–Hijaz terranes (the Bi’r Umq–Nakasib suture), which is dated at ca. 780–750 Ma (Johnson et al. 2011). This suture

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Fig. 8.2 Four hemispherical reconstructions modified from the full-plate tectonic reconstruction of Merdith et al. (2021), focused on the region of the future Arabian–Nubian Shield (ANS). At 780 Ma, the earlier (now NE–SW) sutures are subduction zones that, as the intervening oceans are closed, suture to form the ANS as a separate continent within the Mozambique Ocean. In the subsequent 700 Ma reconstruction, the later (now NNW–SSE striking) Keraf and Nabitah sutures are subduction zones that facilitate the closure of the ocean separating the now amalgamated ANS with the Sahara Metacraton and the southward movement of India and closure of the Mozambique Ocean. A–A Afif–Abas, Ar Ar Rayn terrane, Aus Australia, Av Avalonia, Az Azania, Co Congo/Tanzania/Bangweulu continent, East Rodinia combined continent with Australia/Mawson and North China Craton, KA Kalahari Craton, L Lhasa, NAC North Australian Craton, nANS northern ANS, NC North China, NI Neoproterozoic India, SAC combined South Australian and West Australian Craton, Sah Sahara Metacraton, sANS southern ANS, SC South China, Si Sinai, T Tarim. Subduction zones marked as red lines ornamented with red triangles, divergent and transform margins marked as black lines

created the kernel of a late Tonian microcontinent. The Midyan–Eastern Desert collided with this kernel ca. 715 Ma along the Yanbu-Sol Hamed suture (Robinson et al. 2015) (Fig. 8.2). Both of these sutures evolved from SE-dipping subduction zones (Robinson et al. 2015). The older NE–SW sutures are bound by younger NNW– SSE Cryogenian to Ediacaran sutures and terranes (Fig. 8.1) that represent a fundamental kinematic change in Mozambique Ocean subduction. The oldest of these is the 680– 640 Ma Nabitah suture, which forms the eastern margin of the intra-Mozambique Ocean island-arc terrane microcontinent (discussed above), against Tonian–Cryogenian continental arcs built on the Afif–Abas microcontinent. This now

enlarged Afif–Abas microcontinent collided with the active margin of the Sahara Metacraton along the Sudanese Keraf Suture (Abdelsalam et al. 1998; Blades et al. 2015). This collision occurred in late Cryogenian to early Ediacaran times (ca. 650–580 Ma) (Abdelsalam et al. 1998). Further to the east, in the most easterly exposed terrane, the Saudi Ar Rayn Terrane, juvenile calc-alkaline magmatism, stretches from ca. 690 Ma to 615 Ma (Doebrich et al. 2007). Turbiditic sediment deposition in the Ad Dawadimi basin that separates the Ar Rayn Terrane from the Afif–Abas microcontinent continued until at least 620 Ma, but was locally intruded by ca. 630 Ma adakitic magmas (Cox et al. 2019). This sequence was metamorphosed to greenschist facies

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grades at ca. 620 Ma (Cox et al. 2012). Further east still, broad N–S magnetic highs, beneath the Arabian Phanerozoic sedimentary sequence (Johnson and Stewart 1995), suggest younger arc terranes now buried beneath the Rub al-Khali Basin. The transition to post-tectonic magmatism within the eastern terranes of the ANS begins at ca. 605 Ma (Doebrich et al. 2007) and pull-apart basins developed along the large strike-slip faults that cut the region (Nettle et al. 2014). The final collision between Neoproterozoic India and the, by then amalgamated, Azania/Congo Craton occurred at ca. 570–540 Ma, closing the final strand of the Mozambique Ocean (Collins and Pisarevsky 2005; Schmitt et al. 2018). This suture lies beneath the Phanerozoic cover between the exposed Saudi and Yemen basement and Mirbat in SW Oman. It appears to be imaged by shear wave anisotropy variations seen directly west of Mirbat (Al-Lazki et al. 2012). In western Oman, latest Ediacaran–Cambrian deformation occurs in the subsurface, its limit is known as the Western Deformation Front and the deformation associated with this is known as the ‘Angudan event’ (Loosveld et al. 1996). The sub-Rub al-Khali suture has been traced south within reconstructed Gondwana to Madagascar where it has been correlated with the Antsaba shear zone of NW Madagascar (Armistead et al. 2019), the Betsimisaraka suture of Collins and Windley (2002) and into the Palghat–Cauvery Suture of southern India (Collins et al. 2007). This Palghat– Betsimisaraka–Antsaba–Western Deformation Front suture represents the final suture of the Mozambique Ocean (Collins and Pisarevsky 2005; Merdith et al. 2017a).

that the Sahara Metacraton did not exist as a coherent ancient entity, but instead formed a Neoproterozoic series of accretionary terranes that were folded into great oroclinal bends during Neoproterozoic orogenesis.

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8.7

The Western Margin of the EAO (the Eastern Saharan Metacraton)

The Saharan Metacraton is still very poorly known, but extensive late Mesoproterozoic subduction-related magmatism is found in Chad and east and north Sudan (Blades et al. 2021; de Wit and Linol 2015; Isseini et al. 2012; Shellnutt et al. 2020). To the west of this, in eastern Sudan and western Ethiopia magmatism associated with early Neoproterozoic subduction characterizes terranes that are thought to have formed over westward dipping subduction zones (Blades et al. 2015, 2019b). This longevity of subduction, which also includes that seen in the Sinai (Be’eri-Shlevin et al. 2012; Eyal et al. 2014), demonstrates that the EAO extends back into the Stenian, or even earlier, when terrane accretion and subduction zone magmatism initiated against the Palaeoproterozoic kernel of the ‘metacraton’. The NE margin of the Congo Craton, in Uganda, preserves orogenesis that begins with Tonian subduction zone magmatism in the Karamoja Belt that is coeval with terranes in Sudan (Westerhof et al. 2014). Recently, Şengör et al. (2020) linked all of these together to support previous assertions

8.8

Discussion

The variations in geometry of preserved plate tectonic boundaries within the ANS are reflected by the change in orientation of the main subduction magmatic belts and intervening suture zones. These allow the tectonic geography of the region to be back calculated, and these are then able to be used to infer changes in ancient plate kinematics in a full-plate tectonic model. The approximately 90° change in suture zone orientation at approximately 715 Ma suggests a major plate kinematic reorganization at this time in this area (Collins et al. 2021). In the Phanerozoic earth, plate tectonics is characterized by long periods of semi-continuous plate motions separated by short periods of plate kinematic reconfigurations due to changes in the coupling between the lithosphere and the deep earth. Since Gondwana broke up, there have been two main plate reorganization events, at ca. 140–120 Ma and at ca. 50 Ma (Muller et al. 2016; Whittaker et al. 2007). Full-plate reconstructions allow us to investigate plate reorganization events in deep time. The ca. 715 Ma reorganization of plate kinematics in the ANS occurs at approximately the same time as Neoproterozoic India began to move south from the high latitudes it inhabited in the Tonian (Torsvik et al. 2001), to begin a journey that would bring it colliding into central Africa and sliding against Western Australia by the Cambrian (Collins and Pisarevsky 2005; Powell and Pisarevsky 2002). Elsewhere in the world, central Rodinia was breaking apart and ca. 715 Ma is plausibly the time of ocean crust production in the Pacific Ocean as Laurentia broke away from Australia/Mawson (Merdith et al. 2017b). This is recorded as a time of extensive rifting in the Adelaide Superbasin (Lloyd et al. 2020). Therefore, the plate motion change seen in the ANS may well reflect global plate reorganization event at approximately the time of one of the largest climate (Hoffman et al. 2017), seawater chemistry (Cox et al. 2016) and biological perturbations (Brocks et al. 2017) as the planet entered the Cryogenian. Gernon et al. (2016) suggested that increased, shallow, mid-ocean ridge production may have caused for this planetary pivot. Other proposed plate tectonic triggers include increased magmatism during breakup of Rodinia (Cox et al. 2016; Goddéris et al. 2003), decreased continental arc magmatism reducing CO2 influx into the atmosphere during the Cryogenian (McKenzie et al. 2016), and even a suggestion that modern-style plate tectonics may have started at this time and caused major perturbations to earth

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surface systems (Stern and Miller 2018). Testing of these models is now possible through these full-plate tectonic reconstructions that are based primarily on geological information from areas such as the ANS.

8.9

Conclusions

The ANS preserves evidence for a major change in plate kinematics at ca. 715 Ma that are demonstrated by the change in trend of volcanic arc terranes and intervening suture zones. These are broadly NE–SW before ca. 715 Ma and *NNW–SSE after ca. 715 Ma. By placing these geological observations into a full-plate tectonic framework, it is possible to reconstruct the subduction history of this part of the Mozambique Ocean and correlate this plate kinematic change with the southern progression of Neoproterozoic India (that likely included South China and many central Asian terranes) from Tonian high latitudes to more equatorial locations when it collided with the Congo craton and Australia/Mawson by the Cambrian. This ca. 715 Ma plate tectonic event is tentatively correlated with other plate tectonic changes elsewhere in the world and may represent a global plate reorganization event that occurred close to the Tonian–Cryogenian boundary. Acknowledgements ASC was supported by Australian Research Council grants FT120100340 and with MLB and JDF is funded by LP160101353, with the support of the Northern Territory Geological Survey, Origin Energy, Santos Ltd and Imperial Oil and Gas. ASM is supported by the Deep Energy Community of the Deep Carbon Observatory. The authors thank Jean-Paul Liégeois for a constructive review that improved the manuscript.

References Abdelsalam MG, Stern RJ, Copeland P, El-Faki E-FM, Al-Hur B, Ibrahim FM (1998) The Neoproterozoic Keraf Suture in NE Sudan: sinistral transpression along the eastern margin of West Gondwana: J Geol 106:133–147 Abdelsalam MG, Liegeois JP, Stern RJ (2002) The saharan metacraton. J Afr Earth Sc 34(3–4):119–136 Alessio BL, Blades ML, Murray G, Thorpe B, Collins AS, Kelsey DE, Foden J, Payne J, Al-Khirbash S, Jourdan F (2018) Origin and tectonic evolution of the NE basement of Oman: a window into the Neoproterozoic accretionary growth of India?: Geol Magaz 155 (5):1150–1174 Al-Lazki A, Ebinger C, Kendall M, Helffrich G, Leroy S, Tiberi C, Stuart G, Al-Toobi K (2012) Upper mantle anisotropy of southeast Arabia passive margin [Gulf of Aden northern conjugate margin]. Oman: Arabian J Geosci 5(5):925–934 Armistead SE, Collins AS, Merdith AS, Payne JL, Cox GM, Foden JD, Razakamanana T, De Waele B (2019) Evolving marginal terranes during neoproterozoic supercontinent reorganization: constraints from the bemarivo domain in Northern Madagascar. Tectonics 38 (6):2019–2035

Be’eri-Shlevin Y, Eyal M, Eyal Y, Whitehouse MJ, Litvinovsky B, (2012) The Sa’al volcano-sedimentary complex (Sinai, Egypt): a latest Mesoproterozoic volcanic arc in the northern Arabian Nubian Shield: Geology 40(5):403–406 Blades ML, Collins AS, Payne JL, Foden JD, Stüwe K, Hassan M, Abu-Alam, T, Makroum F (2021) Age and Hafnium isotope evolution of Sudanese Butana and Chad illuminates the Stenian to Ediacaran evolution of the southeast Sahara Metacraton. Precambrian Research, in press Blades ML, Collins AS, Foden J, Payne JL, Xu XC, Alemu T, Woldetinsae G, Clark C, Taylor RJM (2015) Age and hafnium isotopic evolution of the Didesa and Kemashi Domains, western Ethiopia: Precambr Res 270:267–284 Blades ML, Alessio BL, Collins AS, Foden J, Payne JL, Glorie S, Holden P, Thorpe B, Al-Khirbash S (2019a) Unravelling the Neoproterozoic accretionary history of Oman, using an array of isotopic systems in zircon: J Geol Soc 2018–2125 Blades ML, Foden J, Collins AS, Alemu T, Woldetinsae G (2019b) The origin of the ultramafic rocks of the Tulu Dimtu Belt, western Ethiopia—do they represent remnants of the Mozambique Ocean?: Geol Magaz 156(1):62–82 Brocks JJ, Jarrett AJM, Sirantoine E, Hallmann C, Hoshino Y, Liyanage T (2017) The rise of algae in Cryogenian oceans and the emergence of animals. Nature 548(7669):578– + Collins AS, Pisarevsky SA (2005) Amalgamating eastern Gondwana: the evolution of the Circum-Indian Orogens. Earth Sci Rev 71 (3):229–270 Collins A, Windley B (2002) The tectonic evolution of central and northern Madagascar and its place in the final assembly of Gondwana. J Geol 110(3):325–339 Collins AS, Clark C, Sajeev K, Santosh M, Kelsey DE, Hand M (2007) Passage through India: the Mozambique Ocean suture, high-pressure granulites and the Palghat-Cauvery shear zone system. Terra Nova 19(2):141–147 Collins AS, Blades ML, Merdith AS, Foden JD (2021) Closure of the Proterozoic Mozambique Ocean was instigated by a late Tonian plate reorganization event. Commun Earth Environ 2(1):75 Cox GM, Lewis CJ, Collins AS, Halverson GP, Jourdan F, Foden J, Nettle D, Kattan F (2012) Ediacaran terrane accretion within the Arabian-Nubian Shield. Gondwana Res 21(2–3):341–352 Cox GM, Halverson GP, Stevenson RK, Vokaty M, Poirier A, Kunzmann M, Li ZX, Denyszyn SW, Strauss JV, Macdonald FA (2016) Continental flood basalt weathering as a trigger for Neoproterozoic Snowball Earth. Earth Planet Sci Lett 446:89–99 Cox GM, Foden J, Collins AS (2019) Late Neoproterozoic adakitic magmatism of the eastern Arabian Nubian Shield. Geosci Front 10 (6):1981–1992 Cozzi A, Rea G, Craig J (2012) From global geology to hydrocarbon exploration: Ediacaran-Early Cambrian petroleum plays of India, Pakistan and Oman. Geol Soc London, Special Publications 366 (1):131–162 de Wit MJ, Linol B (2015) Precambrian basement of the congo basin and its flanking terrains. In: de Wit MJ, Guillocheau F, de Wit MCJ (eds) Geology and resource potential of the congo basin. Springer, Berlin Heidelberg, pp 19–37 Doebrich JL, Al-Jehani AM, Siddiqui AA, Hayes TS, Wooden JL, Johnson PR (2007) Geology and metalogeny of the Ar Rayn terrane, eastern Arabian shield: Evolution of a Neoproterozoic continental-margin arc during assembly of Gondwana within the East African orogen. Precambr Res 158:17–50 Domeier M (2018) Early Paleozoic tectonics of Asia: towards a full-plate model. Geosci Front 9(3):789–862 Eyal M, Be’eri-Shlevin Y, Eyal Y, Whitehouse MJ, Litvinovsky B (2014) Three successive Proterozoic island arcs in the Northern

The Arabian–Nubian Shield Within the Neoproterozoic Plate …

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Arabian-Nubian Shield: evidence from SIMS U-Pb dating of zircon. Gondwana Res 25(1):338–357 Fritz H, Abdelsalam M, Ali KA, Bingen B, Collins AS, Fowler AR, Ghebreab W, Hauzenberger CA, Johnson P, Kusky T, Macey P, Muhongo S, Stern RJ, Viola G (2013) Orogen styles in the East Africa Orogens: a review of the Neoproterozoic to Cambrian Tectonic Evolution. J Afr Earth Sci 86:65–106 Gernon TM, Hincks TK, Tyrrell T, Rohling EJ, Palmer MR (2016) Snowball Earth ocean chemistry driven by extensive ridge volcanism during Rodinia breakup. Nat Geosci 9(3):242–283 Goddéris Y, Donnadieu Y, Nédélec A, Dupré B, Dessert C, Grard A, Ramstein G, François LM (2003) The Sturtian ‘snowball’ glaciation: fire and ice. Earth Planet Sci Lett 211:1–12 Gurnis M, Turner M, Zahirovic S, DiCaprio L, Spasojevic S, Muller RD, Boyden J, Seton M, Manea VC, Bower DJ (2012) Plate tectonic reconstructions with continuously closing plates. Comput Geosci 38:35–42 Hoffman PF, Abbot DS, Ashkenazy Y, Benn DI, Brocks JJ, Cohen PA, Cox GM, Creveling JR, Donnadieu Y, Erwin DH, Fairchild IJ, Ferreira D, Goodman JC, Halverson GP, Jansen MF, Le Hir G, Love GD, Macdonald FA, Maloof AC, Partin CA, Ramstein G, Rose BEJ, Rose CV, Sadler PM, Tziperman E, Voigt A, Warren SG (2017) Snowball earth climate dynamics and Cryogenian geology-geobiology. Sci Adv 3(11):e1600983 Isseini M, André-Mayer A-S, Vanderhaeghe O, Barbey P, Deloule E, (2012) A-type granites from the Pan-African orogenic belt in south-western Chadconstrained using geochemistry, Sr–Nd isotopes and U–Pb geochronology. Lithos 153:39–52 Johnson PR, Stewart ICF (1995) Magnetically inferred basement structure in central Saudi Arabia. Tectonophyscics 245:37–52 Johnson PR, Andresen A, Collins AS, Fowler AR, Fritz H, Ghebreab W, Kusky T, Stern RJ (2011) Late Cryogenian-Ediacaran history of the Arabian-Nubian Shield: a review of depositional, plutonic, structural, and tectonic events in the closing stages of the northern East African Orogen. J Afr Earth Sci 61(3):167–232 Li ZX, Bogdanova SV, Collins AS, Davidson A, De Waele B, Ernst RE, Fitzsimons ICW, Fuck RA, Gladkochub DP, Jacobs J, Karlstrom KE, Lu S, Natapov LM, Pease V, Pisarevsky SA, Thrane K, Vernikovsky V (2008) Assembly, configuration, and break-up history of Rodinia: a synthesis. Precambr Res 160(1– 2):179-210 Liégeois J-P, Abdelsalam G, Ennih N, Ouabadi A (2013) Metacraton: nature, genesis and behavior. Gondwana Res 23:220–237 Lloyd JC, Blades ML, Counts JW, Collins AS, Amos KJ, Wade BP, Hall JW, Hore S, Ball AL, Shahin S, Drabsch M (2020) Neoproterozoic geochronology and provenance of the Adelaide Superbasin. Precambr Res 350 Loosveld R, Bell A, Terken J (1996) The tectonic evolution of Interior Oman. Geoarabia 1:28–51 McKenzie NR, Horton BK, Loomis SE, Stockli DF, Planavsky NJ, Lee CTA (2016) Continental arc volcanism as the principal driver of icehouse-greenhouse variability. Science 352(6284):444–447 Meert JG (2003) A synopsis of events related to the assembly of eastern Gondwana. Tectonophysics 362(1–4):1–40 Merdith AS, Collins AS, Williams SE, Pisarevsky S, Foden JD, Archibald DB, Blades ML, Alessio BL, Armistead S, Plavsa D, Clark C, Muller RD (2017a) A full-plate global reconstruction of the Neoproterozoic. Gondwana Res 50:84–134 Merdith AS, Williams SE, Muller RD, Collins AS (2017b) Kinematic constraints on the Rodinia to Gondwana transition. Precambr Res 299:132–150 Merdith AS, Williams SE, Brune S, Collins AS, Muller RD (2019) Rift and plate boundary evolution across two supercontinent cycles. Global Planet Change 173:1–14

Merdith AS, Williams SE, Collins AS, Tetley MG, Mulder JA, Blades ML, Young A, Armistead SE, Cannon J, Zahirovic S, Müller RD (2021) Extending full-plate tectonic models into deep time: linking the Neoproterozoic and the Phanerozoic. Earth-Sci Rev 214:103477 Muller RD, Seton M, Zahirovic S, Williams SE, Matthews KJ, Wright NM, Shephard GE, Maloney KT, Barnett-Moore N, Hosseinpour M, Bower DJ, Cannon J (2016) Ocean basin evolution and global-scale plate reorganization events since pangea breakup. Annu Rev Earth Planet Sci 44:107–138 Muller RD, Zahirovic S, Williams SE, Cannon J, Seton M, Bower DJ, Tetley MG, Heine C, Le Breton E, Liu SF, Russell SHJ, Yang T, Leonard J, Gurnis M (2019) A global plate model including lithospheric deformation along major rifts and orogens since the triassic. Tectonics 38(6):1884–1907 Nettle D, Halverson GP, Cox GM, Collins AS, Schmitz M, Gehling J, Johnson PR, Kadi K (2014) A middle-late Ediacaran volcano-sedimentary record from the eastern Arabian-Nubian shield. Terra Nova 26(2):120–129 Powell CM, Pisarevsky SA (2002) Late Neoproterozoic assembly of East Gondwana. Geology 30:3–6 Ramstein G, Godderis Y, Donnadieu Y, Sepulchre P, Fluteau F, Zhang Z, Zhang R, Su B, Jiang D, Schuster M, Besse J (2019) Some illustrations of large tectonically driven climate changes in earth history. Tectonics 38(12):4454–4464 Robinson FA, Foden JD, Collins AS (2015) Geochemical and isotopic constraints on island arc, synorogenic, post-orogenic and anorogenic granitoids in the Arabian Shield, Saudi Arabia. Lithos 220:97–115 Schmitt RDS, Fragoso RDA, Collins AS (2018) Suturing Gondwana in the Cambrian: the orogenic events of the final amalgamation. In: Siegesmund S, Basei MAS, Oyhantçabal P, Oriolo S (eds) Geology of southwest Gondwana, pp 411–432. Springer International Publishing, Cham Şengör AMC, Lom N, Zabcı C, Sunal G, Öner T (2020) Reconstructing orogens without biostratigraphy: the Saharides and continental growth during the final assembly of Gondwana-Land. Proc Natl Acad Sci 117(51):32278–32284 Shellnutt JG, Pham NHT, Yeh M-W, Lee T-Y (2020) Two series of Ediacaran collision-related granites in the Guéra Massif, South-Central Chad: tectonomagmatic constraints on the terminal collision of the eastern Central African Orogenic Belt. Precambr Res 347: Sobh M, Ebbing J, Mansi AH, Gotze HJ, Emry EL, Abdelsalam MG (2020) The lithospheric structure of the saharan metacraton from 3-D integrated geophysical-petrological modeling. J Geophys Res-Solid Earth 125(8) Stern RJ (1994) Arc Assembly and continental collision in the Neoproterozoic East African orogeny—implications for the consolidation of Gondwana. Annu Rev Earth Planet Sci 22:319–351 Stern RA (2002) Crustal evolution in the East African Orogen: a neodymian isotopic perspective. J Afr Earth Sci 34:109–117 Stern RJ, Miller NR (2018) Did the transition to plate tectonics cause Neoproterozoic Snowball Earth? Terra Nova 30(2):87–94 Torsvik TH, Carter LM, Ashwal LD, Bhushan SK, Pandit MK, Jamtveit B (2001) Rodinia refined or obscured: palaeomagnetism of the Malani igneous suite (NW India). Precambr Res 108(3–4):319–333 Westerhof ABP, Härmä P, Isabirye E, Katto E, Koistinen T, Kuosmanen E, Lehto T, Lehtonen MI, Mäkitie H, Manninen T, Mänttäri I, Pekkala Y, Pokki J, Saalmann K, Virransalo P (2014) Geology and geodynamic development of Uganda with Explanation of the 1:1,000,000 Scale Geological Map. Geological Survey of Finland Whittaker JM, Muller RD, Leitchenkov G, Stagg H, Sdrolias M, Gaina C, Goncharov A (2007) Major Australian-Antarctic plate reorganization at Hawaiian-Emperor bend time. Science 318 (5847):83–86

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202 Williams S, Wright NM, Cannon J, Flament N, Müller RD (2021) Reconstructing seafloor age distributions in lost ocean basins. Geosci Front 12(2):769–780 Blades ML, Collins AS, Foden JD, Payne JL, Stüwe K, Hassan M, Abu-Alam T (submitted) Age and hafnium isotope evolution of the Butana and Chad—does Butana Represent The Eastern Margin Of The Sahara Metacraton? J Afr Earth Sci

A. S. Collins et al. Young A, Flament N, Maloney K, Williams S, Matthews K, Zahirovic S, Muller RD (2019) Global kinematics of tectonic plates and subduction zones since the late Paleozoic Era. Geosci Front 10(3):989–1013 Zhang R, Jiang D, Ramstein G, Zhang Z, Lippert PC, Yu E (2018) Changes in Tibetan Plateau latitude as an important factor for understanding East Asian climate since the Eocene: a modeling study. Earth Planet Sci Lett 484:295–308

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Early Ensimatic Stage of the Arabian-Nubian Shield Mahmoud Hassan, Abdel-Rahman Fowler, Osama Dessouky, and Tamer Abu-Alam

Abstract

Keywords

The majority of geological investigations that deal with the Arabian-Nubian Shield are concerned with the processes of ocean closure, subduction, orogenesis and crustal growth, in relation to the assembly of Gondwanaland in the late Neoproterozoic. Other valuable published works deal with the earlier development of the Arabian-Nubian Shield in the light of the configuration of Rodinia (assembly and rifting) and the Mozambique Ocean. Progress in modern geochronological and structural data from the Arabian-Nubian Shield reveals that some of the Arabian-Nubian Shield rocks were derived from older crustal material and were affected by tectonic events of the early ensimatic stage of the Mesoproterozoic Rodinia breakup. The studies of Arabian-Nubian Shield ophiolites and related mélange rock units, representing remnants or fragments of earliest simatic (maficultramafic) lithosphere, provide essential constraints on the oceanic realm predating the accretionary and collisional stages of the Arabian-Nubian Shield (*780– 600 Ma). Understanding the complete tectonic evolution of the Arabian-Nubian Shield requires providing special attention to the structural, petrological, geochemical and geochronological characteristics of its early primitive stage during the Rodinia breakup.

Rodinia breakup Ensimatic stage peninsula

M. Hassan (&) Geology Department, Faculty of Science, Suez Canal University, Ismailia, Egypt e-mail: [email protected] A.-R. Fowler Geology Department, College of Science, United Arab Emirates University, Al-Ain, United Arab Emirates O. Dessouky Nuclear Materials Authority of Egypt, P.O. Box 530, El Maadi, Cairo, Egypt T. Abu-Alam Universitetsbiblioteket, University of Tromsø—The Arctic University of Norway, 9037 Tromsø, Norway

9.1





Gondwana Ophiolites









Arabian-Nubian shield Mesoproterozoic Sinai

Introduction

The cycle of assembly and breakup of supercontinents is now a widely accepted concept among researchers of Earth history. The Mesoproterozoic supercontinent Rodinia formed during the Grenvillian time (*1300–900 Ma ago) by accretion of most of the earlier continental fragments into a single continental mass (Pesonen et al. 2003). At the beginning of the Neoproterozoic, the Rodinia continent itself began to disintegrate into the original component continents by rifting along older sutures. A renewed assembly (though in a different configuration) of the supercontinents Gondwana and Laurasia followed thereafter. These eventually fused to form Pangaea in the late Neoproterozoic (Meert and Van der Voo 1997; Hoffman 1991, 1999; Pisarevsky et al. 2003). The Arabian-Nubian Shield (ANS) bears fragments of this entire history of supercontinent assembly and breakup and is therefore a key role in understanding this history. The ANS is a constituent of the East African Orogen and formed by collision of juvenile Neoproterozoic intra-oceanic arcs that were generated within and around the margins of a large oceanic tract (the Mozambique Ocean), which formed in association with the Rodinia breakup (Stern 1994; Stein and Goldstein 1996). The crust generated during formation and closure of the Mozambique Ocean was ultimately combined to form western Gondwana during the Neoproterozoic (e.g., Stern 1994; Whitehouse et al. 2001; Meert 2003; Collins and Pisarevsky 2005). The early simple model for the ANS, involving stacking of juvenile Neoproterozoic arcs (Stern 1994), has evolved

© The Author(s), under exclusive license to Springer Nature Switzerland AG 2021 Z. Hamimi et al. (eds.), The Geology of the Arabian-Nubian Shield, Regional Geology Reviews, https://doi.org/10.1007/978-3-030-72995-0_9

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into a more sophisticated one due to the growing ANS database of geochronology and geochemistry (Johnson and Woldehaimanot 2003; Be’eri-Shlevin et al. 2012; Eyal et al. 2014). Tracing the tectonic evolution of the ANS from oceanic lithosphere to island arc and continental lithosphere (Zoheir and Klemm 2007) is a part of the history of supercontinent assembly and breakup. The ANS history begins with the breakup of Rodinia in the earliest Mesoproterozoic, ushering in an ensimatic stage, which was followed by the processes of ocean closure, subduction, orogenesis and crustal growth (Dessouky et al. 2020), leading to the assembly of Gondwana in the late Neoproterozoic (Stern 1994; Johansson 2014). The role of juvenile Neoproterozoic arcs in the formation of the ANS is well attested; however, the relations of these arcs to the early ensimatic stage of Rodinia breakup are a matter of discussion. In standard models, Rodinia was assembled over the interval 1300–900 Ma (Li et al. 2008). The breakup and rifting events were delayed for about 100– 150 My after assembly (Li et al. 2008; Meert and Torsvik 2003; Evans 2009; Nance et al. 2014). Events in the 100– 150 My after Rodinia assembly are not recorded in the Neoproterozoic ANS (Stein and Goldstein 1996; Stein 2003; Evans 2009; Nance et al. 2014). However, the central and southern ANS preserves some terrane protoliths of Tonian age (older than 850 Ma) in addition to few scattered arcs within the ANS (Johnson 2014). The locally strongly deformed, metamorphosed and altered ophiolites and mélanges outline sutures separating the constituent terranes of the ANS, and also separate the ANS from the rest of east and west Gondwana (Kröner et al. 1992; Abdelsalam and Stern 1996; Abdelsalam et al. 2003). These complex detached slices of ophiolitic mélanges are regarded as direct or indirect evidence for oceanic material of the Mozambique oceanic crust that may have been emplaced over arcs and continental margins at destructive plate boundaries (Bakor et al. 1976; Neary et al. 1976; Gass 1977; Kröner 1985; Shackleton 1988; El-Gaby and Greiling 1988; Zoheir and Klemm 2007), while Abd El-Rahman et al. (2012) classified the widely distributed ophiolitic units in their present coordinates as Neoproterozoic intra-oceanic island arc assemblages based on their geochemical signatures. This chapter summarizes the available geological data on the development of the ANS in the light of the pre-rifting configuration of Rodinia and the Mozambique Ocean. In addition, the characteristics and distribution of ophiolites are assessed as physical evidence for their oceanic setting that predated Gondwana assembly.

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9.2

Rodinia Pre-rifting Configurations

The concept of a late Proterozoic supercontinent, proposed in the 1970s—early 1990s (Piper 1976; McMenamin and McMenamin 1990; Dalziel 1991; Hoffman 1991), was suggested due to the existence of several continent-scale collision belts of 1200–1000 Ma age range (broadly Grenvillian and Kibaran) that were used to control the matching of continent margins in this time range. The early attempts could not be rigorously tested by the paleomagnetic data available at the time, which had poor age constraints. Valentine and Moores (1970) called this late Proterozoic supercontinent “Pangaea I” and suggested that the diversification of life forms on Earth that occurred toward the end of the Precambrian, was due to the breakup of this supercontinent. The widespread appearance of continental shelf environments led Bond et al. (1984) to suggest that the breakup of Pangaea I had occurred by the end of the Proterozoic. Subsequently, Pangaea I was renamed “Rodinia” by McMenamin and McMenamin (1990). In standard models, Rodinia incorporated all of the continents into a single mass at the end of the Mesoproterozoic. The process of assembly began at around 1100 Ma and continued until completion at about 900 Ma, followed by breakup, beginning 100–150 Ma later (Meert and Torsvik 2003). The improved paleomagnetic database and more precise dating have confirmed the existence of Rodinia, but have permitted a series of possible competing configurations with significant repercussions for the Rodinia breakup history. All configurations place the later continent of Laurentia in the inner part of Rodinia, with constraints placed by the need to connect the Grenvillian-Kibaran age collision belts across the continents. In order to understand the paleogeographic position of the Mesoproterozoic outcrops of the ANS during the evolution of the Rodinia, we present a short summary of different Rodinia configuration models.

9.2.1 SWEAT Model The earlier configurations by Dalziel (1991, 1997) and Torsvik et al. (1996) are described as “traditional” or “archetypal.” One popular model of this category was forwarded by Moores (1991) in the SWEAT (South Western US and East Antarctica) hypothesis that set a close proximity of the western side of Laurentia (SW US) against East Antarctica (Fig. 9.1). However, the hypothesis of the connection between western Laurentia, Australia and East

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Fig. 9.1 SWEAT configuration of Rodinia at *700 Ma, modified after Hoffman (1991), showing the Mozambique Ocean. In this figure, the subduction zone location was guided by Spencer et al. (2015)

Antarctica was originally suggested by Bell and Jefferson (1987) and Eisbacher (1985), based on similarity of Neoproterozoic stratigraphy. Moores (1991), Dalziel (1991) and Hoffman (1991) suggested that the SWEAT connection existed as early as 1900 Ma. However, paleomagnetic data (e.g., Powell et al. 1993), geological discontinuities across Rodina continent, as suggested within SWEAT (e.g., Borg and DePaolo 1994), and the presence of Grenvillian orogenic belts across and within the suggested SWEAT configuration (e.g., Berry et al. 2005; Fioretti et al. 2005; Borg and DePaolo 1994) argue against the old age (i.e., 1900 Ma) of the SWEAT. Paleomagnetic data and the minimum collision age of the Grenvillian orogenic belts suggest that the SWEAT may have existed during the time interval ca. 1050– 720 Ma (e.g., Powell et al. 1993). Van der Voo et al. (1984) suggested that the SWEAT arrangement existed until 580 Ma or ca. 650 Ma. In the SWEAT model, Greenland, Baltica and Amazonia were connected to the eastern side of Laurentia (e.g., Weil et al. 1998; Pesonen et al. 2012). The SWEAT hypothesis has been criticized, but was also later reinvigorated by isotopic evidence that found Antarctic granites (1400 Ma old) with similar distinctive chemistry to Laurentian granites of the same age (Goodge et al. 2008), suggesting their near proximity at 1400 Ma.

9.2.2 “Missing-Link” Model Detailed geological mapping of the component crustal provinces across Rodinia suggests a mismatch between Australia-East Antarctica and Laurentia within the SWEAT model. On the other hand, continuation and similarities in both Neoproterozoic stratigraphy (e.g., Eisbacher 1985) and crustal provinces of South China (i.e., the Yangtze Block) and southeastern Australia, western and southern Laurentia make the South China Block a good candidate to fill the mismatch between Australia-East Antarctica and Laurentia within the SWEAT model as suggested by Li et al. (1995). This insertion of the South China Block into the SWEAT model configuration is called the “Missing-Link” model (Fig. 9.2; Li et al. 1995). The missing-link model is supported by several other geological and geophysical observations. For example, the Cathaysia block (southeastern South China) has crustal composition similar to the Belt Basin of western North America in southwestern Laurentia (Ross et al. 1992; Li et al. 2002a, b; Ross and Villeneuve 2003). Non-Laurentian detrital grains from the Belt Basin (Ross and Villeneuve 2003) constrain the formation age of the “Missing-Link” model to lie between 1140 and 900 Ma and can explain the occurrence of Grenvillian orogenic belts

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Fig. 9.2 “Missing-link” configuration showing the setting of Arabia and Nubia microcontinents within the Mozambique Ocean after breakup of Rodinia (720 Ma) (modified from Li et al. 2008)

along the western Laurentia, northern Queensland, King Island-Transantarctic Mountains and the Albany-FraserMusgrave belt (Fig. 9.2) (Li et al. 2003c; Ling et al. 2003). Support for the “Missing-Link” hypothesis came from recognition of mantle plume activity and associated ultramafic dykes and sills in South China that could be traced in the model into Australia (Li et al. 1999, 2008). The plume was also implicated in the mechanism of breakup of Rodinia at *820 Ma (Wingate et al. 1998; Li et al. 2003a).

9.2.3 Other Models Two other models explain the connection between Australia and Laurentia through Southwest US or Mexico, which are referred to as Australia–Western US (AUSWUS) and Australia–Mexico (AUSMEX), respectively. The AUSWUS model was suggested to explain similar lineament fractures along the margins of the eastern Australian craton and western Laurentia (Brookfield 1993). There are several lines of observations that argue against the AUSWUS model, including that these linear fractures were formed around 600 Ma and are therefore not related to Rodinia (Direen and Crawford 2003). Wingate et al. (2002) suggested that the AustraliaLaurentia connection was in existence at ca. 1070 Ma,

through a connection between northern Queensland of Australia and Mexico of southern Laurentia (i.e., the AUSMEX model; Fig. 9.3). However, this model is based on an assumption that the paleomagnetic pole at 1070 Ma was at Bangemall Basin in the Edmund Fold Belt of Western Australia, and neglects other pole locations suggested for the same age from central Australia. Further work on the central Australian poles (Schmidt et al. 2006) argued against the 1070 Ma pole in the Bangemall Basin. The “SAMBA” model of Johansson (2014) discusses the change from Rodinia to Gondwana in the frame of “orthoversion” breakup and assembly of the supercontinent of Mitchell et al. (2012). Johansson (2009, 2014) proposed a kinematic scenario of two stages of counter-clockwise rotation around fixed point. The first 90° counterclockwise rotation was around a pole located at the corner of Kalahari, when the eastern South American and southern African cratons turned away from each other. The second 120° counterclockwise rotation around a pole located within the Kalahari craton, when East Antarctica, Australia and India (future East Gondwana) rotated away from the Laurentia and Amazonia opening of the Proto-Pacific Ocean. The Adamastor and Mozambique oceans were opened and closed during these rotations. Johansson (2014) assigned the Neoproterozoic island arcs stage of the ANS to this initial separation period. He also discussed the unusual amount of

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Fig. 9.3 AUSMEX configuration of Rodinia, showing the Mozambique Ocean as an embayment in the western side of the supercontinent (Pisarevsky et al. 2003; Wingate et al. 2002)

its juvenile Neoproterozoic continental crust (Stern 1994) with respect to the high rate of subduction and consumption of oceanic crust during these rotational movement.

9.3

Rodinia Rifting

Rifting of a supercontinent is commonly associated with formation of dyke swarms and syn-rift magmatism related to deep mantle activities, e.g., mantle plumes (Li et al. 2008). Dyke swarms from Rodinia show different ages, e.g., ca. 825 Ma and ca. 780 Ma for the Gairdner-Amata dyke swarm in central and southeastern Australia; and the radiating Gunbarrel dyke swarms from western Laurentia, respectively, (Sun and Sheraton 1996; Wingate et al. 1998; Harlan et al. 2003a, b). Episodes of rift-related magmatism are observed and can be correlated from both southern China and eastern Australia. These rift-related magmatic events

occurred on a time interval covering the period 830–720 Ma (Powell et al. 1994; Li et al. 1995, 1999, 2002a, b, 2003b; Preiss 2000; Wang and Li 2003). This may indicate the rifting of the Rodinia supercontinent as having occurred between 830 and 720 Ma. Rodinia breakup scenarios mostly agree on the rift separation of Australia–Antarctica from Laurentia at about 800 Ma to form the Proto-Pacific Ocean (Moores 1991; Dalziel 1991; Hoffman 1991). This event was followed a little later by the separation of Laurentia from Amazonia-Rio Plata and Baltics, to form the Iapetus Ocean. The process of opening of the Pacific and Iapetus freed Laurentia from the core of Rodinia. The northern and southern clusters (Fig. 9.4) of continental masses that flanked Laurentia in Rodinia formed the nuclei of later eastern and western Gondwana, respectively. Thereafter, the reunification of East and West Gondwana closed the Mozambique Ocean to form the East African Orogen at 500 Ma. The ANS, as the

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Fig. 9.4 Rodinia configuration from Scotese (2009) showing the northern and southern clusters of continental masses that flanked Laurentia and the Mozambique Ocean at 750 Ma. The Mozambique Ocean is shown separated from the Adamastor Ocean by the Congo-Sao Francisco-Saharan continent. Symbols are: NC = North China, Ar = Arabia, Cm = Cimmeria, SC = South China, In = Indochina, Sb = Sibumasu terrane, Au = Australia, Ind = India, M = Madagascar, H = Hijaz arc, An = Antarctica, CC = Congo, K = Kalahari, L = Laurentia, Sa = Saharan metacraton, Sf = Sao Francisco, RP = Rio Plata, A = Amazonia, B = Baltica, S = Siberia, Cd = Cadomian arc, WA = West Africa

northern part of the East African Orogen, is in tectono-lithological contact with the eastern margin of the Saharan Metacraton (Küster et al. 2008).

9.4

Origins and Configurations of the Mozambique Ocean—Rift or Remnant?

The current understanding of the opening of the Mozambique Ocean varies according to the preferred initial configuration and later breakup history of the Rodinia supercontinent. In the SWEAT configurations (Fig. 9.1), the Mozambique Ocean lay to the west of Rodinia (Hoffman 1991) and was presumably being subducted beneath Rodinia at 750 Ma (Spencer et al. 2015). In this scenario, the Mozambique Ocean consisted of >1100 Ma oceanic crust and was coincident with the Pan-Rodinia ocean, called the Mirovoi Ocean. Dalziel (1991, 1997) gave the name Mozambique Ocean to the remnants of the Miravoi Ocean, following closure of parts of the Miravoi during the breakup of Rodinia. Cawood (2005) similarly regarded the Mozambique Ocean as a remnant of the Mirovoi Ocean.

In the AUSWUS and AUSMEX configurations (Fig. 9.3), the more southerly position of Australia–East Antarctica with respect to Laurentia interposes them between the Mozambique Ocean and the Adamastor Ocean to its east. In these configurations, the Mozambique Ocean occupies an embayment at the western side of Rodinia (Fig. 9.3). The Mozambique Ocean had an equatorial location, with northern margins constrained by Antarctica–India–Mozambique and southern shores against the Congo and Kalahari continents. Scotese (2009) suggested that both the Mozambique and the Adamastor Oceans were at the western side of the Rodinia, where the Congo continent formed the main barrier between the two oceans at ca. 750 Ma (Fig. 9.4). In all of these models, the Mirovoi-cum-Mozambique Ocean was floored by pre-Rodinia ocean crust. In the “Missing-Link” configuration of Li et al.’s (2008), there are two small continental masses labeled “Arabia” and “Nubia” within the Mozambique Ocean (Fig. 9.2), referring to the Afif-Abas and Bayuda terranes (Fig. 9.5). Li et al. (2003c) and Hargrove et al. (2004) related the igneous rocks with age range 850–750 Ma in these terranes to the Neoproterozoic Rodinia superplume activity. Scotese (2009) illustrates these as the substrate for a southward vergent

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Fig. 9.5 Current geographic coordinates showing the locations, outcrops, terranes, suture zones and shear zones mentioned in the text (modified after Johnson et al. (2011) and references therein)

subduction-related arc referred to as the “Hijaz Arc,” which is also referred to as the “ANS arc” by Meert and Torsvik (2003).

The position of the ANS in most of Rodinia models lies between the Indian and Saharan cratons to the northeast of Madagascar (Li and Powell 2001; Collins and Pisarevsky 2005;

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Li et al. 2008), which is relevant for the subsequent late Neoproterozoic configuration of the amalgamated Gondwanaland. The ANS took the place of the Mozambique Ocean within fragmented Rodinia; therefore, the main physical evidence of an oceanic setting for the ANS should be the ophiolite nappes, melanges and complexes that have been dated to 890– 690 Ma (Stern 1994; Stern et al. 2004). However, lithogeochemistry and associated chromite compositions suggest that most, if not all, of the ANS ophiolites (Fig. 9.6) are arc related and derived from fore-arc and back-arc basins (Farahat 2010; Azer 2014; Obeid et al. 2016) or other related basins (Gamal El Dien et al. 2016) and are not fragments of the Mozambique Ocean. Stern (1994) noted that there was no certain data about the characteristics of the oceanic crust of the Mozambique Ocean. However, Stern (op. cit.) concluded on other grounds that the Mozambique Ocean must have existed and was probably quite large, even Pacific in scale. These conclusions were based on the fact that back-arc and fore-arc basins require older (by *100 My) denser oceanic crust to have subducted beneath the arc in order to form. The abundance of these basins in the ANS record required substantial subduction and a correspondingly large amounts of Mozambique Oceanic crust. With commencement of the ocean diminishing at *780 Ma, the Mozambique Ocean would have to have contained oceanic volcanic arcs deposited on its thin juvenile crust with age 870 Ma (Johnson and Woldehaimanot 2003). The earliest recorded magmatic events in the ANS were delayed to ca. 854 Ma (Hargrove et al. 2006), concurrent with the deformation and metamorphism along the greater Mozambique Belt, with their peak ranges geochronologically between 720 and 620 Ma (Meert 2003; Tohver et al. 2006). Even greater age is possible to mark the initiation of the proto-ANS if parts of the Mozambique Ocean were extant at the time of Rodinia assembly as oceanic material outboard from Rodinia, as suggested by other authors (Cawood 2005). Opening of the Mozambique Ocean, as a consequence of Rodinia breakup, is mentioned by Fritz et al. (2013) for the southern ANS at *900–850 Ma, and this concept of the Mozambique opening remains popular in ANS studies. Some support for a rift opening stage of the Mozambique comes from the opening of the “Neo-Mozambique” Ocean between Azania (Proto-Madagascar) and the Congo (Fitzsimons and Hulscher 2005; Jöns and Schenk 2008).

9.5

Remnants of Rodinia Within the ANS

Practical knowledge about the origin of the ANS has accumulated from revised scenarios discussing the magmatic and deformational events in its juvenile part, and the geological history of its older parts. Most of the older ages are interpreted to represent crustal reworking of assimilated older basement or inherited zircon grains from middle-

Neoproterozoic and pre-Neoproterozoic crust, which correlate with the timing of supercontinent assembly or breakup (Wang et al. 2020). The following section presents a summarized overview of ANS rock units, with geological and/or geochronological evidence that may indicate their connection with the Rodinia supercontinent. We will begin with rocks showing continental signatures, followed by those with oceanic signatures.

9.5.1 Rodinia Continental Signatures Within the ANS Fragmentary pre-Neoproterozoic Rodinia continental blocks have survived and participated in a number of subsequent orogenies. The East African Orogeny formed after Rodinia breakup as the ‘‘Transgondwanan supermountain range’’, and was associated with closure of the Mozambique Ocean (Fritz et al. 2013). The amalgamation of the East African Orogen began from the Neoproterozoic to early Cambrian, where it currently extends from northeastern Africa and western Arabia to the southern Mozambique and Madagascar (Stern 1994). The ANS represents an accretion-type orogeny at the northern section of the East African Orogen (Fritz et al. 2013; Abu-Alam et al. 2014; Abd El-Wahed and Hamimi 2020). The ANS is subdivided into crustal blocks representing various tectonostratigraphic terranes decorated with ophiolite sutures zones (e.g., Figs. 9.5 and 9.6; Berhe 1990). These blocks originally formed by deposition of volcano-sedimentary formations in an oceanic or marginal marine environment between dispersed continental fragments (Nance et al. 1986; Murphy and Nance 2003). Influence of the pre-Neoproterozoic continental crust or Rodinia-related events within the ANS was deduced based on geological observations (Agar 1985; Fowler and Hassen 2008; Abu-Alam and Stüwe 2009; Hassan et al. 2014; Fowler et al. 2015, 2018), age dating information (Calvez et al. 1983; Be’eri-Shlevin et al. 2009; Ali et al. 2009, 2010, 2012; Eyal et al. 2014; Abd El-Rahman et al. 2019) and Pb, Nd and Sr isotopic signatures (Baubron et al. 1976; Stacey et al. 1980; Fleck and Hadley 1982; Stacey and Stoeser 1983; Stacey and Hedge 1984; Stacey and Agar 1985; Windley et al. 1996). The pre-Neoproterozoic continental crust is sometimes referred to as “contaminated shield” that shows continental 207Pb/206Pb isotopic signatures (Type III Pb; Stoeser and Stacey 1988), low or commonly negative initial eNd values and old TDM ages (Hargrove et al. 2006). U–Pb single zircon geochronology databases in recently published reviews (Stern et al. 2010; Johnson 2014; Abd El-Rahman et al. 2019; Hamimi et al. 2021) feature four peaks of inherited zircon ages: late Mesoproterozoic (0.95– 1.15 Ga), early Proterozoic (1.7–2.1 Ga), late Archean (2.4– 2.8 Ga); and early Archean (>3.2 Ga). The oldest age dating

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Fig. 9.6 Distribution of the ophiolites in the ANS (modified after Abu-Alam and Hamdy 2014 and references therein)

results from the ANS were obtained from the granitic gneisses of the Al-Mahfid terrane (Fig. 9.5) in Yemen indicating Archean origin (2550 and 2560 Ma—Whitehouse et al. 1998). The Afif-Abas terrane and the Khida subterrane (2000 Ma; Li et al. 2008; Tucker et al. 2011) represent a transition from continental to oceanic setting, based on Pb and Nd isotope data (Johnson and Woldehaimanot 2003).

Similar Paleoproterozoic crustal material is reported from the Eastern Desert of Egypt and Halfa regions of the Nubian part of the shield, as indicated by the existence of pre-Neoproterozoic crust (Ali et al. 2013, 2018). Other Paleoproterozoic ages (1660 and 1680 Ma) were detected from the Khida granites. Small bodies of biotite granite gneiss of the South Libab orthogneiss of the Arabian part of

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the ANS show Paleoproterozoic age (1900–1750 Ma) estimated from the outer zircon rims while the inner cores give age of 2100–2400 Ma (Stoeser et al. 2001). The South Libab orthogneiss is less affected by metamorphism than the neighboring late Archean-early Paleoproterozoic anorthosite of the Muhayil suite (1850–1670 Ma, Stoeser et al. 2004). Some deformed pre-tectonic granites in the Central Eastern Desert of Egypt show upper intercept ages of ca. 1650 Ma that were interpreted by Sultan et al. (1990) as the maximum age of the inherited material that was supported by the higher initial 87Sr/86Sr and 207Pb/204Pb ratios. Anorthosite of Gebel El Asr (Fig. 9.5), Western Desert of Egypt, which represents a reworked part of the Saharan Craton (Harms et al. 1990; Sultan et al. 1994), yielded upper intercepts of 1922–2141 Ma via TIMS U–Pb zircon concordia age. The negative eNd(t) values (−4.3 to −12.9) and the Palaeoproterozoic model ages (2430–2010 Ma) from the meta-sedimentary protolithic material of the leucocratic gneisses and schists of the Bayuda Desert, are evidence of their pre-Neoproterozoic continental origin, and can be ascribed to the Saharan Metacraton (Küster and Liégeois 2001; Küster et al. 2008). The Erkowit pluton of the Haya terrane (Kröner et al. 1991; Reischmann et al. 1992) in Sudan, and the Makkah Suite (Fleck 1985) in the Jiddah terrane, show geochronological signatures of upper Tonian age. The distribution of these inheritances does not yet provide a coherent picture, due to the wide spacing of data in the ANS; however, the clustering of 900–1100 Ma inherited zircons ages in the northernmost ANS may be symptomatic. The northernmost ANS (Fig. 9.5) encloses a few widespread 900–1100 Ma inherited zircons ages (such as Sa’al metavolcanics, Taba and Rutig Conglomerate) suggesting the influence of the Neo-Mesoproterozoic older crust of assembled Rodinia. The Sa’al metavolcanics show late Mesoproterozoic-early Tonian single zircon grains of *1.03–0.93 Ga age (Be’eri-Shlevin et al. 2009, 2012; Eyal et al. 2014). Be’eri-Shlevin et al. (2009, 2012) and Eyal et al. (2014) constrained the depositional age of the protolith of the Ra’ayan formation schists in the Sa’al area to be 950 Ma. The Sa’al area is surrounded by quartz diorite and quartz monzonite gneiss with crystallization age of ca. 930 Ma and 819 ± 4 Ma, respectively. The high 87Sr/86Sr ratios and the low eNd values of the meta-sedimentary rocks (Abu Anbar et al. 2009) and the whole-rock eNd(T) of the Kibaran-aged schists (+2; Be’eri-Shlevin et al. 2009) indicate an old crust of pre-Neoproterozoic origin. Based on the aforementioned geochronological data in addition to calculated geothermal gradient and structural analyses, the firstly developed foliations (S1) in older rocks of Sa’al-Zaghara metamorphic complex strongly signified tectonic origin via NNW–SSE directed crustal extension (Hassan et al. 2014; Fowler et al. 2015). A similar observation from the Feiran-Solaf metamorphic complex indicates that the D1

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event is an extension event (Fowler and Hassen 2008; Hassan et al. 2021). The structural data is interpreted to be correlated to the breakup of Rodinia (Hassan et al. 2014; Fowler et al. 2015; Hassan et al. 2021) and is supported by geochronological ages (838 ± 16 Ma; Abu Anbar et al. 2004) and petrological observations (Abu-Alam and Stüwe 2009). In the Nubian part of the ANS, Abd El-Rahman et al. (2019) assigned the cluster of negative eHf(t) values of detrital zircons of the Um Had Conglomerate to older crustal source rocks. Latest Mesoproterozoic (Kibaran age) crust is also exposed in the Bayuda Desert (Küster et al. 2008). Perelló et al. (2020) referred to ages of >860 Ma in the Eritrean highlands of the Nakfa terrane (southern ANS), as dating the initiation of the Mozambique Ocean, consistent with Rodinia breakup that rapidly transfigured to subduction and island arc stages, while Abd El-Rahman et al. (2017) pointed to convincing contemporaneous existence of a continental arc system and an intra-oceanic arc system in the northern Nubian Shield.

9.5.2 Mozambique Ocean Signatures Within the ANS 9.5.2.1 Setting and Occurrence The ANS began to materialize at ca. 870 Ma, where oceanic volcanic arcs were formed on the thin crust of the Mozambique Ocean. The shield was collected through different amalgamation events, including subduction, basin closure, ophiolite obduction (e.g., Harms et al. 1990; Abdelsalam et al. 2003; Abu-Alam et al. 2014; Abd El-Wahed and Hamimi 2020). Ophiolite outcrops are widespread over the entire ANS, marking sutures between major terranes and crustal blocks (Fig. 9.6). These dismembered ophiolites occupy a total area of about two million square kilometers (Stern et al. 2004) and were previously recognized as allochthonous pieces of oceanic crust (Genna et al. 2002). More deformed and metamorphosed outcrops of these fragments are found in the tectonic mélanges. Reconstruction of some ophiolitic successions from the ANS indicates a thickness of 2.5–6 km of the ophiolite succession. The ANS ophiolites are frequently associated with a 1–3 km thick sequence of dense cumulate ultramafic rocks, which pertain to a transition zone between the seismic and petrologic Moho. The ophiolites of the northern part of the ANS have been disrupted by the Najd Fault System (NW-trending strike-slip crustal scale shear zone; Sultan et al. 1988). Complete ophiolitic sequences have been described from Gabal Ess in Saudi Arabia, Wadi Ghadir and Gabal Gerf in Egypt, Wadi Onib in Sudan (Johnson et al. 2004; El-Bayoumi 1983; El Akhal 1993; Hussein 2000; Zimmer et al. 1995), which include peridotites, gabbros, sheeted dykes, pillow lavas and sedimentary rocks that reflect a

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deep-sea environment. The peridotites are predominantly tectonized harzburgite, and rarely lherzolite and dunite. Numerous ANS ophiolites have well-preserved transition zones that distinguish the mafic lower crust gabbros from the ultramafic upper mantle peridotites. These transition zones are characterized by interlayered pyroxenite, wehrlite, lherzolite, dunite and/or chromite at the base rising to parts gradually dominated by gabbro. Other ophiolites have thin to non-existent transition zones (e.g., at Fawkhir; El-Sayed et al. 1999). The ophiolitic gabbros are widespread in the ANS ophiolites. Original igneous textures can be recognized, but metamorphic textures that indicate recrystallization under greenschist to amphibolite facies conditions are common. The ophiolitic gabbros are mainly represented by pyroxene gabbro, while olivine gabbros are far less common. Clinopyroxene generally dominates over orthopyroxene. Layered gabbroic rocks exist and are represented mainly by melagabbro and anorthositic gabbro. The layered gabbros are composed of alternating plagioclase and amphibole-rich layers (e.g., Ess ophiolite; Shanti 1983). The accumulation of sequenced layered gabbro stratigraphy shows significant crystallization from olivine ± chromite-clinopyroxeneplagioclase (Price 1984), to olivine ± chromiteclinopyroxene-orthopyroxene-plagioclase (Nassief et al. 1984), or, less commonly, olivine ± chromiteorthopyroxene-clinopyroxene-plagioclase (Abdel-Rahman 1993). At the mineral level, plagioclase compositions in Sol Hamed gabbros change from An70–85 at the base of the accumulation to more sodic compositions at the top of the sequence (Fitches et al. 1983). Sheeted dikes are a common component of ANS ophiolites. Wherever observed, they typically transition downwards into the ophiolitic gabbros and grade upwards into pillow basalts. Sheeted dikes are reported from the following ophiolites (Figs. 9.5 and 9.6): Rahib (Abdel-Rahman et al. 1990), Ess (Shanti and Roobol 1979), Sol Hamed (Fitches et al. 1983), Ingessana (Price 1984), Wadi Ghadir (El-Bayoumi 1983), Gerf (Zimmer et al. 1995) and Thurwah (Nassief et al. 1984). Sheeted dykes for other ANS ophiolites are not identified or are poorly developed (e.g., Al Ays and Fawkhir: Bakor et al. 1976; El-Sayed et al. 1999). Lavas with well-preserved pillow structures are a characteristic feature of ANS ophiolites (Stern et al. 2006). These are of significant importance due to the fact that lava compositions provide valuable clues about tectonic setting and the nature of melt generation.

9.5.2.2 Magma Type Investigating the geochemistry and magma type of the best preserved and well-exposed ophiolites in the ANS may provide definite clues and establish their connection to the tectonic model of the ANS. The ANS ophiolites were

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derived from magmas that are often classified as subalkaline suites, characterized by low-K and moderate-Ti contents that are moderately fractionated (Pallister et al. 1988; Zimmer et al. 1995; Wolde et al. 1996; Stern et al. 2004; Abd El-Rahman et al. 2009a, b; Gahlan et al. 2015; Abdel-Karim et al. 2016). The source magma of the gabbros, sheeted dykes and the pillow lavas was fractionated, with high Mg# content, but less than the content of Mg# for mantle peridotite. Eritrean and Ethiopia suspected ophiolitic pillow lavas typically contain less TiO2, Y and Zr and have higher K2O, Mg#, Cr and Ni than most of the other ANS ophiolitic lavas. On the other hand, the pillow lavas from the Sekerr ophiolite are also distinct, with much higher TiO2, Y and Zr along with lower Mg#, Cr and Ni than most of the other ANS ophiolitic sequences. Obviously, the lavas from ANS ophiolites are mostly tholeiitic and show little affinity toward calc-alkaline nature and include a significant subordinate proportion of boninites (Zimmer et al. 1995; Takla et al. 2002; Abd El-Rahman et al. 2009a, b). The chemical features of the ANS cumulate sequences indicate that the magmatic systems, represented by the ophiolites, have undergone fractionation from a system dominated by primitive mafic magmas to ones dominated by highly evolved magmatic liquids. The major element characteristics are similar to a wide range of oceanic lavas, including MORB, some intra-oceanic arc lavas and back-arc basin basalt. Geochemically, most of the ophiolites of the ANS were formed in supra-subduction zone settings (SSZ) (Bakor et al. 1976; Nassief et al. 1984; Price 1984), with chemical signatures indicate fore-arc tectonic setting, while MORB signatures were recorded as back-arc basin oceanic crust in the southern ophiolites of the ANS (Bakor et al. 1976; Frisch and Al-Shanti 1977; Kröner 1985; Pallister et al. 1988; Berhe 1990; Ahmed et al. 2001; Johnson and Kattan 2001; Stern et al. 2004; Azer and Stern 2007; Abd El-Rahman et al. 2009a, b; Farahat et al. 2010; El Bahariya 2021).

9.5.2.3 Tectonic Setting Geochemical characteristics of the ANS ophiolites and field relations that indicate obduction along destructive plate boundaries (Fig. 9.7) are strong evidence for the formation of these ophiolites by plate-tectonic processes similar to those of Phanerozoic tectonics (Stern et al. 2004). Zimmer et al. (1995) suggested that the ANS ophiolites represent fragments of normal oceanic crust. Other ophiolites, e.g., Darb Zubaydah in Saudi Arabia and Dahanib in Egypt may represent roots of island arcs or autochthonous layered intrusions, respectively, (Dixon 1981). Flexible tectonic styles are accepted for these widely distributed ophiolites in the ANS, even for the constituents of same broadly extended individual exposure. U–Pb zircon and Pb–Pb evaporation techniques on zircons separated from

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Fig. 9.7 Three-dimensional model illustrating the tectonic evolution of the ANS ophiolites (modified after Abu-Alam and Hamdy 2014). a formation of MORB, fore-arc and back-arc ophiolites at different stages of evolution of the ANS. b Thrusting and obduction of the ophiolitic sequence

gabbros and plagiogranites (Stacey et al. 1984; Pallister et al. 1988, Kröner et al. 1992; Zimmer et al. 1995) and Sm–Nd mineral and whole-rock techniques (Claesson et al. 1984; Zimmer et al. 1995; Worku 1996) discriminate between the different evolutionary stages that signify the tectonic settings of the ophiolites of the ANS. The older ophiolites in the western side of the Arabian part of the shield and the associated intrusions (*870–830 Ma) imply ocean-floor magmatism preceding the ocean, diminishing at *780 Ma (Stern et al. 2004). Their oceanic affinity infers an ensimatic island arc attribute, due to development in a large ocean basin (especially the Bir Umq suture; Fig. 9.5) as the oldest ophiolites (Ahmed and Hariri 2008). Stern (2004) considered that the zircon data document oceanic crust generation between *810 and *730 Ma ago in the Nubian segment of the ANS, whereas younger ophiolites in the central and eastern Arabian Shield (Hulayfah-Ruwah, Halaban and Al Amar; Fig. 9.5) developed within supra-subduction zones of smaller basins, revealing volcanic arc affinity (Dilek and Ahmed 2003) along the latest amalgamated arcs of the Ad

Dawadimi and Ar Rayn terranes, which are believed to be the youngest contributor terranes in the continental growth of the ANS (Stern et al. 2004; Hamimi et al. 2015).

9.6

Conclusions and Open Questions

The presented overview directs attention to the most commonly accepted scenarios for the ANS formation in the early stage, where our understanding of the tectonic evolution of the ANS needs more improvement. The ANS, geochronologically, is considered as juvenile, although some igneous rocks in the northern segments of the shield contain inherited zircon from pre-Neoproterozoic sources and support the influence of assimilated pre-Neoproterozoic older continental materials in the ANS. These rocks have continental and oceanic signatures affiliated to the Rodinia supercontinent and the Mozambique Ocean, respectively, as supposed by several contributors. In our contribution, we raise questions about the evolution of the Mozambique Ocean in relation to

9

Early Ensimatic Stage of the Arabian-Nubian Shield

the early stage of the ANS. The absence of paleomagnetic data limits the ability of the researchers to constrain the paleogeographic location of the proto-ANS during the Rodinia time. The paleogeographic location of the proto-ANS will help to understand the nature, the extent and the size of the Mozambique Ocean. The abundance of ophiolites in northeastern Africa and Arabia, even with the limited isotopic data (mid-Neoproterozoic age 690–890 Ma; mean = 781 ± 47 Ma), supports their origin in supra-subduction zone, back-arc and fore-arc tectonic settings. The mixed subalkaline characteristics (tholeiitic and calc-alkaline) of ophiolitic lavas and the significant presence of subordinate boninites are characteristic of the same origin. Collectively, the cumulate ophiolite sequences with these mineral and lava compositions imply fractionated magmatic systems of the ANS ophiolites. Tectonically and geochronologically two different types of ophiolites reconstruct the progressive events of the development of the ANS. Although the MOR signatures signify some ophiolitic rocks in the ANS, the majority of the ophiolites were formed in subduction-related settings either as fore-arc or back-arc ophiolites. This challenges the existence and the size of ophiolites formed in the oceanic basins. Comprehensive integrated investigations are needed to resolve reservations about ophiolites in relation to the premature events of the evolution of the ANS and the real nature of the ophiolites (supra-subsubduction or MOR).

References Abd El-Rahman Y, Polat A, Dilek Y, Fryer BJ, El-Sharkawy M, Sakran S (2009a) Geochemistry and tectonic evolution of the Neoproterozoic incipient arc-fore arc crust in the Fawakhir area, Central Eastern Desert, Egypt. Precambr Res 175:116–134 Abd El-Rahman Y, Polat A, Dilek Y, Fryer BJ, El-Sharkawy M, Sakran S (2009b) Geochemistry and tectonic evolution of the Neoproterozoic Wadi Ghadir ophiolite, Eastern Desert, Egypt. Lithos 113:158–178 Abd El-Rahman Y, Polat A, Dilek Y, Kusky TM, El-Sharkawi M, Said A (2012) Cryogenian ophiolite tectonics and metallogeny of the Central Eastern Desert ofEgypt. Int Geol Rev 54:1870–1884 Abd El-Rahman Y, Seifert T, Gutzmer J, Said A, Hofmann M, Gärtner A, Linnemann U (2017) The South Um Mongul Cu–Mo–Au prospect in the Eastern Desert of Egypt: from a mid-Cryogenian continental arc to Ediacaran post-collisional appinite-high Ba–Sr monzogranite. Ore Geol Rev 80:250–266 Abd El-Rahman Y, Abu-Anbar M, Li X-H, Li J, Ling X-X, Wu L-G, Masoud AE (2019) The evolution of the Arabian-Nubian Shield and survival of its zircon U–Pb–Hf–O isotopic signature: a tale from the Um Had Conglomerate, central Eastern Desert, Egypt. Precambr Res 320:46–62 Abd EL-Wahed M, Hamimi Z (2020) Neoproterozoic tectonic events of Egypt. Acta Geologica Sinica-English Edition. https://doi.org/10. 1111/1755-6724.14410 Abdel Rahman EM, Harms U, Schandelmeier H, Franz G, Darbyshire DPF, Horn P, Müller-Sohnius D (1990) A new ophiolite

215 occurrence in NW Sudan—constraints on Late Proterozoic tectonism. Terra Nova 2(4):363–376 Abdel-Karim A-AM, Ali S, Helmy HM, El-Shafei HM (2016) A fore-arc setting of the Gerf ophiolite, Eastern Desert, Egypt: evidence from mineral chemistry and geochemistry of ultramafites. Lithos 263:52–65 Abdel-Rahman AM (1993) Geochemical and geotectonic controls of the metallogenic evolution of selected ophiolite complexes from the Sudan. Berliner geowissenschaftliche Abhandlungen 145:175 Abdelsalam MG, Stern RJ (1996) Sutures and shear zones in the ANS. J Afr Earth Sci 23:289–310 Abdelsalam MG, Abdelrahman EM, El-Faki EM, Al-Hur B, El-Bashier FM, Stern RJ, Thurmond AK (2003) Neoproterozoic deformation in the northeastern part of the Saharan Metacraton, northern Sudan. Precambr Res 123:203–221 Abu Anbar M, Solis Pichardo G, Hernandez Bernal MS (2004) Sm–Nd and Rb–Sr isotopes of Feiran gneisses and amphibolites: evidences of pre-Pan-African continental crust in Sinai, Egypt. 6th International conference on geochemistry. Alexandria University, Egypt, pp 727–745 Abu Anbar M, Finger F, Solis Pichardo G (2009) Evolution of the metamorphosed volcano-sedimentary rocks in Sa’al area, southern Sinai, Egypt: implications from lithology, geochemistry, mineralogy, P-T condition and age dating. Egypt J Geol 53:1–35 Abu-Alam TS, Hamdy M (2014) Thermodynamic modelling of Sol Hamed serpentinite, South Eastern Desert of Egypt: implication for fluid interaction in the Arabian-Nubian Shield ophiolites. J Afr Earth Sci 99:7–23 Abu-Alam TS, Stüwe K (2009) Exhumation during oblique transpression: the Feiran-Solaf region, Egypt. J Metam Geol 27:439–459 Abu-Alam TS, Hassan M, Stüwe K, Meyer S, Passchier C (2014) Multistage tectonism during Gondwana collision: Baladiyah Complex, Saudi Arabia. J Petrol 55:1941–1964 Agar RA (1985) Stratigraphy and palaeogeography of the Siham group: direct evidence for a late Proterozoic continental microplate and active continental margin in the Saudi Arabian Shield. J Geol Soc lond 142(6):1205–1220 Ahmed AH, Arai S, Attia AK (2001) Petrological characteristics of podiform chromitites and associated peridotites of the Pan African ophiolite complexes of Egypt. Miner Deposita 36:72–84 Ahmed Z and Hariri MM (2008) Neoproterozoic ophiolites as developed in Saudi Arabia and their oceanic and pericontinental domains. Arab J Sci Eng 33(1C):17–54 Ali KA, Stern RJ, Manton WI, Kimura J-I, Khamees HA (2009) Geochemistry, Nd isotopes and U-Pb SHRIMP zircon dating of Neoproterozoic volcanic rocks from the Central Eastern Desert of Egypt: new insights into the *750 Ma crust-forming event. Precambr Res 171:1–22 Ali KA, Stern RJ, Manton WI, Kimura J-I, Whitehouse M, Mukherjee SK, Johnson PR, Griffin WR (2010) Geochemical, U–Pb zircon and Nd isotopic investigations of the Ghawjah Metavolcanics of Northwestern Saudi Arabia: the *750 Ma crust-forming event correlated across the Red Sea. Lithos 120:379–392 Ali KA, Simon WA, Stern RJ, Moghazi AM, Mahbubul Ameen SM (2013) Hf isotopic composition of single zircons from Neoproterozoic arc volcanics and postcollisiongranites, Eastern Desert of Egypt: Implications for crustal growth and recycling in the Arabian-Nubian Shield. Precambr Res 239:42–55 Azer MK (2014) Petrological studies of Neoproterozoic serpentinized ultramafics of the Nubian Shield: spinel compositions as evidence of the tectonic evolution of Egyptian ophiolites. Acta Geol Polon 64:113–127 Azer MK, Stern RJ (2007) Neoproterozoic (835–720 Ma) serpentinites in the Eastern Desert, Egypt: fragments of fore-arc mantle. J Geol 115:457–472

216 Bakor AR, Gass IG, Neary CR (1976) Jabal al Wask, northwest Saudi Arabia: an Eocambrian back-arc ophiolite. Earth Planet Sci Lett 30:1–9 Baubron JC, Delfour J, Violette Y (1976) Geochronological measurements on rocks of the Arabian Shield. Kingdom of Saudi Arabia, BRGM Min. Rep. 76-JED-17, p 19 Be’eri-Shlevin Y, Katzir Y, Whitehouse MJ, Kleinhanns IC (2009) Contribution of pre-Pan-African crust to formation of the Arabian-Nubian Shield: new secondary ionization mass spectrometry U–Pb and O studies of zircon. Geology 37:899–902 Be’eri-Shlevin Y, Eyal M, Eyal Y, Whitehouse MJ, Litvinovsky B (2012) The Sa’al volcano-sedimentary complex (Sinai, Egypt): a latest Mesoproterozoic volcanic arc in the northern Arabian Nubian Shield. Geology 40:403–406 Bell RT, Jefferson CW (1987) A hypothesis for an Australian-Canadian connection in the Late Proterozoic and the birth of the Pacific Ocean. Australian Ins Mining Metall 87:39–50 Berhe SM (1990) Ophiolites in northeast and east Africa: implication for Proterozoic crustal growth. J Geol Soc Lond 147:41–57 Berry RF, Holm OH, Steele DA (2005) Chemical U–Th–Pb monazite dating and the Proterozoic history of King Island, southeast Australia. Australian J Earth Sci 52:461–471 Bond GC, Nickeson PA, Kominz MA (1984) Breakup of a supercontinent between 625 and 555 Ma: new evidence and implications for continental histories. Earth Planet Sci Lett 70:325–345 Borg SG, DePaolo DJ (1994) Laurentia, Australia, and Antarctica as a late Proterozoic supercontinent: constraints from isotopic mapping. Geology 22:307–310 Brookfield ME (1993) Neoproterozoic Laurentia-Australia fit. Geology 21:683–686 Calvez JY, Alsac C, Delfour J, Kemp J, Pellaton C (1983) Geological evolution of western, central, and eastern parts of the northern Precambrian shield, Kingdom of Saudi Arabia. Saudi Arabian Deputy Ministry for Mineral Resources, Open-File Report BRGM-OF-02-26 Cawood PA (2005) Terra Australis Orogen: Rodinia breakup and development of the Pacific and Iapetus margins of Gondwana during the Neoproterozoic and Paleozoic. Earth Sci Rev 69:249–279 Claesson S, Pallister JS, Tatsumoto M (1984) Samarium-neodymium data on two late Proterozoic ophiolites of Saudi Arabia and implications for crustal and mantle evolution. Contrib Mineral Petrol 85:244–252 Collins AS, Pisarevsky SA (2005) Amalgamating eastern Gondwana: the evolution of the Circum-Indian Orogens. Earth Sci Rev 71:229– 270 Dalziel IWD (1991) Pacific margins of Laurentia and East Antarctica as a conjugate rift pair: evidence and implications for an Eocambrian supercontinent. Geology 19:598–601 Dalziel IWD (1997) Neoproterozoic—Paleozoic geography and tectonics: review, hypothesis and environmental speculation. Geol Soc Am Bull 109:16–42 Dessouky OK, Weidong S, Ibrahim WS, Hani HA, Hassan M, Congying Li, Omran AA, Dardier AM (2020) New insights into Cryogenian Arc granitoids hosting Th–U mineralized Ediacaran syenogranite dikes, Ra’s Abdah area in the northern Egyptian Nubian Shield: Constraints from U–Pb Ages and zircon geochemistry. Precam Res https://doi.org/10.1016/j.precamres.2020.105986 Dilek Y, Ahmed Z (2003) Proterozoic ophiolites of the Arabian Shield and their significance in Precambrian tectonics. Ophiolites in Earth History: Geological Society, London, Special Publications 218:685–701 Direen NG, Crawford AJ (2003) The Tasman Line: where is it, what is it, and is it Australia’s Rodinian breakup boundary? Australian J Earth Sci 50:491–502 Dixon TH (1981) Gebel dahanib, a late Precambrian layered sill of komatiitic composition. Contrib Mineral Petrol 76:42–52

M. Hassan et al. Eisbacher GH (1985) Pericollisional strike-slip faults and synorogenic basins, Canadian Cordillera. In: Biddle KT, Christie-Blick N (eds) Strike slip deformation. basin formation and sedimentation. Soc Econ Paleontol Mineral Special Publications 37:265–282 El Akhal H (1993) Transect of a tectonic melange to an island-arc in the Pan-African of SE Egypt (Wadi Ghadir area). Scientific series of the International Bureau/Forschungszentrum Jülich GmbH, p 20 El Bahariya GA (2021) The ophiolite-dominated suprastructure eastern desert, Egypt. In: Hamimi Z, Arai S, Fowler AR, El-Bialy MZ (eds) The geology of the Egyptian nubian shield. Regional Geology Reviews. Springer, Cham., Heidelberg, pp 83–106. https://doi.org/ 10.1007/978-3-030-49771-2_6 El-Bayoumi RM (1983) Ophiolites and mélange complex of Wadi Ghadir Area, Eastern Desert, Egypt. Bulletin of the Faculty of Earth Sciences, King Abdulaziz University, Jiddah 6:329–342 El-Gaby S, Greiling RO (1988) The Pan-African belt of Northeast Africa and adjacent areas. Tectonic evolution and economic aspects of a late proterozoic Orogen. Vieweg and Sohn, Braunschweig/ Wiesbaden, p 276 El-Sayed MM, Furnes H, Mohamed FH (1999) Geochemical constraints on the tectonomagmatic evolution of the late Precambrian Fawakhir ophiolite, central eastern desert, Egypt. J Afr Earth Sci 29:515–533 Evans DAD (2009) The palaeomagnetically viable, long-lived and all-inclusive Rodinia supercontinent reconstruction. In: Murphy JB, Keppie JD, Hynes AJ (eds) Ancient orogens and modern analogues. Geological Society, London, Special Publications, 327:371–404 Eyal M, Be’eri-Shlevin Y, Eyal Y, Whitehouse MJ, Litvinovsky B (2014) Three successive Proterozoic island arcs in the Northern Arabian-Nubian Shield: evidence from SIMS U–Pb dating of zircon. Gondwana Res 25:338–357 Farahat ES (2010) Neoproterozoic arc–back-arc system in the central eastern desert of Egypt: Evidence from supra-subduction zone ophiolites. Lithos 120:293–308 Farahat ES, El Mahalawi MM, Holinkes G, Abdel Aal AY, Hauzenberger CA (2010) Reply to GA El Bahariya’s comments on: “Pillow form morphology of selected Neoproterozoic metavolcanics in the Egyptian central eastern desert and their implications”. J Afr Earth Sci 57:163–168 Fioretti AM, Black LP, Foden J, Visona D (2005) Grenville-age magmatism at the South Tasman Rise (Australia): a new piercing point for the reconstruction of Rodinia. Geology 33:769–772 Fitches WR, Graham RH, Hussein IM, Ries AC, Shackleton RM, Price RC (1983) The late proterozoic ophiolite of Sol Hamed, NE Sudan. Precambr Res 9:385–411 Fitzsimons ICW, Hulscher B (2005) Out of Africa: detrital zircon provenance of central Madagascar and Neoproterozoic terrane transfer across the Mozambique Ocean. Terra Nova 17:224–235 Fleck RJ (1985) Age of diorite-granodiorite gneisses of the Jiddah-Makkah region, Kingdom of Saudi Arabia. Saudi Arabian Deputy Ministry for Mineral Resources Professional Paper PP2, pp 19–27 Fleck RJ, Hadley DG (1982) Ages and strontium initial ratios of plutonic rocks in a transect of the Arabian Shield, U.S. Geological Survey Open File Report USGS-OF-03-38, p 43 Fowler A, Hassan I (2008) Extensional tectonic origin of gneissosity and related structures of the Feiran-Solaf metamorphic belt, Sinai, Egypt. Precambr Res 164:119–136 Fowler A, Hassen IS, Hassan M (2015) Tectonic evolution and setting of the Sa’al complex, southern Sinai, Egypt: a Proterozoic continental back-arc rift model. J Afr Earth Sci 104:103–131 Fowler A, Hassen IS, Hassan M (2018) The Feiran-Solaf metamorphic complex, Sinai, Egypt: evidence for orthogonal or oblique tectonic convergence? J Afr Earth Sci 146:48–65

9

Early Ensimatic Stage of the Arabian-Nubian Shield

Frisch W, Al-Shanti AM (1977) Ophiolite belts and the collision of island arcs in the Arabian Shield. Tectonophysics 43:293–306 Fritz H, Abdelsalam M, Ali KA, Bingen B, Collins AS, Fowler AR, Ghebreab W, Hauzenberger CA, Johnson PR, Kusky TM, Macey P, Muhongo S, Stern RJ, Viola G (2013) Orogen styles in the East African Orogen: a review of the Neoproterozoic to Cambrian tectonic evolution. J Afr Earth Sci 86:65–106 Gahlan HA, Azer MK, Khalil AES (2015) The Neoproterozoic Abu Dahr ophiolite, South Eastern Desert, Egypt: petrological characteristics and tectonomagmatic evolution. Mineral Petrol 109:611– 630 Gamal El Dien H, Hamdy M, Abu El-Ela A, Abu-Alam T, Hassan A, Kil Y, Mizukami T, Soda Y (2016) Neoproterozoic serpentinites from the Eastern Desert of Egypt: insights into Neoproterozoic mantle geodynamics and processes beneath the ANS. Precambr Res 286:213–233 Gass IG (1977) The evolution of the Pan-African crystalline basement in NE Africa and Arabia. J Geol Soc Lond 134:129–138 Genna A, Nehlig P, Le Goff E, Guerrot C, Shanti M (2002) Proterozoic tectonism of the Arabian shield. Precambr Res 117:21–40 Goodge JW, Vervoort JD, Fanning CM, Brecke DM, Farmer GL, Williams IS, Myrow PM, DePaolo DJ (2008) A positive test of East Antactica-Laurentia juxtaposition within the Rodinia Supercontinent. Science 321:235–240 Hamimi Z, Kassem OMK, El-Sabrouty MN (2015) Application of kinematic vorticity techniques for mylonitized rocks in Al Amar suture, Eastern Arabian Shield, Saudi Arabia. Geotectonics 49 (5):439–450 Hamimi Z, Arai S, Fowler AR, El-Bialy MZ (2021) The geology of the Egyptian Nubian Shield. Regional Geology Reviews, Springer Nature Switzerland, p. 700 Hargrove US, Stern RJ, Kimura J-I, Manton WI, Johnson PR (2006) How juvenile is the ANS? Evidence from Ndisotopes and pre-Neoproterozoic inherited zircon in the Bi’r Umq suture zone, Saudi Arabia. Earth Planet Sci Lett 252(3–4):308–326 Hargrove US, Stern RJ, Griffin WR, Johnson PR, Abdelsalam MG (2006) From island arc to craton: timescales of crustal formation along the Neoproterozoic Bi’rUmq suture zone, Kingdom of Saudi Arabia. Technical Report SGS-TR-2006-6. Saudi Geological Survey, 69 pp Harlan SS, Geissman JW, Premo WR (2003a) Paleomagnetism and geochronology of an early Proterozoic quartz diorite in the southern Rind River Range, Wyoming USA. Tectonophysics 362:105–122 Harlan SS, Heaman L, LeCheminant AN, Premo WR (2003b) Gunbarrel mafic magmatic event: a key 780 Ma time marker for Rodinia plate reconstructions. Geology 31:1053–1056 Harms U, Schandelmeier H, Darbyshire DPE (1990) Pan-African reworked early/middle Proterozoic crust in NE Africa west of the Nile: Sr and Nd isotope evidence. J Geol Soc Lond 147:859–872 Hassan M, Abu-Alam TS, Stüwe K, Fowler A, Hassen IS (2014) Metamorphic evolution of the Sa’al–Zaghra Complex in Sinai: evidence for Mesoproterozoic Rodinia break-up? Precambr Res 241:104–128 Hassan M, Abu-Alam T, Fowler AR (2021) The Sinai metamorphic complexes from Rodinia rifting to the Gondwana Collision. In: Hamimi Z, Arai S, Fowler AR, El-Bialy MZ (eds) The geology of the Egyptian Nubian shield. Regional Geology Reviews. Springer, Cham, Heidelberg, pp 83–106. https://doi.org/10.1007/978-3-03049771-2_4 Hoffman PF (1991) Did the breakout of Laurentia turn Gondwanaland inside out? Science 252:1409–1412 Hoffman PF (1999) The break-up of Rodinia, birth of Gondwana, true polar wander and the snowball Earth. J Afr Earth Sci 28:17–33 Hussein IM (2000) Geodynamic evolution of the Pan-African crystalline basement in the northern Red Sea Hills, Sudan, with special

217 emphasis on the Wadi Onib ophiolite and geology to the west of Port Sudan. Dissertation, Johannes Gutenberg Universität-Mainz, Mainz Johansson A (2009) Baltica, Amazonia and the SAMBA connection— 1000 million years of neighbourhood during the Proterozoic? Precambr Res 175(1–4):221–234 Johnson PR, Andresen A, Collins AS, Fowler AR, Fritz H, Ghebreab W, Kusky T, Stern RJ (2011) Late Cryogenian-Ediacaran history of the Arabian-NubianShield: A review of depositional, plutonic, structural, and tectonic events in the closing stages of the northern East African Orogen. J Afr Earth Sci 61:167–232 Johansson A (2014) From Rodinia to Gondwana with the ‘SAMBA’model—a distant view from Baltica towards Amazonia and beyond. Precambr Res 244:226–235 Johnson PR (2014) An expanding Arabian-Nubian Shield geochronologic and isotopic dataset: defining limits and confirming the tectonic setting of a Neoproterozoic accretionary orogen. Open Geol J 8:3–33 Johnson PR, Kattan FH (2001) Oblique sinistral transpression in the Arabian shield the timing and kinematics of a Neoproterozoic suture zone. Precambr Res 107:117–138 Johnson PR, Woldehaimanot B (2003) Development of the ANS. Perspectives on accretion and deformation in the northern East African Orogen and the assembly of Gondwana. In: Yoshida M, Windley BF, Dasgupta S (eds) Proterozoic East Gondwana, supercontinent assembly and breakup. Geological Society, London, Special Publications, 206:290–325 Johnson PR, Kattan FH, Al-Saleh AM (2004) Neoproterozoic ophiolites in the Arabian Shield: Field relations and structure. In: Kusky TM (ed) Precambrian ophiolites and related rocks. Developments in Precambrian Geology, vol 13, pp 129–162. Elsevier, Amsterdam Jöns N, Schenk V (2008) Relics of the Mozambique ocean in the central East African Orogen: evidence from the Vohibory Block of southern Madagascar. J Metam Geol 26:17–28 Kröner A (1985) Ophiolites and the evolution of tectonic boundaries in the late Proterozoic ANS of northeastern Africa and Arabia. Precambr Res 27:277–300 Kröner A, Linnebacher P, Stern RJ, Reischmann T, Manton WI, Hussein IM (1991) Evolution of Pan-African island-arc assemblages in the southern Red Sea Hills, Sudan and in southwestern Arabia as exemplified by geochemistry and geochronology. Precambr Res 53:99–118 Kröner A, Todt W, Hussein IM, Mansour M, Rashwan AA (1992) Dating of late Proterozoic ophiolites in Egypt and the Sudan using the single zircon evaporation technique. Precambr Res 59:15–32 Küster D, Liégeois J-P (2001) Sr, Nd isotopes and geochemistry of the Bayuda Desert high-grade metamorphic basement (Sudan): an early Pan-African oceanic convergent margin, not the edge of the East Saharan ghost craton? Precambr Res 109:1–23 Küster D, Liégeois JP, Matukov D, Sergeev S, Lucassen F (2008) Zircon geochronology and Sr, Nd, Pb isotope geochemistry of granitoids from Bayuda Desert andSabaloka (Sudan): Evidence for a Bayudian event (920–900 Ma) preceding the Pan-African orogenic cycle (860-590 Ma) at the eastern Boundary of the SaharanMetacraton: Precambr Res 164:16–39 Li ZX, Zhang L, Powell CM (1995) South China in Rodinia: part of the missing link between Australia-East Antarctica and Laurentia? Geology 23:407–410 Li ZX, Li XH, Kinny PD, Wang J (1999) The breakup of Rodinia: did it start with a mantle plume beneath South China? Earth Planet Sci Lett 173:171–181 Li ZX, Powell CM (2001) An outline of the Palaeogeographic evolution of the Australasian region since the beginning of the Neoproterozoic. Earth Sci Rev 53:237–277

218 Li XH, Li ZX, Zhou H, Liu Y, Kinny PD (2002a) U–Pb zircon geochronology, geochemistry and Nd isotopic study of Neoproterozoic bimodal volcanic rocks in the Kangdian Rift of South China; implications for the initial rifting of Rodinia. Precambr Res 113:135–154 Li ZX, Li XH, Zhou H, Kinny PD (2002b) Grenvillian continental collision in South China: new SHRIMP U–Pb zircon results and implications for the configuration of Rodinia. Geology 30:163–166 Li XH, Li ZX, Ge WC, Zhou HW, Li WX, Liu Y, Wingate MTD (2003a) Neoproterozoic granitoids in South China: crustal melting above a mantle plume at ca 825 Ma? Precambr Res 122:45–83 Li XH, Li ZX, Zhou H, Liu Y, Liang X, Li W (2003b) SHRIMP U–Pb zircon age, geochemistry and Nd isotope of the Guandaoshan pluton in SW Sichuan: petrogenesis and tectonic significance. Sci China D 46 (Suppl):73–83 Li ZX, Li XH, Kinny PD, Wang J, Zhang S, Zhou H (2003c) Geochronology of Neoproterozoic syn-rift magmatism in the Yangtze Craton South China and correlations with other continents: evidence for a mantle superplume that broke up Rodinia. Precambr Res 122:85–109 Li ZX, Bogdanova SV, Collins AS, Davidson A, De Waele B, Ernst RE, Fitzsimons ICW, Fuck RA, Gladkochub DP, Jacobs J, Karlstrom KE, Lu S, Natapov LM, Pease V, Pisarevsky SA, Thrane K, Vernikovsky V (2008) Assembly, configuration, and break-up history of Rodinia: a synthesis. Precambr Res 160:179– 210 Li XH, Abd El-Rahman Y, Abu Anbar M, Li J, Ling X-X, Wu L-G, Masoud AE (2018) Old continental crust underlying juvenile oceanic arc: evidence from northern Arabian-Nubian Shield, Egypt. Geophys Res Lett 45:3001–3008 Ling W, Gao S, Zhang B, Li H, Liu Y, Cheng J (2003) Neoproterozoic tectonic evolution of the northwestern Yangtze craton South China: implications for amalgamation and break-up of the Rodinia Supercontinent. Precambr Res 122:111–140 McMenamin MAS, McMenamin DLS (1990) The Emergence of Animals; The Cambrian Breakthrough. Columbia Univ. Press, New York, p 217 Meert JG (2003) A synopsis of events related to the assembly of eastern Gondwana. Tectonophysics 362:1–40 Meert JG, Torsvik TH (2003) The making and unmaking of a super-continent: Rodinia revisited. Tectonophysics 375:261–288 Meert JG, Van der Voo R (1997) The assembly of Gondwana 800–550 Ma. J Geodyn 23:223–235 Mitchell RN, Kilian TM, Evans DAD (2012) Supercontinent cycles and the calculation of absolute palaeolongitude in deep time. Nature 482:208–212 Moores EM (1991) Southwest U.S.—East Antarctic (SWEAT) connection: a hypothesis. Geology 19:425–428 Murphy JB, Nance RD (2003) Do supercontinents introvert or extrovert? Sm–Nd isotopic evidence. Geology 31:873–876 Nance RD, Worsley TR, Moody JB (1986) Post-Archean biogeochemical cycles and long-term episodicity in tectonic processes. Geology 14:514–518 Nance RD, Murphy JB, Santosh M (2014) The supercontinent cycle: a retrospective essay. Gondwana Res 25:4–29 Nassief MO, Macdonald R, Gass IG (1984) The Jebel Thurwah upper Proterozoic ophiolite complex, western Saudi Arabia. J Geol Soc Lond 141:537–546 Neary CR, Gass IG, Cavanagh BJ (1976) Granitic association of northeastern Sudan. Geol Soc Am Bull 87:1501–1512 Obeid MA, Khalil AES, Azer MK (2016) Mineralogy, geochemistry, and geotectonic significance of the Neoproterozoic ophiolite of Wadi Arais area, south Eastern Desert, Egypt. Int Geol Rev 58:687–702 Pallister JS, Stacey JS, Fischer LB, Premo WR (1988) Precambrian ophiolites of Arabia: geologic settings, U–Pb geochronology, Pb–

M. Hassan et al. isotope characteristics, and implications for continental accretion. Precambr Res 38:1–54 Perelló J, Sillitoe RH, Brockway H, García A (2020) Metallogenic inception of the Arabian-Nubian Shield: Daero Paulos porphyry copper prospect, Eritrea. Gondwana Res 88:106–125 Pesonen LJ, Elming S-A, Mertanen S, Pisarevsky S, D’Agrella-Filho MS, Meert JG, Schmidt PW, Abrahamsen N, Bylund G (2003) Palaeomagnetic configuration of continents during the Proterozoic. Tectonophysics 375:289–324 Pesonen LJ, Mertanen S, Veikkolainen T (2012) PaleoMesoproterozoic supercontinents—a paleomagnetic view. Geophysica 47(1–2):3–46 Piper JDA (1976) Palaeomagnetic evidence for a Proterozoic supercontinent. Philos Trans Res Soc Lond 280:469–490 Pisarevsky SA, Wingate MTD, Powell CMcA, Johnson SP, Evans DAD (2003) Models of Rodinia assembly and fragmentation. In: Yoshida M, Windley BF, Dasgupta S (eds) Proterozoic East Gondwana: Supercontinent Assembly and Break-up. Geological Society, London, Special Publications, 206:35–55 Powell CM, Li ZX, McElhinny MW, Meert JG, Park JK (1993) Paleomagnetic constraints on timing of the Neoproterozoic breakup of Rodinia and the Cambrian formation of Gondwana; with Suppl Data 9335. Geology 21:889–892 Powell CM, Preiss WV, Gatehouse CG, Krapez B, Li ZX (1994) South Australian record of a Rodinian epicontinental basin and its mid-Neoproterozoic breakup (*700 Ma) to form the palaeo-Pacifi c Ocean: Tectonophysics 237:113–140 Preiss WV (2000) The adelaide geosyncline of South Australia and its significance in Neoproterozoic continental reconstruction. Precambr Res 100:21–63 Price RC (1984) Late precambrian mafic-ultramafic complexes in Northeast Africa. PhD thesis. Open University, Milton Keynes, p 325 Reischmann T, Bachtadse V, Kröner A, Layer P (1992) Geochronology and paleomagnetism of a late Proterozoic island-arc terrane from the Red Sea Hills, northeast Sudan. Earth Planet Sci Lett 114:1–15 Ross GM, Villeneuve M (2003) Provenance of the Mesoproterozoic (1.45 Ga) Belt basin (western North America): another piece of the pre-Rodinia paleogeographic puzzle. Geol Soc Am Bull 115:1191– 1217 Ross GM, Parrish RR, Winston D (1992) Provenance and U-Pb geochronology of the Mesoproterozoic Belt Supergroup (northwestern United States): implications for age of deposition and pre-Panthalassa plate reconstructions. Earth Planet Sci Lett 113:57– 76 Schmidt PW, Williams GE, Camacho A, Lee JKW (2006) Assembly of Proterozoic Australia: implications of a revised pole for the similar to 1070 Ma Alcurra Dyke Swarm, central Australia. Geophys J Int 167:626–634 Scotese CR (2009) Late Proterozoic plate tectonics and palaeogeography: a tale of two supercontinents, Rodinia and Pannotia. J Geol Soc Lond Spec Publ 326:67–83 Shackleton RM (1988) Contrasting structural relationships of Proterozoic ophiolites in Northeast and Eastern Africa. In: El-Gaby S, Greiling RO (eds) The Pan-African belt of Northeast African and adjacent areas. Friedrich Viewig and Sohn, Braunschweig/ Wiesbaden, pp 183–193 Shanti M (1983) The Jabal Ess ophiolite complex. Bulletin of the Faculty of Earth Sciences, King Abdulaziz University, Jiddah 6:289–317 Shanti M, Roobol MJ (1979) A late Proterozoic ophiolite complex at Jabal Ess in northern Saudi Arabia. Nature 279:488–491 Stacey JS, Agar RA (1985) U–Pb isotopic evidence for the accretion of a continental micro-plate in the Zalm region of the Saudi Arabian Shield. J Geol Soc Lond 142:1189–1203

9

Early Ensimatic Stage of the Arabian-Nubian Shield

Stacey JS, Hedge CE (1984) Geochronologic and isotopic evidence for early Proterozoic crust in the eastern Arabian shield. Geology 12:310–313 Stacey JS, Stoeser DB (1983) Distribution of continental and oceanic leads in the ANS. Contrib Mineral Petrol 84:91–105 Stacey JS, Doe BR, Roberts RJ, Delevaux MH, Gramlich JW (1980) A lead isotope study of mineralization in the Saudi Arabian shield. Contrib Mineral Petrol 74:175–188 Stacey JS, Stoeser DB, Greenwood WR, Fischer LB (1984) U-Pb zircon geochronology and geological evolution of the Halaban-Al Amar region of the Eastern Arabian Shield, Kingdom of Saudi Arabia. J Geol Soc Lond 141:1043–1055 Stein M (2003) Tracing the plume material in the ANS. Precambr Res 123:223–234 Stein M, Goldstein SL (1996) From plume head to continental lithosphere in the ANS. Nature 382:773–778 Stern RJ (1994) Arc assembly and continental collision in the Neoproterozoic East African Orogen: implications for consolidation of Gondwanaland. Annu Rev Earth Planet Sci 22:319–354 Stern RJ (2004) Subduction initiation: spontaneous and induced. Earth Planet Sci Lett 226:275–292 Stern RJ, Johnson PR, Kröner A, Yibas B (2004) Neoproterozoic ophiolites of the ANS. In: Kusky TM (ed) Precambrian ophiolites and related rocks. Developments in Precambrian Geology, vol 13, pp 95–128. Elsevier, Amsterdam Stern RJ, Avigad D, Miller NR, Beyth M (2006) Evidence for the Snowball earth hypothesis in the ANS and the East African Orogen. J Afr Earth Sci 44:1–20 Stern RJ, Ali KA, Liégeois JP, Johnson PR, Kozdroj W, Kattan FH (2010) Distribution and significance of pre-Neoproterozoic zircons in juvenile Neoproterozoic igneous rocks of the ANS. Am J Sci 310:791–811 Stoeser DB, Stacey JS (1988) Evolution, U–Pb geochronology, and isotope geology of the Pan-African Nabitah orogenic belt of the Saudi Arabian Shield. In: El-Gaby S, Greiling RO (eds) The Pan-African belt of Northeast African and adjacent areas. Friedrich Viewig and Sohn, Braunschweig/Wiesbaden, pp 227–288 Stoeser DB, Whitehouse MJ, Stacey JS (2001) The Khida terrane-geology of Paleoproterozoic rocks in the Muhayil area, eastern Arabian Shield, Saudi Arabia. Gondwana Res 4:192–194 Stoeser DB, Whitehouse MJ, Stacey JS (2004) Neoproterozoic evolution of the Khida terrane, Saudi Arabia: a detached microplate in the Arabian craton. In 32nd IGC, Florence (Abstracts) Sultan M, Arvidson RE, Duncan I, Stern RJ, El Kaliouby M (1988) Extension of the Najd shear system from Saudi Arabia to the Central Eastern Desert of Egypt based on integrated field and landsat observations. Tectonics 7:1291–1306 Sultan M, Chamberlain KR, Bowring SA, Arvidson RE, Abuzied H, El Kaliouby B (1990) Geochronological and isotopic evidence for involvement of pre-Pan-African crust in the Nubian shield. Geology 18:761–764 Sultan M, Tucker RD, El Alfy Z, Attia R (1994) U-Pb (zircon) ages for the gneissic terrane west of the Nile, southern Egypt. Geologisch Rundschau 83:514–522 Sun SS, Sheraton JW (1996) Geochemical and isotopic evolution. In: Glikson AY et al (eds) Geology of the western Musgrave block, Central Australia, with particular reference to the mafic-ultramafic Giles complex. Australian Geological Survey Organization (AGSO) Bulletin 239:135–143 Takla MA, Basta FF, El Maghraby AMO, Griffin NL (2002) Geology and geochemistry of ophiolites from the Nubian Shield, North East Africa, Egypt. In: International conference geology. The Arab World, Cairo University, Egypt (GAW 6), pp 1–24

219 Tohver E, D’Agrella-Filho MS, Trindade RIF (2006) Paleomagnetic record of Africa and South America for the 1200–500 Ma interval, and evaluation of Rodinia and Gondwana assemblies. Precambr Res 147:193–222 Torsvik TH, Smethurst MA, Meert JG, Van der Voo R, McKerrow WS, Sturt BA, Brasier MD, Walderhaug HJ (1996) Continental breakup and collision in the Neoproterozoic and Paleozoic-a tale of Baltica and Laurentia. Earth Sci Rev 40:229–258 Tucker RD, Roig J-Y, Delor C, Amelin Y, Goncalves P, Rabarimanana MH, Ralison AV, Belcher RW (2011) Neoproterozoic extension in the Greater Dharwar Craton: a reevaluation of the “Betsimisaraka suture” in Madagascar. Canadian J Earth Sci 48 (2):389–417 Valentine JW, Moores EM (1970) Plate-tectonic regulation of animal diversity and sea level: a model. Nature 228:657–659 Van der Voo R, Peinado J, Scotese CR (1984) Was Laurentia part of an Eocambrian supercontinent? Geodynamics Ser 12:131–136 Wang J, Li ZX (2003) History of Neoproterozoic rift basins in South China: implications for Rodinia break-up. Precambr Res 122:141– 158 Wang X-C, Wilde SA, Li Z-X, Li S, Li L (2020) Do Supercontinent-superplume cycles control the growth and evolution of continental crust? J Earth Sci. https://doi.org/10.1007/s12583020-1077-4 Weil AB, Van der Voo R, Mac Niocaill C, Meert JG (1998) The Proterozoic supercontinent Rodinia: paleomagnetically derived reconstructions for 1,100 to 800 Ma. Earth Planet Sci Lett 154:13–24 Spencer CJ, Cawood PA, Hawkesworth CJ, Prave AR, Roberts NMW, Horstwood MSA, Whitehouse MJ, EIMF (2015) Generation and preservation of continental crust in the Grenville Orogeny. Geosci Front 6:357–372 Whitehouse MJ, Windley BF, Ba-Bttat MAO, Fanning CM, Rex DC (1998) Crustal evolution and terrane correlation in the eastern Arabian Shield, Yemen: geochronological constraints. J Geol Soc Lond 155:281–295 Whitehouse MJ, Stoeser DB, Stacey JS (2001) The Khida terrane— geochronological and isotopic evidence for Palaeoproterozoic and Archaean crust in the Eastern Arabian Shield of Saudi Arabia. Gondwana Res 4:200–202 Windley BF, Whitehouse MJ, Ba-Bttat MAO (1996) Early Precambrian gneiss terranes and Pan-African island arcs in Yemen: crustal accretion of the eastern Arabian Shield. Geology 24:131–134 Wingate MTD, Campbell IH, Compston W, Gibson GM (1998) Ion microprobe U–Pb ages for Neoproterozoic basaltic magmatism in south-central Australia and implications for the breakup of Rodinia. Precambr Res 87(3–4):135–159 Wingate MTD, Pisarevsky SA, Evans DAD (2002) Rodinia connections between Australia and Laurentia: no SWEAT, no AUSWUS? Terra Nova 14:121–128 Wolde B, Asres Z, Desta Z, Gonzalez JJ (1996) Neoproterozoic zirconium-depleted boninite and tholeiitic series rocks from Adola, southern Ethiopia. Precambr Res 80:261–279 Worku H (1996) Geodynamic development of the Adola Belt (Southern Ethiopia) in the Neopro-terozoic and Its Control on Gold Mineralization. Verlag Dr, Köster, Berlin, p 156 Zimmer M, Kröner A, Jochum KP, Reischmann T, Todt W (1995) The Gabal Gerf complex: a Precambrian N-MORB ophiolite in the Nubian Shield, NE Africa. Chem Geol 123:29–51 Zoheir BA, Klemm DD (2007) The tectono-metamorphic evolution of the central part of the Neoproterozoic Allaqi-Heiani suture, south Eastern Desert of Egypt. Gondwana Res 12:289–304

Terrane Accretion Within the Arabian-Nubian Shield

10

Ali Farrag Osman and Abdel-Rahman Fowler

Abstract

Accretionary orogens result from plate tectonic processes at subduction zones. These orogens represent significant contributions to crustal growth through Earth history and are the most important factories for generating, recycling, and maturing continental crust. Most accreted terranes are autochthonous, i.e., which are sedimentary or volcanic tectonic elements developed near to the site of terrane collision, such as accretionary prisms, fore-arc basins, and melange bodies. Terrane studies provide important evidence for the migration of parts of the Earth’s crust from one place to another distant place. Terranes transported long distances to the site of accretion are allochthonous, or ‘suspect’ or ‘exotic’ terranes. The style of terrane accretion may involve obduction, overthrusting, subduction, underplating, or simple suturing. Accretion commonly results in crustal thickening and uplift, and trench migration. Three orogenies were key to the assembly of Gondwana: The East African Orogeny (800–650 Ma), the Kuunga Orogeny (c. 550 Ma), and the Brasiliano Orogeny (660–530 Ma). Collectively, these have been referred to as the Pan-African Orogeny (PAO). The PAO comprised a series of deformation belts, including the ANS and the Mozambique Belt (MB), the latter extending from East Antarctica through East Africa, up to the ANS. The PAO and Grenville Orogenies are the largest known orogenic systems on Earth. The East African Orogen marks the collision zone of East and West Gondwana. In its southern parts, it records oblique continent–continent collision and high-grade metamorphic reworking of much older crust. In its northern part, in the ANS, it records A. F. Osman (&) Faculty of Science, Geology Department, Ain Shams University, Cairo, Egypt e-mail: [email protected] A.-R. Fowler Faculty of Science, Geology Department, United Arab Emirates University, Al Ain, United Arab Emirates

stacking of juvenile arc terranes, followed by tectonic escape.

10.1

Introduction

The Arabian-Nubian Shield (ANS) is a classic example of an accretionary orogen, involving terrane amalgamation and accretion during subduction, leading to the progressive formation of superterranes and proto-continents, and finally achieving continental growth. The ANS is the result of a 300 million-year period of crustal growth, involving the accretion of juvenile subduction-related magmatic arcs, periods of sedimentation, and volcanism in basins developed on newly post-amalgamated arc terranes, crustal thickening, the emplacement of large amounts of granitic magma, periodic uplift and erosion, extension, and tectonic escape on transcurrent faults. The ANS assembly involved numerous events: post-amalgamation sedimentation and volcanism, crustal thickening, granite magmatism, uplift and erosion, tectonic extension and escape, and transcurrent faulting. The first stage of ANS began with its southern part, including areas south of the Nakasib-Bir Umq suture and from Nakasib suture to the Yanbu-Allaqi-Heiani suture. These arc terranes collide between 830 and 710 Ma. The second stage involved formation of the Midyan—Eastern Desert terrane by collisions in the range 760–730 Ma, to form the ‘western arc terranes.’ The third stage witnessed the collision of the western arc terranes with the earlier amalgamated terranes to form the proto-ANS, an arc superterrane, at 680–640 Ma. The Nabitah Orogeny marks the climax of this stage. The fourth and final stage involved the accretion of the proto-ANS with the Saharan craton along the Keraf suture, as a result of collisions in the east of the ANS of the Ad Dawadmi and Ar Rayn terranes (650–542 Ma). Recognition of ophiolites introduced the application of plate tectonic theory to the evolution of the ANS. Models for the ANS

© The Author(s), under exclusive license to Springer Nature Switzerland AG 2021 Z. Hamimi et al. (eds.), The Geology of the Arabian-Nubian Shield, Regional Geology Reviews, https://doi.org/10.1007/978-3-030-72995-0_10

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include (1) single immature island arc, (2) successive accretion of island arcs, (3) repeated rifting of old sialic crust, and 4) microplate accretion. A continental margin model for ANS evolution was popular, involving a pre-Pan-African sialic infrastructure, over which ophiolites and arc volcanics were thrusted (a suprastructure). Following emplacement of the suprastructure, continuing subduction resulted in Andean-type magmatism. A more recent model of originally entirely ensimatic setting prior to cratonization is better supported by current data. The ophiolite-decorated terrane boundaries within the ANS have encouraged interpretation of these as sutures. However, there is a rarity of Penrose-type pseudostratigraphy, suggesting that these may not be true ophiolites. Many have the chemical characteristics of fore-arc oceanic crust. The sutures can be divided into arc–arc sutures (Allaqi-Heiani-Yanbu; Nakasib; Baraka-Tulu Dimtu; Adola-Moyale plus perhaps some ophiolitic mélange in the Eastern Desert) and arc–continent sutures (Al Amar; Keraf; Sekerr) Post-accretion structures include N–S to NW–SE trending intense strain zones (Hamisana and Oko) that have also been originally interpreted as shear zones. The timing of development of the Hamisana and Oko shear zone is poorly constrained to date somewhere between 700 and 560 Ma. In both zones, the early structures are N-S to NNW-trending upright tight folds, followed by NW–SE trending shear zones. Fault and shear zone systems in the ANS are mainly N-S oriented in the southern parts and NW–SE oriented (with a minor NE-trending set) in the northern parts. The NW-trending sinistral Najd Fault System (NFS) is associated with elongated gneiss domes and pull-apart sedimentary basins. The Najd activity has been divided into a 630– 600 Ma stage of dextral shear, followed by a 600–530 Ma sinistral strike-slip shear history with *240 km sinistral displacement estimates. The origins of the NFS remain somewhat controversial with indentation tectonic and continental transform interpretations proposed. Postamalgamation volcano-sedimentary basins also developed in the ANS. These basins lie unconformably upon the suture zones. They are intruded by large volumes of late- to post-orogenic granitoids. Post-accretion structures include N-S to NW-SE trending intense strain zones (Hamisana and Oko) that have also been originally interpreted as shear zones. The timing of development of the Hamisana and Oko shear zone is poorly constrained to date somewhere between 700 and 560 Ma. In both zones the early structures are N-S to NNW-trending upright tight folds, followed by NW-SE trending shear zones. Fault and shear zone systems in the ANS are mainly N-S oriented in the southern parts and NW-SE oriented (with a minor NE-trending set) in the northern parts. The NW-trending sinistral Najd Fault System (NFS) is associated with elongated gneiss domes and pull-apart sedimentary basins. The Najd activity has been divided into a 630–

A. F. Osman and A.-R. Fowler

600 Ma stage of dextral shear, followed by a 600–530 Ma sinistral strike-slip shear history with *240 km sinistral displacement estimates. The origins of the NFS remain somewhat controversial with indentation tectonic and continental transform interpretations proposed. Post-amalgamation volcano-sedimentary basins were also developed in the ANS. These basins lie unconformably in the suture zones. They are intruded by large volumes of lateto post-orogenic granitoids.

10.2

The Rationale

One of the most significant mechanisms of crustal growth is the development of accretionary orogens at oceanic lithospheric subduction sites (Condie 2002; Foster et al. 2009). The accretionary orogens consist of magmatic-arc systems along with material accreted from the downgoing plate and eroded from the upper plate. These orogens form at intraoceanic and continental margin convergent plate boundaries. They include the supra-subduction zone fore-arc, magmatic arc, and back-arc components (Cawood et al. 2009). The ANS is considered as one of the classic examples of accretionary orogens. Accretionary orogenesis has played a significant role in the growth of continental crust (Stern and Scholl 2010; Clift et al. 2009; Cawood et al. 2009). Accreted terranes can be defined as either autochthonous or allochthonous. Autochthonous terranes comprise most of terrane accretion and are composed of accretionary prism, mélange, and ophiolite units. Allochthonous terranes, also called suspect terranes (Howell 1985), were exotic (i.e., distant) crustal units on the subducting oceanic plate, that were transported to the subduction zone resulting in accretion. Present-day examples of accretionary orogenesis are recognized in the Pacific, where numerous oceanic plateaux and submarine ridges meet subduction zones along the margins of island arcs or continents (e.g., Solomon Islands, Izu Bonin—Japan, Roo Rise—Java, Ogasawara Plateau— Izu Bonin Mariana arc). Accretion of island arcs in such tectonic environments is a major contributor to continental crust growth (Stern and Scholl 2010; Clift et al. 2009; Cawood et al. 2009). The accretion of crustal units occurs when there is a weak detachment layer within the allochthonous terranes, such as island arcs, oceanic plateaus, submarine ridges, and continental fragments (Tetreault and Buiter 2012). The depth of detachment controls the amount of crust accreted onto the overriding plate, and lithospheric buoyancy does not prevent allochthonous terranes subduction during constant convergence. Island arcs, oceanic plateaus, and continental fragments will completely subduct, despite being composed of buoyant lithosphere, if they have rheologically strong crusts. Weak basal layers, representing pre-existing weaknesses or detachment layers will either lead

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to underplating of faulted blocks of allochthonous terrane crust to the overriding plate, or collision and suturing of an unbroken crust. The different types of accretionary processes also affect deformation and uplift patterns in the overriding plate, trench migration and jumping, and the dip of the plate interface. Recent geological and geophysical studies have shown that much of the crust of the ANS has grown through the accretion of discrete tectono-stratigraphic terranes. It is nowadays universally accepted that orogenic belts comprise diverse assemblages of rocks of different ages and tectonic settings that are juxtaposed because of orthogonal or oblique convergence along subduction zones, or because of lateral transport along major transcurrent faults (Johnson 2012). Such packages of rocks have different geologic histories and developed to a lesser or greater degree independently of each other. Their present juxtaposition is the result of orogenic processes, not the result of original depositional or intrusive relationships. Such packages are referred to as tectonostratigraphic terranes.

10.2.1 Terrane Terminology A terrane is a fragment of crustal material formed on, or broken off from, one tectonic plate and accreted or ‘sutured’ to crust lying on another plate. The concept of accreted terranes was first introduced in the 1970s (Monger et al. 1972; Irwin 1972; Coney 1978, 1980; Jones et al. 1982; Snoke and Barnes 2006). Irwin (1972) was the first to define the term ‘terrane’ in the geologic lexicon as ‘an association of geologic features, such as stratigraphic formations, intrusive rocks, mineral deposits, and tectonic history, some or all of which lend a distinctive character to a particular tract of rocks and which differ from those of an adjacent terrane.’ Terranes were a mystery to geologists until plate tectonic theory revealed how pieces of the crust could be moved and added to a continent far from their origins. Terranes that originated in an oceanic setting, but later became part of continents, definitely point to accretion, and are key evidence for plate tectonic theory. Terrane accretion is most common at convergent plate boundaries, but it may be possible for a terrane to be brought from an exotic location along a transform plate boundary. It is also possible for a newly developed divergent plate boundary to rift a continent apart. Part of the continent may then drift away on a moving plate to become an accreted terrane on another continent. Terranes are separated from the rocks around them by major faults, including sutures. They range in size from a few square kilometers to thousands of square kilometers. The crustal block that forms the terrane preserves its own distinctive geologic history, which contrasts with that of the

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surrounding areas—hence the term ‘exotic’ terrane. Accreted terranes may also be called suspect or tectono-stratigraphic terranes (Howell et al. 1985; Howell 1989). An important characteristic of terranes is that their present spatial relations are not compatible with their inferred geologic histories. In the case of adjacent terranes possessing strata of the same age, it must be demonstrable that their respective geologic evolutions are different and incompatible, and there must be no intermediate lithofacies which could link the strata. The term ‘geologic history’ is genetic, involving differences in lithology and degrees of volcanism or magmatism, differences in metamorphism and structure, differences in tectonic settings, and, if the rocks possess fossils, differences in faunal contents. Individual terranes can be classified into three types: (1) Stratigraphic terranes: composed of coherent sequences that represent depositional environments of continental fragments, ocean, or continental margin basins, and/or volcanic arcs. Stratigraphic terranes are characterized by coherent sequences of strata, in which depositionally different relations between successive or adjacent lithologies can be demonstrated. If crystalline basement rock is within the terrane boundary, its characteristics are helpful in classifying stratigraphic terranes into four broad categories: (a) fragments of continents, (b) fragments of continental margins, (c) fragments of volcanic arcs, and (d) fragments of ocean basins (Howell 1993). Oceanic crust, island arc crust, and continental crust can make up accreted terranes. Sometimes ophiolite sequences (which represent pieces of oceanic lithosphere) accrete onto the margins of a continent. The ophiolite sequences include, in their complete form, from the highest levels down: oceanic sediments, pillow basalt, sheeted dikes, gabbro, layered gabbro, and ultramafic rocks at the base. (2) Disrupted terranes: characterized by blocks of heterogeneous lithology and age, set in a matrix of foliated sandstone or serpentinite. (3) Metamorphic terranes: represented by structural blocks with a regional penetrative metamorphic fabric that obscures and is more distinctive than the original lithotypes. (4) Fragments of volcanic-arc terranes: composed predominantly of volcanic rocks, plutonic roots of arcs, and the sedimentary debris derived from volcanoes.

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Instances where fragments of volcanic arcs, fore-arc basins, and subduction complexes occur along a continental margin in an orogenic belt do not prove either that these assemblages formed in their present position, or that they are components of the same subduction complex. Some aspects of a volcanic arc complex are represented in almost all continental margin orogenic belts. Fragments of ocean basins terranes are characterized by sequences of mafic and ultramafic rocks typical of oceanic crust, along with overlying deep-sea sedimentary deposits. Many sequences may have younger continental margin strata capping the ophiolitic sequences, reflecting either back-arc-basin settings or the translational history of the crust from a mid-ocean setting to a continental margin location prior to accretion. The word terrane, with its distinctive spelling, is technically a tectono-stratigraphic terrane. A group of terranes that are accreted together prior to arrival at their final destination is called a superterrane. The term ‘tectonostratigraphic’ derives from how terranes originate. A tectono-stratigraphic terrane has both structural (tectonic) and stratigraphic criteria—a fault-bounded package of strata that is genetically unrelated to the adjoining stratigraphic packages genetically distinct from the other tectonostratigraphic terranes (Howell et al. 1985). The faults bounding the terrane resulted from tectonic processes lead to accretion. Some mountain belts, such as the North American Cordillera, are characterized by collisions between a single continent and numerous exotic terranes, rather than between two continents (Moores and Twiss 1995). Terranes converge and join by the process of amalgamation and in some cases eventually accrete to older continental crust. Post-amalgamation basins and stitching granites are emplaced at various stages during this tectonic sequence and constrain the timing of terrane suturing. There is broad consensus that plate tectonics operated during the Neoproterozoic, resulting (in the case of the ANS and regions to the south along the EAO) in the creation of the world’s largest expanse of juvenile Neoproterozoic crust. Many studies suggest that terrane accretions have occurred along the margin of most of ANS geologic entities of regional extent (e.g., Cawood and others 2009; Stoeser and Camp 1985; Johnson and Woldehaimanot 2003, Fritz et al. 2013; Johnson 2014). The term ‘terrane analysis’ refers to the methods of properly identifying crustal blocks as allochthonous relative to each other, and the investigations of the tectonic process of accretion (Jones et al. 1983; Howell 1995). Once terrane maps have been drawn, it is possible to construct plate tectonic models with regards to location and polarity of subduction zones and the formation of fore-arc, main arc, and back-arc crustal units.

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10.2.2 Formation of Terranes Terranes may vary enormously in size, from subcontinental dimension to block-size. The shapes of terranes and the patterns of their distribution reflect one, or a combination, of several tectonic phases. Terranes begin growing in an oceanic setting, as in the case of island arcs, or by some calving process that rifts a fragment from pre-existing crust. A buoyant piece of crust attached to an oceanic plate will ultimately collide with other buoyant pieces of crust. In order to distinguish between two similar processes of collision tectonics (the joining of terranes in an oceanic setting and the joining along a continental margin), it is useful to use ‘amalgamation’ for the former and ‘accretion’ for the latter, even though the tectonic processes of both may be similar. Following all collision events, disruption is likely to continue, resulting in the dispersion of the accreted fragments. Dispersion tectonics in many orogenic belts has dramatically affected the distribution of the accreted terranes. Terrane boundaries are rheologically weak zones between terranes, e.g., faults or complex fault zones and suture zones. Faults or sutures may be inferred between areas with different rock units if the boundary is not exposed, and the units on either side cannot be linked by reasonable lithofacies characteristics. Details of the processes of accretion and amalgamation are still largely unknown. Thrust faulting seems to play an important role in accretion, but many thrust faults are later modified by folding and high-angle faulting. Confusion may arise in discriminating between faults or suture-bounded terranes and fault-bounded packages within a terrane, such as a series of thrust sheets stacked one upon the other in what is commonly called a duplex configuration. Similarly, distinct lithologic packages that are separated by unconformities do not constitute different terranes. A volcanic sequence built on an ophiolite that is overlain by thick pillow basalt, red beds, shallow- marine, and finally deep-marine pelagic strata, does not necessarily indicate two or more terranes.

10.3

Formation and Amalgamation of Gondwana Supercontinent

Gondwana was the largest mass of continental crust on Earth for more than two hundred million years. Gondwana is the name for the southern half of an ancient supercontinent known as Pangaea that exist some 300 Ma, along with a northern supercontinent known as Laurasia (Meert 2012). The term Gondwana was first used in geological literature to refer to a plant-bearing series in India and was afterward extended to the Gondwana system (Feistmantel 1876;

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Medlicott and Blanford 1879, and references therein). The continent broke-up about 180 million years ago and eventually split into the landmasses we recognize today: South America, Africa, Madagascar, Sri Lanka, India, the Arabian Peninsula, Antarctica, and Australia (Fig. 10.1). Several models have been proposed for the Gondwana assembly. Some view the collision as a simple unification of East Gondwana (Fig. 10.2): India, East Antarctica, Madagascar, Australia, and Sri Lanka) with West Gondwana Africa and South America. These models (Yoshida 1995; Yoshida and Upreti 2006; Squire et al. 2006) tend to oversimplify the geologic data and ignore earlier (Neoproterozoic) reconstructions that demonstrate latitudinal offsets between East Gondwana blocks (Torsvik et al. 2001a, b; Meert 2003; Veevers 2004; Collins and Pisarevsky 2005). Gondwana supercontinent in eastern and southern Africa evolved by collision and amalgamation of two crustal plates, provisionally named East Gondwana and West Gondwana and consumption of the Mozambique Ocean between 841 and 632 Ma. East Gondwana comprised the ANS shield and older crystalline basement in Madagascar, India, Antarctica, and Australia. West Gondwana was composed of most of Africa and South America. Gondwana supercontinent existed from the Neoproterozoic (about 550 million years ago)

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until the Carboniferous (about 320 million years ago). Gondwana became the largest piece of continental crust of the Paleozoic Era, covering an area of about 100 million km2 (Torsvik and Cocks 2013). During the Carboniferous, Gondwana amalgamated with Euramerica to form a larger supercontinent called Pangaea. The amalgamation of Gondwana took place during one of the most dynamic periods known in the evolution of the climate (Hoffman et al. 1998; Evans 2000) and the deep Earth (Kirschvink et al. 1997; Evans 1998). Understanding supercontinent, amalgamation requires detailed knowledge of the orogens that assembled the former continental fragments together. This requires knowledge of the component lithofacies, the gross crustal framework, the geometry of the major fault and shear zones as well as the thermal and temporal aspects of deformation, metamorphism, and magmatism. A polyphase assembly of Gondwana during the East Africa, Brasiliano, Kuungan, and Damaran orogenies resulted in an extensive mountain chain which delivered detritus into a shifting oceanic realm. Gondwana (and Pangaea) gradually broke-up during the Mesozoic Era. During Silurian, Gondwana supercontinent (Australia, Antarctica, India, Arabia, Africa, and South America, Florida, southern Europe, and the Cimmerian

Fig. 10.1 Gondwanaland supercontinent. West Gondwana is shaded in light blue, and East Gondwana is shaded yellow (modified after Gray et al. 2008). Neoproterozoic orogenic belts cross the supercontinent. Those associated with the final amalgamation of the supercontinent are the East African Antractica Orogen (AEAO) (750–620 Ma; red), the Brasiliano—Damara Orogen (630–520 Ma; light blue), and the Kuungan Orogen (570–530 Ma; green). Black arrows show inferred convergence and location of the most intense continent–continent collision. Open arrow shows the possible direction of a tectonic escape. Modified after Stern and Johnson (2010), Meert and Lieberman (2008)

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Fig. 10.2 Collision between the various Gondwana cratons along the Brasiliano and Pan-African orogens at 650–550 Ma; juvenile crust formation followed by arc collision and northwards extrusion in the Arabian-Nubian sector and extensive reworking of the Saharan metacraton (after Johansson 2014). Initial breakup between Laurentia, Amazonia and Baltica at 600 Ma (after Abd El-Wahed and Hamimi 2021)

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terranes, namely Turkey, Iran, Afghanistan, Tibet, and the Malay Peninsula) were centered over the South Pole. The Rheic Ocean was an east-west ocean that separated the southern European part of Gondwana from northern Europe (Baltica) and was basically a southwestern extension of the Paleotethys Sea. During the Caledonian Orogeny, about 410 million years ago (during the Devonian), the lesser supercontinent, Euramerica, was created by collision between the Avalonia cratons, Baltica, and Laurentian. The Caledonian orogeny was a mountain building era recorded in the northern parts of Ireland and Britain, the Scandinavian Mountains, Svalbard, eastern Greenland, and parts of north-central Europe (Torsvik and Cocks 2013). In the Permian, the Euramerica supercontinent became a part of the major supercontinent Pangaea. In the Jurassic, when Pangaea rifted into two continents, Gondwana and Laurasia, Euramerica was a part of Laurasia. Some 180 million years ago, in the Jurassic Period, the western half of Gondwana (Africa and South America) separated from the eastern half (Madagascar, India, Australia, and Antarctica). About 140 million years ago, in the Cretaceous, Africa separated from South America and the South Atlantic Ocean opened. At about the same time, the central Indian Ocean opened as India broke away from Madagascar, and Australia slowly drifted away from Antarctica. Also, Laurasia is divided into the North America and Eurasia continents. Baltica became a part of Eurasia, and the Laurentian craton became a part of North America, while Avalonia was split between the two. The remnants of Gondwana make up about two-thirds of today’s continental area (Torsvik and Cocks 2013). Fifty million years ago, India collided with Eurasia forming the Himalayan Mountains, while the northward-moving Australian plate had just begun its collision along the southern margin of Southeast Asia—a collision that is still progressing today. The amalgamation of Gondwana started at ca. 630 Ma and extended to ca. 550–530 Ma, when subduction along its proto-Pacific margin was already established (Dalziel 1997; Cordani et al. 2003; Meert and Torsvik 2003; Cawood 2005; Collins and Pisarevsky 2005; Cawood and Buchan 2007). Likewise, Pannotia was considered as a late Neoproterozoic ‘short-lived’ supercontinent that included Laurentia and Gondwanan domains, prior to Gondwana final configuration (Powell and Young 1995; Dalziel 1997). On the other hand, the breakup of Rodinia took place in two phases at ca. 800–700 Ma and after ca. 600 Ma, the latter being coeval with the timing of Gondwana assembly (Cordani et al. 2003; Cawood 2005; Li et al. 2008). Deciphering the construction of East Gondwana supercontinent, comprising the Arabian-Nubian Shield and older crystalline basement in Madagascar, India, Antarctica, and Australia, requires understanding of the Pan-African/EAO. On the other hand, development of the West Gondwana supercontinent (composed of most of Africa and a mosaic of

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South American cratons, linked by a complex set of fold belts) requires an understanding of the Brasiliano/ Pan-African orogenies. Collision and amalgamation of East and West Gondwana produced the EAO. Several events, collectively known as the PAO, led to the amalgamation of most of the continental fragments to form a much older supercontinent, Rodinia, which began to form at c. 1.23 Ga by accretion and collision of fragments produced by breakup of an even older supercontinent, Columbia, which assembled during 2.0–1.8 Ga collisional events (Zhao et al. 2002, 2004). The assembly of Rodinia during the Grenville Orogeny, incorporated most of the Precambrian continental blocks, and lasted ca. 400 m.y. (from 1300 to 900 Ma). Rodinia assembly was associated with collisions that produced late Mesoproterozoic (1300 to 1000 Ma) mobile belts (Pisarevsky et al. 2003). Condie (2002) concluded that Rodinia formation took place between 1300 and 900 Ma and considered that the continental breakup occurred between 950 and 600 Ma. Its breakup was presumably triggered by a mantle superplume that took place between 830 and 650 Ma (Bogdanova et al. 2009; Li et al. 2008; Meert 2012). Rodinia was surrounded by an ocean called Mirovia. Rodinia broke-up in four stages between 825 and 550 Ma (Bogdanova et al. 2009) (Fig. 10.3). The superplume that initiated the breakup generated crustal arching, intense bimodal magmatism, and accumulation of thick rift-type sedimentary successions recorded in South Australia, South China, Tarim, Kalahari, India, and the ANS. Rifting progressed in the same cratons 800–750 Ma and spread into Laurentia and perhaps Siberia. India (including Madagascar) and the Congo-Säo Francisco Craton were either detached from Rodinia during this period or were simply never were part of the supercontinent. As the central part of Rodinia reached the Equator around 750–700 Ma, a new pulse of magmatism and rifting continued the disassembly in western Kalahari, West Australia, South China, Tarim, and most margins of Laurentia. By 650–550 Ma, several events coincided as follows: the opening of the Iapetus Ocean; the closure of the Braziliano, Adamastor and Mozambique Oceans, and the PAO. The result was the formation of Gondwana. Assembling of the continental blocks of West Gondwana started around 900–700 Ma interval, with final amalgamation of the whole Gondwana around 550–530 Ma, based on paleomagnetic, geologic, and isotopic data (Meert and Van Der Voo 1997; Rogers and Santosh 2004; Teixeira et al. 2007). Since the 1990s, workers have recognized three orogenies: the EAO (650–800 Ma) and Kuunga Orogeny (including the Malagasy Orogeny in southern Madagascar) (550 Ma)—the collision between East Gondwana and East Africa in two steps, and the Brasiliano Orogeny (660– 530 Ma)—the collision between South American and African Cratons (Meert and Van Der Voo 1997).

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Fig. 10.3 Stages in the tectonic evolution of the Arabian-Nubian Shield and the East African Orogen, in the supercontinent cycle bracketed by the breakup of Rodinia and the assembly of Gondwana. After Stern and Johnson (2010)

10.3.1 Neoproterozoic Pan-African Orogeny The term ‘Pan-African’ (Kennedy 1964) refers to a sequence of tectono-thermal events at 500 * 100 Ma within Africa and adjacent Gondwana elements. The term was broadened by Kröner (1984) to include orogenic events of the same time range (950–450 Ma) on a more global scale. In the intervening years, the Pan-African orogenic cycle has been more narrowly defined both spatially and temporally, such that it is possible to recognize individual orogenic events within Gondwana (Trompette 1997; Stern 1994; Meert et al. 1995). Later, when plate tectonics became generally accepted, the term Pan-African was extended to all of the Gondwana supercontinent. Because the formation of Gondwana encompassed several continents and extended from the Neoproterozoic to the early Palaeozoic, Pan-African could no longer be considered a single orogeny (Meert 2003), but rather an orogenic cycle that included the opening and closing of several large oceans and the collisions of several continental blocks (Kröner and Stern 2004). Other terms

were used for similar orogenic events on other continents, e.g., Brasiliano in South America, Adelaidean in Australia, and Beardmore in Antarctica. Furthermore, the Pan-African events are contemporaneous with the Cadomian orogeny in Europe and the Baikalian orogeny in Asia, and crust from these areas was probably part of Pannotia (i.e., Gondwana in its earliest stages) during the Precambrian (Kröner and Stern 2004). This tectono-thermal event was later recognized to constitute the final part of an orogenic cycle, leading to orogenic belts which are currently interpreted to have resulted from the amalgamation of continental domains during the period 870 to 550 Ma. Pan-African events culminated in the formation of the late Neoproterozoic supercontinent Gondwana (Fig. 10.4). The PAO was a series of majors. Neoproterozoic orogenic events, related to the formation of the supercontinents Gondwana and Pannotia (Fig. 10.5), about 600 million years ago. This orogeny is also known as the Pan-Gondwanan or Saldanian Orogeny. The PAO and the Grenville orogeny are the largest known systems of orogenies on Earth (Kröner

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Fig. 10.4 Map of Gondwana at the end of Neoproterozoic time 540 MA) showing the general arrangement of Pan-African belts. AS Arabian Shield: BR, Brasiliano: DA, Damara: OM, Dom Feliciano: DR, Denman Darling: E–W, Ellsworth-Whitmore Mountains: GP. Gariep: KB, Kaoko: MA, Mauretanides; MB, Mozambique Boelt: N–S, Nubian Shield; PM, Peterman Ranges; PB, Pryolz Bay, PR Pampean Ranges: PS, Paterson, QM, Queen Maud Land: RB, Rokolides; SD, Saldania: SG, Southern Granulite Terrane: TS. Trans Sahara Belt: WB, West Congo: ZB, Zambezi (after Kusky et al. 2003)

et al. 2008 and Van Hinsbergen et al. 2011). The Pan-African orogenic cycle was the result of ocean closure, arc and microcontinent accretion, and final suturing of continental fragments to form the supercontinent Gondwana. Kröner and Stern (Kröner and Stern 2004) opined that the African and South American cratons were never part of Rodinia. The Pan-African orogeny (PAO) comprised different orogenic belts (Kröner and Stern 2004) including the ANS; the Mozambique Belt (MB), extending from east Antarctica through East Africa, up to the ANS; the Zambezi Belt, branching off from the MB in northern Zimbabwe and extending into Zambia; the Damara Belt, exposed in Namibia between the Congo and Kalahari cratons and continuing southwards into the coastal Gariep and Saldania Belts, and northwards into the Kaoko Belt.

10.3.2 The East African Orogen (EAO)

Fig. 10.5 The ANS in the supercontinent Pannotia c. 570 million years ago, before the opening of the Red Sea (after Meert 2002)

The assembly of the eastern part of Gondwana (eastern Africa, Arabian–Nubian shield (ANS), Seychelles, India, Madagascar, Sri Lanka, East Antarctica and Australia)

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resulted from a complex series of orogenic events spanning the interval *750 to *530 Ma (Meert 2002). A detailed examination and geochronologic studies eastern Gondwana suggest a multiphase assembly. Two main periods of orogenesis within eastern Gondwana have been identified. The older orogen resulted from the amalgamation of arc terranes in the Arabian—Nubian shield region and oblique continent—continent collision between eastern Africa (Kenya–Tanzania and points northward) with a still ill-defined collage of continental blocks, including parts of Madagascar, Sri Lanka, Seychelles, India and East Antarctica, during the interval *750 to 620 Ma. The second major episode of orogenesis took place between 570 and 530 Ma and resulted from the oblique collision between Australia, plus an unknown portion of East Antarctica, with the elements previously assembled during the East African Orogen. This episode is referred to as the Kuunga Orogeny, following the suggestion of Meert et al. (1995). The EAO is one of Earth’s great collision zones, where East and West Gondwana collided to form the supercontinent ‘Greater Gondwana’ or ‘Pannotia’ at the end of Neoproterozoic time (Stern 2002) (Fig. 10.5). Pannotia was an enigmatic short-lived supercontinent that came into existence at the end of the Neoproterozoic because of the reassembly of fragments of Rodinia (Scotese 2009). The EAO is an extensive Neoproterozoic accretionary orogen and collisional zone within Gondwana (Stern 1994; Collins and Windley 2002; Cawood et al. 2009). It was realized that high-grade metamorphism in East Africa was not older than Neoproterozoic (Coolen et al. 1982), despite the Mesoproterozoic to Archean protolith ages of some of the rocks caught up in the orogen (Möller et al. 1998, Tenczer et al. 2006). Greenwood et al. (1980) extended the (MB) northward into the ANS, and Berhe (1990) described ANS-type ophiolite-decorated north-trending shear zones extending south into central Kenya, confirming that the MB and the Arabian—Nubian Shield are correlatives. The Neoproterozoic orogenic belts were present in other parts of greater Gondwana at the end of the Precambrian as the result of similar events of crustal accretion (e.g., Grunow 1999; Abdelsalam et al. 2003; Collins and Pisarevsky 2005; Brito Neves and Cordani 1991). The EAO records the amalgamation of arc terranes in the Arabian–Nubian shield region and the collision of a major landmass to the south. Stern and Abdelsalam (1998) argue that terrane amalgamation in the ANS region began around 750 Ma. Several geochronologic studies suggest that the entire region was assembled by 630 Ma, with continental escape in the ANS following shortly thereafter (Fig. 10.6). Farther south in Kenya and Tanzania, the available geochronology indicates a major episode of high-grade metamorphism at *640 Ma, although others have

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suggested that this metamorphism was unrelated to continental collision (Möller et al. 2000; Kröner et al. 2000). Kröner et al. (2000) suggest that east Gondwana was a collage of terranes that amalgamated during a 100 My time period, but they did not provide details on the make up of the individual blocks that collided during the 640–530 Ma interval. The collision in the southern part of the EAO involved oblique continent–continent collision, and the EAO further north records the accretion of juvenile terranes and tectonic escape (Stern 1994). There is no evidence of a ca. 640 Ma orogenic episode in India, Sri Lanka or East Antarctica, but this may also be due to the fact that collision in those regions was younger than southern and central Madagascar and that these areas were relatively unaffected by the older deformation. The MB is one of the orogenic belts formed during 800– 650 Ma and was originally interpreted as the suture between East and West Gondwana. The Mozambique Ocean separated the Congo–Tanzania–Bangweul Block of central Africa from Neoproterozoic India (India, the Antongil Block in far eastern Madagascar, the Seychelles, and the Napier and Rayner Complexes in East Antarctica). The Azania continent (much of central Madagascar, the Horn of Africa and parts of Yemen and Arabia) was an island in the Mozambique Ocean (Collins and Pisarevsky 2005). Azania is a name that has been applied to various parts of southeastern tropical Africa (Collins and Pisarevsky 2005). The final formation of Gondwana occurred about 500 million years ago. Gondwana is superficially divided into a western half (South America and Africa) and an eastern half (Madagascar, Antarctica, Sri Lanka, India and Australia) and Ediacaran-Cambrian age, by coalescence of East Gondwana with West Gondwana by closing of Mozambique ocean and development of the MB (Stern 1994; Meert 2012; Johansson 2014).

10.4

The Arabian-Nubian Shield (ANS)

The ANS is most prominent and type example of juvenile crustal province of Neoproterozoic age (1000–542 Ma) (e.g., Bentor 1985; Johnson and Woldehaimanot 2003; Stern 1994, 2002) and has, in addition, small amounts of Archean and Paleoproterozoic continental crust. It is exposed on the flanks of the Neoproterozoic Saharan Metacraton in western Egypt, Sudan and Eritrea (Abdelsalam et al. 2002) and in Yemen and west-central Saudi Arabia (Whitehouse et al. 2001b) (Fig. 10.7). These older rocks represent parts of the Rodinia supercontinent, which was assembled through global scale orogenic events between 1300 Ma and 900 Ma, and diachronously broke-up at the beginning of the Neoproterozoic, between ca. 900 Ma and 740 Ma (Stern 1994). Rifting of the supercontinent created the Mozambique

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Fig. 10.6 The ANS in relation to East Africa as component parts of the EAO. Although ANS extends as far south as southern Kenya, the geographic focus of this review is on the ANS between northern Ethiopia and Sinai and Jordan (after Johnson et al. 2011)

Ocean, which was the site of formation of the Neoproterozoic rocks of the ANS, and at the end of the Neoproterozoic, during the PAO, was closed to form the EAO (Fig. 10.8) (Stern 1994; Meert 2003; Genna et al. 2002). The ANS

extends over 3500 km north-south and more than 1500 km east-west at its widest, underlying an area of *2.7  106 km2 in the northern half of the (EAO) (Johnson 2014). The ANS is generally viewed as a combination of island arc

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Fig. 10.7 Arabian–Nubian Shield (ANS), showing its relationship to adjacent areas of older continental crust. The border with the Saharan Metacraton, on the west, is the effective contact between ANS and West Gondwana: A border, on the east, with putative East Gondwanan crust, is not certain. NED = North Eastern Desert; CED = Central Eastern Desert; SED = South Eastern Desert (after Johnson et al. 2011)

complexes accreted during the closure of the Mozambique Ocean between East and West Gondwana (Bentor 1985; Stern 1994). The juvenile character of island arcs and ophiolites that comprise the ANS core is exposed along the Red Sea margins, from Jordan in the north, to Ethiopia in the south. The shield is divided by variably oriented sutures, shear zones, and fold belts into northern and southern

sectors. The southern shield is dominated by northerly trends, while the northern sector shows a variety of trends (Johnson et al. 2011 and Fritz et al. 2013). The ANS consists of pre-Phanerozoic, mostly Neoproterozoic, rocks that crop out on both sides of the Red Sea in western Arabia and northeastern and eastern Africa within Saudi Arabia, Yemen, Egypt, Sudan, Ethiopia, Eritrea, and Somalia and Kenya.

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Fig. 10.8 ANS in relation to older crust on its margins (Fritz et al. 2013). Inferred contact with the Saharan Metacraton shown by a thick dashed line. Ophiolites schematically shown after Berhe 1990: A. Allaqi; Ad Adola; Ak Akobo; B. Baragoi; B.U. Bi’r Umq; E. Jabal Ess; G. Gebel Gerf; H. Halaban; K. Kinyiki; M. Moyale; M.S. Moroto-Sekerr; N. Nuba; Na Nakasib; S. Sol Hamed; T. Jabal Thurwah; Ta Jabal Tays; Tu Bi’r Tuluhah; T.Y. Tuludimtu-Yubdo; U. Jabal Uwayjah; W. Jabal Wask. Inset is the late Proterozoic belts in Gondwana (modified after Johnson 2014)

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The ANS consists dominantly of juvenile Neoproterozoic rocks (Stern et al. 1984) and small amounts of Archean and Paleoproterozoic continental crust exposed on the flanks of the Neoproterozoic Saharan Metacraton (Abdelsalam et al. 2002) and in Yemen and west-central Saudi Arabia (Whitehouse et al. 2001a).

10.4.1 Terrane Analysis, Amalgamation, and Accretion of the ANS It is now widely accepted that the (ANS) is formed by terrane accretion. The shield is divided into lithostratigraphic terranes, though the number of terranes and the identity of their boundaries are debated, and opinion has changed with time. The older components of the shield include Archaean and Palaeoproterozoic continental crust, and Neoproterozoic (c.870–670 Ma) continental marginal and juvenile intraoceanic magmatic-arc terranes that accumulated in an oceanic environment referred to as the Mozambique Ocean (Yoshida et al. 2003). Subduction, beginning c. 870 Ma, and initial arc–arc convergence and terrane suturing at c. 780 Ma, marked the beginning of ocean-basin closure and Gondwana assembly. Terrane amalgamation continued until c. 600 Ma, resulting in the juxtaposition of East and West Gondwana across the deformed rocks of the shield. Final assembly of Gondwana was achieved by c. 550 Ma, following overlapping periods of basin formation, rifting, compression, strike-slip faulting, and the creation of gneiss domes in association with extension and/or thrusting. Most post-amalgamation basins contain terrestrial molasse deposits, but those in the eastern Arabian Shield and Oman also have marine to glaciomarine deposits, which indicate marine penetration across the orogen soon after orogeny. The contacts between terranes in the ANS (Johnson 2012) include (l) sutures composed of ophiolite-decorated shear zones; (2) cryptic or shear zones that may be original sutures, faults superimposed on original sutures or post-suturing structures; (3) post-amalgamation fault zones that may be unrelated to original suturing events. Terranes of the ANS vary enormously in size, from small to continental scale blocks. Many terranes are now disjunct, having been dismembered by post-accretionary sutures, fault zones and strike-slip faulting (e.g., Najd fault system). At many places in the ANS, low-temperature metamorphism and melange formation accompanied accretion. The terrane-forming rocks are folded and metamorphosed and formed a newly amalgamated basement into which late- to post tectonic plutons were intruded and on which younger sedimentary and volcanic basins were deposited (Johnson 2012). Terranes in the shield largely consist of volcanic, volcaniclastic, and intrusive assemblages that formed as arcs above subduction zones in the

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Mozambique Ocean, except one terrane in the eastern part of Arabian Shield (the Ad Dawadimi terrane) that represents a sedimentary accretionary prism or fore-arc deposit (the Abt formation). Early studies of mafic-ultramafic complexes and their identification as ophiolites led to models of back-arc basin development in arc complexes (Bakor et al. 1976). Displacement between terranes need not be large but must be sufficient to juxtapose dissimilar rocks and disrupt original facies trends. Determining offsets along faults and/or shear zones has established displacements for some terranes on the order of a few hundred kilometers. The distribution and nature of ophiolites in the ANS, depicted on the terrane maps (Fig. 10.9), provide insights into the accretion process. The growth and shaping of continents and microcontinents of the ANS may be viewed as resulting from both terrane accretion and terrane dispersion. In contrast to terranes accretion, which results in continental growth or outbuilding, the terrane dispersion results in the diminution of continents. The terranes of the ANS are presented and mapped by different authors (e.g., Calvez et al. 1983; Johnson and Vranas 1984; Stoeser and Camp 1985; Stoeser and Stacey 1988; Vearncombe 1983; Berhe 1990; Ries et al. 1992; Mosley 1993; Kröner et al. 1991; Abdelsalam and Stern 1996; Stern et al. 2010; Johnson et al. 2011). The essence of the terrane maps is the presentation of the packaging of terranes that compose the area being addressed. Some detailed terrane maps show, in addition to the terrane sutures, the distribution and limits of depositional overlap sequences or plutons. The chronology of these features may be depicted directly on the map or in an accompanying tectonic assembly diagram (Stern and Johnson 2010). A feature of all terrane maps is the composition of the terranes, tectonic settings (e.g., island arc, seamount, or continental affinities, or mobile belts of terranes or individual terranes in the context of an accretion sequence). The structural relationships among ophiolites and belts of volcanic rocks led to the idea that the shield is divided into segments by ophiolite zones or sutures, with the implication that the shield was built up of generations of juxtaposed volcanic arcs (Frisch and Al Shanti 1977). Stoeser and Camp (1985) subdivided the Arabian Shield into five terranes: Midyan, Hijaz, Ar Rayn, Asir and Afif (Fig. 10.10). Johnson and Vranas (1984), Johnson and Woldehaimanot (2003), and Genna et al. (2002), mapped eight terranes; and Stoeser and Frost (2006) suggested as many as 14 terranes. A critical question about terrane identification is how to correlate and join up the ophiolitic bodies of the shield along sutures (Church 1988). Separate terranes are recognized also in the southernmost part of the Arabian Shield, in Yemen, and in the Nubian Shield, as described, for example, by Windley et al. (1996), Vail (1985), and Johnson and Woldehaimanot (2003).

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Fig. 10.9 Tectonic map of the ANS (after Johnson and Woldehaimanot (2003), showing the distribution of tectonmagmatic terranes, sutures, and mafic-ultramafic (ophiolite) suites. Ages indicated are terrane protoliths (after Hargrove et al. 2006). Similar patterns indicate the correlation of terranes across the Red Sea

Stoeser and Stacey (1988) subsequently referred to the Asir and Afif as ‘composite terranes’ based on the concept that both crustal units were in fact composed of multiple terranes which had accreted into unified crustal blocks or terranes prior to collision along the Nabitah suture zone. The Asir composite terrane, as defined by Stoeser and Camp (1985), is composed of at least four ensimatic arc assemblages, the An Nimas, Bidah, and Jiddah terranes, which are older than about 800 Ma and the Al Qarah terrane which is less than 740 Ma. Johnson and Kattan (2001) and Johnson and Woldehaimanot (2003) have continued reference to the

Asir terrane, but with the Jiddah as a separate terrane and including the Tathlith-Malahah terrane in the south, and at least two or three Neoproterozoic ensimatic arc terranes in the north. The arc terranes have also been subdivided into two classes: volcanic arc assemblages and basinal arc assemblages, such that the first represents volcanic assemblages formed on or near the main line of arc magmatism, whereas the latter is more likely formed in fore- or back-arc environments. All of these terranes, except one, have been interpreted to represent oceanic-arc assemblages (Stoeser and Stacey 1988; Pallister et al. 1988; Harris et al. 1990),

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Fig. 10.10 Tectonic sketch map of the Arabian Shield showing microplates and suture zones (after Stoeser and Camp 1985)

either as primary volcanic arcs or flanking sedimentary basins, which may be either fore-arc or back-arc in origin. The one exception is the Khida terrane of the eastern Shield, which is interpreted be a microplate of reworked older continental crust (Stacey and Agar 1985; Stoeser and Stacey 1988). The Arabian Shield has been analyzed in terms of tectono-stratigraphic terranes since Stoeser and Camp (1985), Johnson and Vranas (1984), and Vail (1985), who built on the earlier paper by Camp (1984), mentioned above, describing how island arcs in the shield came to be juxtaposed along subduction zones. The lithostratigraphic terranes of the shield form small to large crustal blocks that are bounded by major shear zones. Some shear zones are sutures, i.e., regions in which oceanic crust was consumed during the process of subduction and magmatic-arc

convergence, for example, the Bi’r Umq suture between the Jiddah and Hijaz terranes. Others are transcurrent faults that may reflect the original locations of sutures but are basically zones of late Neoproterozoic strike-slip strain, such as the Ad Damm fault between the Asir and Jiddah terranes, and the Al Amar fault between the Ad Dawadimi and Ar Rayn terranes. The terranes comprise the earliest formed rocks in any given part of the shield and mostly originated in a juvenile Neoproterozoic ocean. As a result of ongoing subduction and the consumption of the intervening oceanic crust, the terranes in the shield converged, amalgamated, and sutured, reflecting an orogenic process characterized by metamorphism, deformation, and syntectonic intrusion. Terranes in the shield are also recognized based on variations in isotopic characteristics and on structural, geochronologic, and lithostratigraphic differences, suggesting

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Fig. 10.11 Principle of tectono-stratigraphic terrane amalgamation, showing the stages and appropriate nomenclature of amalgamation and eventual accretion to continental crust (after Johnson 2012)

different geologic histories in different parts of the shield. The terranes were amalgamated via suturing by convergence and collision above subduction zones. They are bounded by serpentinite-decorated shear zones, which mark the sites of suturing, or by major transcurrent shear zones and are ‘stitched’ together by granitic plutons that intrude rocks on either side of the suture zones and constrain the minimum age of suturing (Fig. 10.11).

10.4.2 Assembly of the ANS and Crustal Growth The ANS is the result of a 300 million year period of crustal growth, involving the accretion of juvenile subductionrelated magmatic arcs, periods of sedimentation and volcanism in basins developed on newly post-amalgamated arc terranes, crustal thickening, the emplacement of large amounts of granitic magma, periodic uplift and erosion, extension, and tectonic escape on transcurrent faults. The ANS constitutes a southward narrowing belt, internally structured by individual terranes. Its western margin is

defined by juxtaposition of ophiolite-decorated volcanosedimentary sequences and juvenile Neoproterozoic arc magmatic terranes with the Eastern Granulite complex of the MB, the Archean Congo Craton, and the Sahara Metacraton. The EAO is subdivided into the ANS in the north, composed largely of juvenile Neoproterozoic crust (e.g., Stern 1994, 2002; Johnson and Woldehaimanot 2003; Johnson et al. 2011), and the MB in the south comprising mostly pre-Neoproterozoic crust with a Neoproterozoic early Cambrian tectono-thermal overprint (Figs. 10.6 and 10.12). In the north, the western margin of the ANS is not defined because it is covered by Mesozoic to Cenozoic sedimentary rocks; however, it extends along the line of the Nile Valley and crops out in the Keraf arc-continent suture in northern Sudan (Abdelsalam et al. 1998). South of this, the margin is defined by a line of sutures and ophiolite belts, namely the Kabus suture of the Nuba Mountains (Abdelsalam and Dawoud 1991), the Sekerr ophiolite of northwestern Kenya (Vearncombe 1983; Berhe 1990; Ries et al. 1992; Mosley 1993) and the Kinyiki ophiolite of southern Kenya (Frisch and Pohl 1986) (Fig. 10.13).

238 Fig. 10.12 Distribution of crustal domains in the EAO. SM, Sahara Metacraton; CTB, Congo– Tanzania–Bangweulu Cratons; ZKC, Zimbabwe– Kalahari Cratons; I, Irumide Belt; A, Antogil Craton; M, Masora Craton; ANS, (after Abd El-Wahed and Hamimi 2021)

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Fig. 10.13 Crustal age domains in the northern EAO and crustal growth phases (accretion stages) in the ANS. CT, Congo–Tanzania Craton; EG, Eastern Granulites (after Fritz et al. 2013)

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The southern tip of the ANS is represented by a belt of 955–845 Ma old subduction-related amphibolites and gneisses adjacent to the Galana-Athi shear zone (Hauzenberger et al. 2007; Bauernhofer et al. 2009). The same shear zone defines the western margin of the Galana terrane considered to be part of Azania. The eastern ANS margin, the inferred contact with Azania, is not well-defined in Kenya, Ethiopia, and Somalia. However, it is marked by the Mutito-Bruna Shear Zone (Mosley 1993), which separates the Barsaloi-Adola Moyale ophiolitic belts from the poorly exposed older suites of the Burr Complex in southern Somalia (Warden and Horkel 1984). The Qabri Bahar and Mora Complexes (Fig. 10.13) containing preNeoproterozoic zircons (Kröner and Sassi 1996) may represent Azania in northern Somali. The ANS-Azania boundary in Yemen is probably defined by arc-continent sutures along the Abas and Al Mahfid terranes. The Afif terrane in eastern Saudi Arabia, including the Paleoproterozoic Khida subterrane, likely represents a crustal block within the ANS, since its boundaries are defined as arc–arc sutures (Whitehouse et al. 2001a; Johnson et al. 2011). Fritz et al. (2013) suggests that the ANS evolved through four main phases of crustal growth. The southern ANS, extending from southern Kenya to the Nakasib-Bir Umq suture (Fig. 10.13), formed first (Johnson et al. 2003). This part of the ANS constitutes terranes that are internally partitioned by arc–arc sutures (Kröner et al. 1991; Abdelsalam and Stern 1996; Stern et al. 2010; Johnson et al. 2011). The terranes are known as Tokar–Barka, Butana, Haya terranes in Eritrea, Ethiopia, and Sudan and Abbas, Asir, Jiddah, and Hijaz terranes in Yemen and Saudi Arabia. Protolith ages from these terranes extend back to 900–830 Ma. The tectonic evolution of the southern ANS through a Wilson Cycle is well illustrated in the Tuludimtu belt of central Ethiopia (Woldemichael et al. 2010). Here, early rifting was initiated between 900 and 860 Ma; the transition from rifting to ocean floor spreading occurred between 860 and 830 Ma; subduction and formation of arc- and back-arc basins occurred between 830 and 750 Ma; basin closure by accretion of island arcs commenced between 750 and 650 Ma. To the north, amalgamation of arc systems by closure of internal arc–arc sutures (e.g., Barka and Bir’ Umq-Nakasib sutures) occurred between 800 and 700 Ma (Abdelsalam and Stern 1996; Yibas et al. 2002; Woldemichael et al. 2010). The central ANS, lying between the Nakasib-Bir Umq suture in the south and the Yanbu-Onib-Sol Hamed-Gerf-AllaqiHeiani suture in the north (Fig. 10.13), developed by arc formation and accretion between 830 and 710 Ma (e.g., Johnson et al. 2003, 2011). The ages for the central ANS igneous rocks overlap with those obtained from the southern ANS but also extend to younger ages.

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The northern ANS formed during a second growth phase between *760 and 730 Ma when the Midyan-Eastern Desert terrane was formed. This terrane subsequently collided and amalgamated with the earlier formed older terranes along the Yanbu-Onib-Sol Hamed-Gerf-Allaqi-Heiani suture (Ali et al. 2010; Johnson et al. 2011). The resulting geologic entity is commonly referred to as the ‘western arc or oceanic terranes’ of the ANS (e.g., Stoeser and Frost 2006; Ali et al. 2009; Johnson et al. 2011). However, the northernmost ANS in Sinai also contains older rocks (1025 Ma rocks of the Sa’al Complex; Be’eri-Shlevin et al. 2012). In the third growth phase, the ‘western arc or oceanic terranes’ of the ANS subsequently collided and amalgamated between 680 and 640 Ma with the Afif and Tathlith terranes, creating a neocontinental crustal block referred to as the proto-ANS (Johnson et al. 2011). The 680–640 Ma assembly of the p-ANS was associated with metamorphic, deformational, and intrusive events that, in the Arabian Shield, are referred to as the ‘Nabitah orogeny’ (680– 640 Ma) (Stoeser and Stacey 1988). The orogeny is named after the Nabitah fault, the suture between the Asir and Tathlith terranes and a ductile shear zone within the Asir terrane. The fault gives its name to the Nabitah mobile belt (Stoeser and Stacey 1988) (Fig. 10.14), a broad zone of deformation and metamorphism that trends N–S across the Arabian Shield. A suturing event along the eastern margin of the Afif terrane is marked by the formation of the Halaban ophiolite at *680–670 Ma (Al-Saleh et al. 1998). The youngest ANS terranes, with late Cryogenian to Ediacaran protoliths, are the Ad Dawadimi and Ar Rayn terranes in the easternmost part of the ANS in Arabia (Fig. 10.13). Ad Dawadimi and Ar Rayn terranes are in contact along the serpentinite-decorated Al Amar fault, which is interpreted as a suture. This suture can be traced magnetically in the subsurface far to the north and is known as the Central Arabian Magnetic anomaly. The anomaly may mark the eastern limit of the ANS (Stern and Johnson 2010), but this is controversial and even younger sutures may lie farther east (Cox et al. 2012). The pre-Phanerozoic crust in eastern Arabia is concealed by Phanerozoic sediments but, where locally exposed in Oman, it appears to have stabilized in the Cryogenian (*750–700 Ma) (Stern and Johnson 2010). In the fourth growth phase, during the late Neoproterozoic (650–542 Ma), the (ANS) experienced final assembly and accretion to the Saharan Metacraton concurrent with the assembly of eastern and western Gondwana (Johnson et al. 2011). By the end of the Neoproterozoic, the ANS lay at one end of the East African Orogen, with its northern margin forming a low-relief stable shelf facing an open ocean, and to the south the ANS transitioned into the Mozambique Belt. The geologic history of the ANS during this period provides

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Fig. 10.14 Terrane assembly in the ANS showing inferred ages of suturing and fault movements, and trajectories of amalgamation (after Johnson et al. 2011)

insight into the closing developmental stages of one of the world’s largest accretionary orogens. Following a 680– 640 Ma orogenic event reflecting amalgamation of a core grouping of island arc terranes (the proto-Arabian–Nubian Shield), the region underwent extensive exhumation, erosion, and subsidence. Depositional basins formed in the northern and eastern proto-Arabian–Nubian Shield, with those in the east below sea level, and connected to an ocean. Most pre-650 Ma rocks in the ANS are parts of Tonian– middle Cryogenian island arc terranes (Fig. 10.15) that developed in the Mozambique Ocean between rifted blocks of the Rodinia supercontinent and other cratons (Stoeser and Camp 1985; Genna et al. 2002; Johnson and Woldehaimanot 2003; Stoeser and Frost 2006). The oldest Neoproterozoic crust is in the Tokar/Barka– Asir terrane, the Jiddah–Haya terrane, and the Hijaz–Gebeit terrane. Although the crust in Eritrea is divided by many workers into many individually named domains (Drury and De Souza Filho 1998), the composite term ‘Tokar/Barka terrane’ is used here for convenience. The terranes converged

and amalgamated during the middle Cryogenian along the Barka and Bir’ Umq– Nakasib sutures (780–750 Ma) (Fig. 10.14), resulting in a core grouping of terranes. By 700 Ma, the Midyan–Eastern Desert terrane had collided and amalgamated with the western arc terranes of the ANS (e.g., Stoeser and Frost 2006). The terranes converged and amalgamated because of intraoceanic subduction-driven arc–arc and ultimately arc–continent collisions. Most terrane boundaries are high-strain shear zones that commonly contain dismembered ophiolites (Berhe 1990) and refolded recumbent folds. The shear zones are widely interpreted as sutures that formed at the time of terrane amalgamation, although identification of the shears zones as sutures has been challenged (Church 1991), and some shear zones are younger strike-slip shear zones that modified or reworked original sutures (e.g., Kusky and Matsah 2003). The Sol Hamid–Allaqi–Heiani suture (Stern et al. 1989; Stern 1994; Abdelsalam and Stern 1996) separates the South Eastern Desert (SED) (Gerf) terrane on the north (Kröner et al. 1987a; Greiling et al. 1994; Shackleton 1994;

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Fig. 10.15 ANS tectono-stratigraphic terranes, showing protolith ages in the NED; CED; SED (after Johnson et al. 2011)

Abdelsalam and Stern 1996; Kusky and Ramadan 2002) from the 830 to 700 Ma Gebeit and Gabgaga terranes on the south. It strikes roughly east, but swings to the south in its middle section as it merges with and is overprinted by the NNE-trending Hamisana shear zone (de Wall et al. 2001). Relatively high-grade gneissic rocks of the SED terrane are interpreted as island arc and ophiolitic nappes, uplifted as they were thrust south over the Gebeit and Gabgaba terranes (Greiling et al. 1994; Kusky and Ramadan 2002). The Gabgaba terrane contains an island arc assemblage (El-Nisr 1997), including metavolcanic rocks and bands of marble, interpreted as deformed shallow-water carbonates that originally fringed the arc volcanics (Greiling et al. 1994).

10.5

Tectonic Models of the ANS

10.5.1 Earlier Models of Evolution of ANS The presence of relatively well-preserved ophiolites in the ANS encouraged many authors to propose plate tectonic models for the evolution of the Shield. Several suites of ophiolitic blocks, volcanic, and volcaniclastic rocks with island arc affinities and granitoid plutons were described. Similarities in the rock assemblages and tectonic evolution between the Arabian and Nubian shield forced some authors to take into consideration the evolutionary models of the

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Arabian Shield when dealing with the Nubian shield. Nearly two-thirds of the rocks exposed in the Arabian Shield consist of volcanic, volcaniclastic, sedimentary, and plutonic rocks that make up belts of deformed and metamorphosed rocks, identified as magmatic arcs. Several tectonic models for the evolution of the ANS have been proposed and are described in Sects. 10.4.1.1–10.4.1.6.

10.5.1.1 A Single Immature Island Arc Evolution (Island Arc-Accretion Model) Greenwood et al. (1973) first suggested that the southern part of the Shield was developed from an oceanic crust in an intraoceanic island arc environment. Greenwood et al. (1976) assumed a northwestern dipping subduction zone situated southwest of the exposed part of the Arabian Shield. 10.5.1.2 Successive Formation and Accretion of Ensimatic Island Arcs Model Gass (1981) modified the single-arc model and recognized that the shield consists of several late Proterozoic, island arc systems (as many as seven arcs) once separated by oceanic lithosphere, that had been swept together, along with their sedimentary aprons and occasional slices of oceanic lithosphere, to form new continental crust (cratonized) between 1100 and 500 Ma ago (Bakor et al. 1976; Frisch and Al Fig. 10.16 Cartoon depicting stages in the development of the ANS. a depicts the situation in the lower Pan-African with many immature arc systems. By middle Pan-African times b, the arcs have matured and coalesced but have not attained continental dimensions. By upper Pan-African times c, the arcs had coalesced into continents, but these still overlay subduction zones and magmatic activity had calc-alkaline affinity. Figure d depicts the post Pan-African (500–600 Ma) situation. When the continent was fully developed, subduction had ceased, and magmatism was peralkaline and of within-plate affinity (after Gass 1981)

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Shanti 1977). Figure 10.16 represents a cartoon depicting the proposed stages in the development of the crystalline basement in NE Africa and Arabia. The lower Pan-African is illustrated with having many immature island arc systems. By middle Pan-African times, the arcs had matured and coalesced, though have not attained continental dimensions. By upper Pan-African times, the arcs had coalesced into continents, but were still underlain by subduction zones, and the magmatic activity was calc-alkaline. The theme of the post-Pan-African (500–600 Ma) included the fully developed continent, cessation of subduction, and magmatism of peralkaline within plate affinity.

10.5.1.3 Repeated Rifting of an Older Sialic Crust Garson and Shalaby (1976) were the first to recognize the ophiolitic nature of the widespread mafic-ultramafic complexes in the Eastern Desert. They suggested that the Arabo-Nubian Shield is developed by successive opening and closure of a system of small marginal basins above a westward dipping subduction zone against an older African craton. 10.5.1.4 Microplate Accretion Model In this model the Arabian Shield is subdivided into five distinct microplates (terranes) that are separated into four ophiolitebearing suture zones (Stoeser and Camp 1985). Three are

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intraoceanic island arc terranes of ensimatic character (Asir, Hijaz, and Midyan in the western Shield), whereas the two terranes Afif and ArRayn in the eastern Shield are continental (Fig. 10.10). Four suture zones of two types separate the above terranes: (1)Yanbu and Bir Umq sutures, formed by island arc—island arc collision; and (2) The Nabitah and Al Amar sutures, formed by arc-continent collision.

10.5.1.5 Reconstruction of the ANS Model Vail (1983, 1985) described dismembered ophiolites and ophiolitic melange in distinct shear zones that mark sutures between previously separate crustal blocks with the sutures of the Red Sea Hills that continue into Saudi Arabia. He recognized comparable terranes in the Nubian Shield and established coeval tectonic events across the Red Sea closure line. Vail (1985) reconstructed the Shield in North Africa and Arabia along three analogous suture zones: Yanbu—Sol Hamed suture, Bir Um Nakasib suture, and Asir-Baraka suture. Kröner et al. (1987a) delineated five distinct terranes separated by three sutures: the Gerf terrane in the north—a part of the crustal block including most of the Eastern Desert; the Gebeit terrane, bounded by Onib—Sol Hamed and Nakasib Amar sutures; the Haya terrane—the region between the Nakasib-Amur and Baraka sutures; the Tokar terrane—a wedge-shaped area of the southern Red Sea Hills east of the Baraka suture; and the Gabagba Terrane—the region west of Hamisana shear zone and south of the Allaqi Heiani ophiolite belt (Fig. 10.17). 10.5.1.6 Pan-African Continental Margin Model El Gaby et al. (1988) identified three main groups of rock units within the Nubian Egyptian shield, according to their space and time relationships (Fig. 10.18). A. Pre-Pan-African rocks (Infrastructure): These are Archaean to early Proterozoic granites, gneisses, and schists and their mylonitized and remobilized equivalents, cropping out at Gabal Uweinat in the southwestern corner of Egypt. In the Central and South Eastern Deserts, these infrastructures crop out in gneisses domes. B. Pan-African rock association (Suprastructure): These are formed in two successive stages: B.1 The ophiolites and island arc stage association: This comprises a tectonically lower ophiolite sequence and a tectonically upper island arc association. Ophiolites always occur as allochthonous and dismembered ultrabasic to basic bodies. The island arc assemblage comprises a series of

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weakly metamorphosed calc-alkaline andesites, dacites, and volcaniclastics of comparable composition. B.2 The Cordilleran stage association: This stage is characterized by the onset of an active phase of calc-alkaline activity. The Cordilleran stage associations comprise penecontemporaneous: a—Subduction related, Andean-type, calc-alkaline magmatic rocks (about 680–570 Ma ago). This magmatic activity is manifested in the emplacement of syn- to late-orogenic calc-alkaline granites of tonalitic to ideal granitic composition and in the eruption of their volcanic and subvolcanic analogues—the Dokhan Volcanics and Post-Hammamat Felsites. b—Molasse-type Hammamat sediments These were essentially derived from the erupting continental margin calc-alkaline volcanics accumulated in the intermontane basins between the raised geanticlines. c—Intrusive, mantle derived ultrabasic to basic intrusions, commonly fresh peridotites, gabbros, and diorites. d—Phanerozoic alkaline rocks: subalkaline to peralkaline A-type granite bodies.

10.5.2 Recent Models of ANS Evolution: Arc-Accretion Model The ANS is an accretionary orogen at the northern end of the EAO (Cawood et al. 2009), and the timing of accretion is extensively described by Stoeser and Camp (1985) and Johnson and Woldehaimanot (2003). Its western boundary is a contact with pre-Neoproterozoic crust belonging to the Saharan Metacraton and Congo-Tanzania Craton. Further south, the ANS is in contact with the Congo-Tanzania Craton and Mozambique Belt. A southeastern ANS margin is recognized as a contact with pre-Neoproterozoic gneisses belonging to a crustal block referred to as Azania. An eastern ANS margin, long a topic of debate, is suspected beneath Phanerozoic cover in central Arabia, and a northwestern margin with late Mesoproterozoic crust is suspected in Sinai. In general, the ANS is composed of a series of juvenile island arc terranes (Stern and Johnson 2010; Johnson et al. 2011; Robinson et al. 2015a, b) accreted during the Cryogenian and early Ediacaran (Johnson et al. 2011; Cox et al. 2012) with at least one pre-Neoproterozoic crustal block (Whitehouse et al. 2001a; Collins and Windley 2002; Hargrove et al. 2006). The ANS is modeled as a collage of oceanic island arcs that evolved during the Neoproterozoic above a multiplicity of subduction zones in the Mozambique Ocean. Recent reviews of the orogenic history of the shield (Fritz et al. 2013; Johnson 2014) provide comprehensive information

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Terrane Accretion Within the Arabian-Nubian Shield

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Fig. 10.17 Tectonic sketch map of southern Eastern Desert and Red Sea Hills showing ophiolite belts and suggested terranes and possible margin of the African craton (modified after Kröner et al. 1987a)

about the deposition, deformation, and tectonic evolution of the region. The Mozambique Ocean was a large basin on the margin of supercontinent Rodinia located between the crustal blocks of India, the Saharan Metacraton, and the Congo-Tanzanian Craton. ANS arc formation occurred over a *300 million year period encompassing supercontinent Rodinia breakup and the assembly of supercontinent Gondwana. The oldest rocks associated with this cycle are about 870 million years old and are found in eastern Sudan and SE Arabia (Whitehouse et al. 2001a; Collins and Windley 2002; Hargrove et al. 2006). Some of the oldest rocks are ophiolites, which testify that formation of ANS continental crust began with formation of oceanic crust by seafloor spreading, followed by the development of

subduction zones and Island arcs (Stern and Johnson 2010; Johnson et al. 2011). The earliest magmatic phases in the ANS include emplacement of thick oceanic tholeiite sequences in intraoceanic settings followed by 950–650 Ma bimodal volcanism with island arc chemistry. These phases involved large volumes of mafic crust and lithospheric mantle within the ANS and additional crustal overprinting (Bentor 1985; Stein and Goldstein 1996; Stein 2003). The upwelling plumes possibly provided the lithophile elementenriched sources for later calc-alkaline (640–590 Ma) and alkaline (590–550 Ma) magmatism (Robinson et al. 2014). The various island arcs collided, and these tectonic terranes were sutured during the time 780 to 620 Ma to form an increasingly broad and thick nucleus of juvenile continental

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Fig. 10.18 Cartoon illustrating the tectonic evolution of the Eastern Desert of Egypt. a island arc stage. b overthrusting of back-arc ophiolites and island arc volcanics over the old continental margin (after El Gaby et al. 1988)

crust. This thickening was associated with the formation of several suture zones, marked by obduction of ophiolites and intense deformation (Johnson et al. 2011; Cox et al. 2012). Crustal thickening was also accompanied by melting and magmatic fractionation of mafic magmas that ponded deep in

the crust. These melts rose upward to be emplaced as granitic plutons. Magmatism during this episode is characterized by tholeiites and calc-alkaline suites (Robinson et al. 2014). The arcs have tholeiitic to calc-alkaline, and locally MORB chemistry and include basalt, boninite, andesite and rhyolite,

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Terrane Accretion Within the Arabian-Nubian Shield

voluminous volcaniclastic and sedimentary rocks, and large amounts of diorite, tonalite, trondhjemite, and granodiorite. Oceanic crust in the shield, represented by ophiolite complexes, ranges from 845 to 675 Ma (Johnson 2014); arc assemblages range from *870 to 615 Ma. The oldest arc rocks in any given terrane are referred to middle Cryogenian, but Tonian rocks are locally preserved in the central and southern ANS, middle to late Cryogenian rocks is found in the north, and late Cryogenian-Ediacaran units crop out in the far

Fig. 10.19 Structural and metamorphic map of the ANS (after Fritz et al. 2013) showing tectonotratigraphic terranes, suture zones, the boundary between eastern and western arc terranes in the Arabian Shield (after Stoeser and Frost 2006), and boundaries between the ANS and flanking older crustal blocks. Arrows show displacement trajectories and sense-of-shear during transpressive orogenic phases in the region (after Johnson 2014)

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east and west (Johnson 2014). Arc amalgamation and suturing occurred between *780 and 600 Ma, and accretion between the ANS and the Saharan Metacraton, reflecting terminal collision of the ANS and western Gondwana blocks, occurred *650–580 Ma. Metamorphic grades in the ANS range from granulite facies in the south to greenschist facies over much of the north and developed during periods of transpressional east-west shortening, north-south extension, strike-slip shearing, and tectonic escape (Fig. 10.19).

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Johnson et al. (2013) stated that the growth of the Arabian Shield can be viewed as crustal development within a supercontinental cycle. Stages of growth can be recognized, made up of many interrelated events. These include the development of arcs across the shield; periodic development of fore-arc and back-arc ophiolites; the surge in post-amalgamation basin development, following middle Cryogenian terrane convergence and suturing; the transition from I-type arc-related magmatism to A-type, within-plate, extensional magmatism; episodes of suturing and shearing; and periodic exhumation and uplift reflecting extension, tectonic escape, and orogenic collapse. Genna et al. (2002) and Johnson and Woldehaimanot (2003) presented tectonic models that are similar in broad outline but differ in detail and emphasis. Blasband and others (2000) and Collins and Pisarevsky (2005) and others discussed the tectonic history of the Arabian Shield within the context of the global history of Gondwana amalgamation, and Stoeser and Frost (2006) compiled isotopic data that constrain tectonic speculation. The discovery of ophiolite assemblages in the late Proterozoic accretionary terrain of the ANS has led to a number of evolutionary models that envisage collision of intraoceanic island arcs and exotic microcontinental fragments after closure of several marginal basins (Bakor et al. 1976; Gass 1981; Al Shanti and Gass 1983; Kröner 1985; Stoeser and Camp 1985; Pallister et al. 1987). In these models, the ophiolites are envisaged as remnants of back-arc oceanic crust, an interpretation-based largely on the LIL-enriched geochemistry of their pillow basalts and sheeted dyke complexes (Bakor et al. 1976; Kröner 1985; Pallister et al. 1988). They were obducted, rather than subducted, in line with the current theory that back-arc basins reach maximum ages of about 40 Ma before being closed. This implies that young, hot, positively buoyant oceanic lithosphere, or at least the upper part of it, is thrust over the neighboring terrains during basin closure, similar to obduction of the Semail ophiolite nappe in Oman (Searle and Stevens 1984). Alternatively, ophiolite obduction was part of a backthrusting process following collapse of the arc systems during collision (Silver et al. 1983; Silver and Reed 1988). The ophiolite complexes or their remnants in the ANS constitute well-defined belts, often associated with strong deformation, and have therefore been interpreted as marking the sites of major suture zones, along which the arc terranes collided (Stoeser and Camp 1985; Kröner 1985; Vail 1985; Shackleton 1986; Kröner et al. 1987a; Pallister et al. 1987; Stoeser and Stacey 1988). In the Eastern Desert of Egypt (ED), most of the ophiolite assemblages are part of extensive nappe complexes (Shackleton et al. 1980) that were thrust over continental margin-type sediments (El Ramly et al. 1984) or arc terrains

A. F. Osman and A.-R. Fowler

(Ries et al. 1983). Those ophiolite fragments which preserve their internal structures rather well were discovered early (e.g., Jabal al Wask and Jabal Ess, Saudi Arabia—Bakor et al. (1976), Shanti and Roobol (1979), Wadi Ghadir, Egypt —El-Sharkawy and El-Bayoumi (1979), El-Bayoumi (1980, 1983), and the Fawakhir ophiolite—Shackleton et al. (1980) and Nasseef et al. (1980); Sol Hamed and Onib—Hussein (1977), Fitches et al. (1983), Hussein et al. (1984)). Those strongly fragmented and further disrupted during post-collisional strike-slip deformation were only recognized in recent years as remnants of oceanic lithosphere. Throughout the Arabian–Nubian Shield, several deformed linear belts of ophiolitic rocks have been observed, and these were interpreted as sutures (e.g., Ries et al. 1983; Vail 1985; Abdelsalam and Stern 1996). The Al Amar, Yanbu, and Bir Umq belts represent examples of the ophiolitic linear belts in the late Proterozoic shield of Saudi Arabia (Fig. 10.20). These belts separate little-deformed domains that are distinguished from each other based on different petrology, geochemistry, and ages. This led several authors to interpret the Asir, Midyan, Afif, Ar Ryan, and Hijaz Domains as accreted terranes and helped to identify the linear belts as ophiolitic sutures (Vail 1985; Stoeser 1986; Johnson et al. 1987). Several sutures display significant strike-slip movement (Quick 1991; Abdelsalam and Stern 1996; Johnson et al. 2004), which appears to post-date the initial stages of the suturing (see below). Granodioritic intrusions were associated with the arc-accretion phase (e.g., Stoeser and Camp 1985; Stoeser 1986; Quick 1991). The Nabitah Belt in Saudi Arabia is the largest and most complicated feature that was formed during the arc-accretion phase in the ANS. It is some 1200 km long, 100–200 km wide, and trends north–south. The Nabitah Belt consists of ophiolites and sheared ophiolites in the form of steeply dipping mylonites and phyllonites (Quick 1991). The oldest folding phase in this belt indicates WNW– ESE to NW–SE compression (Quick 1991). This phase is overprinted by a phase of left-lateral transpression along north– south trending strike-slip faults which were active at 700– 650 Ma (Quick 1991). Similar features have been observed in Sudan and Egypt. The NE-SW trending Nakasib Suture in Sudan separates island arc terranes in the east from a pre-Neoproterozoic continental terrane in the west (Abdelsalam and Stern 1996). It displays relicts of an east–west compressional regime, which represents the suturing phase. The strike-slip movement overprints the earlier structures and represents the later stages of arc-accretion (Abdelsalam and Stern 1996). The activity on the sutures took place at c. 750–650 Ma. The strike-slip movement, related to the later stages of the arc-accretion event, started at c. 670 Ma (Abdelsalam and

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Fig. 10.20 Tectonic map of the ANS showing the locations and extents of terranes, sutures, and post-accretionary structures. Modified after Johnson and Woldehaimanot (2003), Hargrove et al. (2006) and Abdelsalam (2010)

Stern 1996). The recognition of Precambrian ophiolite suites and their dismembered remnants in association with intraoceanic island arc volcanic and plutonic terrains across much of the ANS of eastern Egypt, Sudan, Ethiopia, Yemen, western Saudi Arabia, and Sinai has been used by many authors to support the hypothesis of crustal accretion during late Proterozoic time (950–550 Ma). Reassembly of the various fragments provides a mosaic of Proterozoic microplates in a regular pattern in which at least five oceanic terrains, bounded by the remains of ophiolite belts, lie between remobilized continental plates to east and west. The ophiolite lineaments may represent ancient subduction zones, although there is some doubt and much discussion (e.g., Gass 1981; Vail 1983) as to whether this is so and in which direction they would be inclined. On the other hand, it is also

believed (Bentor 1985; Stern et al. 1984) that not all the ultramafic assemblages are ophiolites and certainly not necessarily subduction derived. Vail (1985) indicated that the isotopic age determinations are insufficient to prove a sequential history between or within the volcanic terranes. In the Arabian Shield, Stoeser and Camp (1985) considered the Midyan terrane to be youngest (700–600 Ma), the Hijaz terrane to have developed about 800–700 Ma, and the southern Asir terrane to contain the oldest arc assemblages at > 900–800 Ma. In Sudan, it has been postulated (Embleton et al. 1982) that the volcano-sedimentary sequences would be oldest in the Northern Red Sea Hills, about 720 Ma in the Central block (Vail et al. 1984), and youngest in the Southern Red Sea Hills terrane. Stern and Hedge (1985) have proposed a threefold division of the Eastern Desert-Northern Red Sea Hills block in

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which the basement becomes younger to the north (from about 765 Ma in the south to 540 Ma in the north, as indicated by the tectonic igneous events). The pre-cratonic elements were brought together during a collision that occurred about 690 Ma ago producing a very thick turbidite succession (Delfour 1979; Delfour et al. 1982). The rocks are more distal and siliciclastic than the schist in other basins, which is commonly composed of tuff and cinerite. This difference led some authors to interpret the Abt Schist as an original accretionary wedge sediment body (Laval and Le Bel 1986). Volcanic arcs and ophiolitic assemblages typical of a subduction context indicate that convergence occurred in an oceanic environment (Camp 1984; Laval and Le Bel 1986; Pallister et al. 1987). However, few kinematic elements are known for this first period, and they cannot be extrapolated (Johnson 1996). The ocean-closure mechanisms are still poorly understood, but the main ones are the Al Ays, Abt, Halaban, Jiddah, Bahah, Baish, and Hali groups. Several authors (e.g., Hadley and Schmidt 1980; Jackson and Ramsay 1980; Delfour 1980a) proposed classifications that assume two or three major tectonic events occurred during this first long period. The basement underlying these early sedimentary formations is either poorly defined or not defined at all. Stacey and Hedge (1984) called it early Proterozoic, but other authors merely considered it as an ‘older basement’ of gneiss and migmatite or ophiolites (e.g., Delfour 1980b; Jackson and Ramsay 1980).

10.6

tectonic significance. Suture zones, the contacts between amalgamated terranes, are the most prominent shear zones in the ANS. Some of the shear zones were created during brittle-ductile deformation associated with transpressive collision and mark the sites of amalgamation between the tectono-stratigraphic terranes that make up the ANS. Other structures are brittle-ductile transcurrent faults that displace and dislocate the amalgamated terranes and were created during a process of continental dispersal or escape at the time of final convergence between eastern and western Gondwana. Some structures have complex histories that reflect early shortening and suturing and later reactivation by strike-slip. Other structures reflect strong shortenings perpendicular to their strikes. Orthogneiss and paragneiss of various ages are associated with the shear zones, cropping out as linear belts of high-grade metamorphic rock and elongate antiformal domes along or on either side of major shear zones (Johnson 2012). The gneisses are originated during periods of terrane suturing as well as during later orogenic collapse and crustal exhumation. Gneisses that have old protoliths mostly represent the deformation and metamorphism of pre-existing (pretectonic) volcano-sedimentary assemblages and plutons during terrane suturing. Some gneisses are syntectonic plutons reflecting magmatism contemporary with arc deformation and terrane suturing. Other gneisses are the product of metamorphism, mylonitization, and uplift of pre-existing and contemporary plutonic rocks along Najd faults in the Arabian Shield.

Faults and Shear Zones in the ANS

Faults and shear zones are present throughout the ANS, and their distribution and trends reveal interesting patterns (Fig. 10.21). They vary from simple brittle fractures to complex zones of ductile deformation and vary in orientation and age. The faults and shear zones are fundamental elements in the structural architecture of the shield. The southern and extreme eastern parts of the shield are dominated by north-trending structures (folds and shear zones), NNW-trending structures, and NE-trending structures. The northerly structural grain of the southern part of the Arabian Shield continues into the southern part of the Nubian Shield in Eritrea and Ethiopia and farther south into southern Ethiopia and northern Kenya. It is the defining feature of what has been differentiated as the Southern ANS (SANS) (Johnson et al. 2011). The northern limit of SANS is the Ruwah and Ad Damm fault zones in the Arabian Shield (Fig. 10.21). North of these fault zones, the structural grain changes to NW and NE trends mentioned above, and the region is referred to as the Northern ANS (NANS) (Johnson et al. 2011). Within the various regions of the shield, the faults and shear zones tend to have similar geological histories; but in detail, each fault and shear zone have its own history and

10.7

Ophiolite Zones as Evidence for Sutures in the ANS

Ophiolites in orogenic systems are generally considered to mark suture zones between collided plates and/or accreted terranes. The geology of the spatially associated lithological units and the regional tectonic history provides significant information on the mode and nature of mechanisms of the incorporation of ophiolites in orogenic systems. Nicolas (1989) used a different approach in classifying ophiolites based on the tectonic settings of their emplacement. The three main types of ophiolites in Nicolas’ classification are (1) Ophiolites tectonically resting on continental passive margins (e.g., Semail in Oman, Papuan ophiolite in Papua-New Guinea), (2) ophiolites incorporated into the active continental margins of the Circum-Pacific belt (e.g., ophiolites in the Franciscan Complex in California), and (3) suture zone ophiolites occurring in continent–continent or arc–continent collision zones (e.g., ophiolites in the Alpine-Himalayan orogenic system, Caledonian ophiolites, Hercynian and Uralian ophiolites). However, Moores (1982) classified ophiolites as Tethyan versus Cordilleran, based

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251

Fig. 10.21 Trends of shear zones and prominent foliation strikes in the ANS plotted on a base map showing the principal tectono-stratigraphic terranes and suture zones. The map illustrates the continuity of structures across the Red Sea and differences in orientation that form the basis for division of the ANS into northern and southern sectors (after Johnson et al. 2011)

upon the presence or absence of a continental substrate (i.e., passive margin of a continental plate or fragment), arc volcanic edifices, and/or accretionary mélanges. In this classification, Tethyan-type ophiolites are commonly observed to rest on passive continental margins along tectonic contacts and are considered to have formed at mid-ocean ridges; Cordilleran type ophiolites are spatially and temporally associated with island arcs and arc volcanic edifices, volcaniclastic rocks, and accretionary mélanges and are

considered to have formed in convergent margin settings (i.e., fore-arc, arc, or intra-arc). Ophiolites in the ANS are ubiquitously deformed, with the consequence that typical ophiolite successions are not preserved at every occurrence. Of these, Jabal Ess is one of the most complete, and Jabal Tays is the least complete. The ophiolites developed over a 170 million-year period ranging from Bi’r Tuluhah (*845 Ma) and Bi’r Umq (*840 Ma) to Halaban (*675 Ma), respectively, the oldest and youngest

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Fig. 10.22 Simplified geological map of the ANS shield showing the distribution of Neoproterozoic ophiolite, volcanic arcs, juvenile Neoproterozoic, and pre-Neoproterozoic. Data sources Palliser et al. (1988), Johnson (2001), Reischmann (2000) and Stem (2002). Abbreviations for ophiolite names are as follows: H. Al-Heiani, Br. Barramiya, G. Cerf, Gh. Ghadir, H. Hagar I. Baruda, D. Dare Tebet, T.D. Tulum Zager, A. Arja, A.A. Al’As, B.T. Bir Tuluha, Bu. Bir Um, D.Z. Darb Zubaydak, H.B. Halahan, J.E. Jabal Ess, J.G. Jabal Ghurrab, M. Mophia Th, Jabal Thurwah, Na, Natura, T.L. Tathlith S. Hamisana, 1. Ingessana, K. Keral. OSH, Onib Sol Hamed

ophiolites known in the entire ANS. Ophiolites in this mosaic occur in curvilinear belts (Fig. 10.22) with distinctive lithotectonic assemblages, age distributions, and chemical fingerprints, indicating discrete pulses of oceanic crust generation and its incorporation at the continental margins of early ANS throughout the late Proterozoic (Coleman 1984). These ophiolite-bearing belts have been identified as suture zones that mark the sites of tectonic amalgamation of

disparate terranes, predominantly of volcanic arc origin (Berhe 1990; Stern et al. 1990). However, the original tectonic setting and the root zones of these Proterozoic ophiolites are uncertain, mainly because of the reworking of the shield crust during multiple deformational episodes (Church 1988). The ophiolitic belts associated with suture zones in the ANS comprise Neoproterozoic oceanic-arc plutonic,

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Terrane Accretion Within the Arabian-Nubian Shield

volcano-sedimentary, and ophiolitic rocks that record some of the earliest magmatic and tectonic events of the East African Orogen (Fig. 10.22). Hargrove et al. (2006) have suggested that the 40 km-thick crust beneath much of the Arabian Shield is juvenile (mantle derived) and that most of it was extracted from depleted mantle during the interval *740–830 Ma. The ANS is an outstanding natural laboratory for studying the origin and evolution of juvenile continental crust: it arguably contains the best-preserved juvenile continental crust of Neoproterozoic age (1000– 542 Ma) (Gradstein et al. 2004). Neoproterozoic ophiolites, ranging in age from c. 870 Ma to c. 627 Ma, occur in several discrete suture and/or fault zones within the ANS and display a record of rift, drift, seafloor spreading, and collision tectonics during the evolution of the East African Orogen (Dilek and Ahmed 2003). The ophiolites in the Arabian Shield within the Yanbu and Bir Umq suture zones in the west are among the oldest (870–740 Ma). They locally show a Penrose-type complete pseudostratigraphy and have chemical compositions typical of modern fore-arc oceanic crust. They are spatially associated with coeval and younger volcanic arc assemblages and were incorporated into the Arabian Shield during a series of collisional events that amalgamated these ensimatic arc terranes. The ophiolites of the Hulayfah-Ruwah suture zone in the central Arabian Shield are coeval with and/or slightly younger (c. 843–821 Ma) than the ophiolites in the west and probably developed in a rifted ensimatic arc system that evolved as a volcanic archipelago near the Afif continental plate. Younger ophiolites (*694 Ma) of the Halaban and A1 Amar suture zones in the eastern Arabian Shield were incorporated into a subduction-accretion complex that evolved at the Andean-type active margin along the eastern edge of the Afif continental plate. The Halaban suture zone ophiolites represent fore-arc oceanic crust, whereas the A1 Amar suture zone ophiolites are scraped-off fragments of Mozambique ocean floor, seamounts and/or ocean island(s); the Abt Schist between them corresponds to a Franciscan-type accretionary prism of the ‘Halaban’ subduction zone. The incorporation of these ophiolites and the continental plates (Afif and Ar Rayn) into the Arabian Shield during 640–620 Ma marks a major shift in the direction of convergence (from northerly to westerly) during the assembly of the shield and distinct episodes of continental collisions during closure of the Mozambique Ocean. The ophiolites of the Nabitah-Hamdah fault zone within the Asir terrane are the youngest (c. 627 Ma) in the shield, post-collisional in origin, and display mid-ocean ridge basalt chemical affinity and represent Ligurian-type oceanic crust developed in an intracontinental pararift zone. The ophiolite tectonics of the Arabian Shield indicate an eastward progression of continental growth through time as the East

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African Orogen was built during the late Neoproterozoic, following the breakup of Rodinia. The recognition of ophiolite complexes is now widespread throughout ANS. They have been described in the Eastern Desert of Egypt (Garson and Shalaby 1975; El-Sharkawi and El-Bayoumi 1979; Gass 1981; Shackleton et al. 1980; Ries et al. 1983, etc.) and in Sinai (Shimron 1980), in eastern Sudan (AbdelMagid and AbdelRahman 1981; Embleton et al. 1982; Hussein et al. 1982). Several ophiolites are known in Ethiopia (Kazmin 1976; Kazmin et al. 1979; Warden et al. 1982) and also in Kenya (Vearncombe 1983). They are also known in Saudi Arabia, where they form extensive belts in the Proterozoic shield, and many have been studied in great detail (Bakor et al. 1976; Al Shanti and Mitchell 1976; Shanti and Roobol 1979; Gass 1981; Shanti 1982; Greenwood et al. 1980, 1982). Because the ophiolites occur along suture zones, their structure and geochronology are important constraints on the history of suturing (Church 1988). The term ‘suture’ refers to zones along which oceans and back-arc basins have closed (Burke et al. 1977). The closing of the oceanic basin is accompanied by a full range of structures, usually localized along linear zones of high strain. Suture is often represented at the surface by an orogen or mountain range. The accreted fragments may be continents, microcontinents, or island arc/back-arc basin complexes. Terrane accretion in the ANS took place along arc–arc sutures (Fig. 10.23) developed between 800 and 700 Ma (Stoeser and Camp 1985; Pallister et al. 1988; Ayalew et al. 1990; Kröner et al. 1992). The ANS was subsequently emplaced between the continental blocks of east and west Gondwana at 750–650 Ma along arc-continent sutures (Stern 1994). Based on the structural styles, ages, and tectonic settings, the thrust belts, fold belts, and strike-slip fault systems (deformational belts) in the ANS (Abdelsalam and Stern 1996) (Table 10.1) are divided into: i. Sutures, including those separating individual arc terranes (arc–arc sutures), and those separating the ANS from the pre-Neoproterozoic continental blocks to the east and west (arc-continental sutures); and ii. Post-accretionary structures, including north-trending shortening zones, and major northwest trending, sinistral and minor northeast trending, dextral strike-slip faults.

10.7.1 Arc–Arc Sutures Many studies have been made to connect ophiolites in the ANS into major arc–arc sutures, separating terranes of presumably different ages (Bakor et al. 1976; Shackleton 1979;

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Fig. 10.23 Sketch tectonic map of the ANS (after Abdelsalam and Stern 1996)

Nassief et al. 1984; Duyvermann 1984; Vail 1985; Berhe 1990; Stern 1993; Shackleton 1994). The following alignment is adopted (Fig. 10.17):

a suture in the Eastern Desert of Egypt. The belts above (except for the Hafafit culmination) qualify as arc–arc sutures for the following reasons:

i. The Allaqi-Heiani-Onib-Sol Hamed-Yanbu (YOSHGAH) suture. ii. The Nakasib-Bir Umq suture. iii. The Baraka-Tulu Dimtu suture; and iv. The Adola-Moyale suture.

1. These belts are associated with zones of high strain. Major structures such as fold and thrust belts and/or strike-slip fault system which continue for hundreds of kilometres coincide with belts listed above as arc–arc sutures. 2. Fragments of ophiolite rocks occur within these high-strain zones. These ophiolitic fragments, in most cases, represent eroded folded nappes which had

In addition to the above sutures, the Hafafit culmination will also be discussed as a possible structure associated with

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Table 10.1 Orientations, ages, and structural styles of major sutures and post-accretionary structures in the ANS (after Abdelsalam and Stern 1996)

Sutures

Arc–Arc

Arc-Continental

Post-accretionary structures

Shortening Zones

NW faults

Deformational belt

Orientation

Age of deformation (Ma)

Structural style

YOSHGAH

E to NE

750–720

Early S to SE-verging ophiolitic nappes. Late E to NE-trending upright folds

Nakasib-Bir Umq

NE

800–750

Early SE-verging ophiolitic nappes. Late NE-trending upright folds

Baraka-Tulu Dimtu

N

820–760

N-trending sinistral transpression

Adola-Moyale

N

830–620

Early E-or W-verging ophiolitic nappes. Late N-trending upright folds, or N-trending sinistral transpression

Keraf

N

700–610

N-trending sinistral transpression

Kabus

NNE

Sekerr

N

820–620

W-verging nappe

Al Amar (?)

N

680–640

Early E-or W-verging ophiolitic nappe. Late N-trending folds

Nabitah (?)

N

720–680

Early N-trending folds. Late E-verging thrusts. or N-trending sinistral transpression

Hamisana shear zone N

N

660–610

Early N-trending upright folds. Late NE-trending dextral strike-slip faults

Oko shear zone

N to NW

700–560

Early N-trending upright folds. Late NW-trending sinistral strike-slip faults

Najd fault system

NW

630–530

Early dextral strike-slip faults & shear zones Late sinistral strike-slip faults & shear zones

travelled for distances from their corresponding sutures. Most of these ophiolitic fragments, however, lie within broad deformational belts, which are the manifestation of collision between terranes; and 3. These belts separate arc terranes with different pre-suturing history and sometimes with different ages, as exemplified by the Nakasib suture, which separates the 900–800 Ma Haya Terrane from the 830–720 Ma Gebeit Terrane (Fig. 10.1; Stern and Kröner 1993).

10.7.2 Arc-Continent Sutures The Pan-African Orogeny in Arabia and northeast Africa is interpreted to have resulted from the opening and closing of the Mozambique Ocean, which developed between east and west Gondwana (Stern 1994). The arcs associated with this ocean are preserved as the ANS, which is separated by tectonic boundaries from older crustal blocks to the east (Schmidt et al. 1979; Stacey et al. 1984; Stoeser and Stacey 1988) and west (Vail 1983, 1985, 1988; Abdelsalam and Dawoud 1991). These boundaries will be referred to as arc-continental sutures, as they juxtapose arc terranes with

E-verging ophiolitic nappe. W-verging nappes

continental blocks after consumption of marginal oceanic basins.

10.7.3 The Eastern Margin of the ANS U/Pb zircon ages and Pb and Sr isotopic data indicate that at least the southern part of the Afif Terrane (Fig. 10.17) incorporates pre-Neoproterozoic crust (Stacey et al. 1980; Stacey and Stoeser 1983; Stacey and Hedge 1984). In general, this boundary is poorly defined, however, the presence of *690–670 Ma arc volcanics of the Ar Rayn Terrane to the east of the Afif Terrane, is problematic in that it indicates the presence of Neoproterozoic crust to the east of what is supposed to be the western margin of East Gondwana. The Ar Rayn Terrane is interpreted as the leading edge of a continent further east (Schmidt et al. 1979; Fleck et al. 1980; Davies 1984) but isotopic studies by Calvez et al. (1984), Stacey and Stoeser (1983), and Stacey et al. (1984) indicate that the terrane is a Neoproterozoic magmatic arc lacking older basement. This led Stoeser and Stacey (1988) to suggest that the pre-Neoproterozoic basement exists further east in Oman and Yemen as demonstrated by isotopic data reported by Stacey et al. (1980) and Stacey and Stoeser

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(1983). The evolution of the Al Amar and Nabitah deformational belts will be discussed as possible exposures of the eastern margin of the ANS, on the basis that the Afif Terrane represents a microcontinent that was rifted from East Gondwana during the Neoproterozoic Pan-African Orogeny. Some attempts have been made to define the eastern boundary of the ANS to the south of the Arabian Peninsula. Berhe (1990) linked the Al Amar suture to the Adola-Moyale suture of southern Ethiopia. Stern (1994) inferred the boundary in Yemen to be somewhere to the west of the Aden group which gave a < 3.0 Ga TDM model Nd age (Stoeser et al. 1991) and further south to the east of the Adola-Moyale and Baragoi ophiolites, where it connects with the Mozambique Belt in northern Kenya (Fig. 10.17). Lenoir et al. (1994) concluded that the basement in northern Somalia is dominated by juvenile Neoproterozoic crust, whereas southern Somalia is underlain by highly reworked pre-Neoproterozoic material with little addition of juvenile material. This is further supported by Haider and Berhe (1995) who indicated that, based on model ages, the margins of east Gondwana occur along the Burr massif in southern Somalia. The high-grade gneisses around the Adola-Moyale suture gave ages comparable to those in other part of the ANS (Teklay et al. 1993). However, geochronological data for gneisses from eastern Ethiopia indicate involvement of pre-Neoproterozoic continental crust (Teklay et al. 1993). These data suggest that the eastern boundary of the ANS lies somewhere in central Somalia and further south to the east of the Adola-MoyaIe suture. The north-trending Al Amar suture developed between 680 and 640 Ma due to collision between the Afif Terrane in the west and the Ar Rayn Terrane to the east (Stoeser and Stacey 1988). The suture is defined by two north-trending discontinuous ophiolitic melange zones sandwiching the Abt schist belt (Al Shanti and Gass 1983). Models which have been proposed to explain the relationships between the Afif Terrane, the Abt schist, the Al Amar suture, and the Ar Rayn Terrane agree in interpreting the Afif Terrane as a pre-Neoproterozoic microcontinent and the Abt schist as an accretionary wedge on the leading edge of a major continent to the east (Al Shanti and Mitchell 1976; Nawab 1979; Schmidt et al. 1979; Al Shanti and Gass 1983; Stacey et al. 1984; Stoeser and Stacey 1988). Al Shanti and Gass (1983) identified two sets of co-axial north-trending folds from the Abt schist. This led Al Shanti and Gass (1983) to conclude that the two ophiolitic melange zones marking the Al Amar suture initially occurred as a sheet-like layer underlying the Abt schist. The exposure of the ophiolitic melange to the east and west of the Abt schist is due to folding of the two units into a north-trending synform. The Nabitah suture (Fig. 10.17) is a north-trending deformational belt extending for *1000 km in central Arabia. The belt is marked by ophiolitic fragments. Early models interpreted the Nabitah

A. F. Osman and A.-R. Fowler

belt as an arc–arc suture (Frisch and Al Shanti 1977; Schmidt et al. 1979). However, geochronological and isotopic data led Stoeser et al. (1984) and Stoeser and Stacey (1988) to suggest that the Nabitah suture has been the site of collision between arc terranes in the west (the Asir and Hijaz Terranes) and a microcontinent to the east (the Afif Terrane, Fig. 10.17) at *690–680 Ma. Agar (1985) and Stacey and Agar (1985) added that an Andean-type arc was developed on the western margin of the Afif Terrane between *720 and 685 Ma. Structural data from the Nabitah suture were interpreted in two models (Abdelsalam and Stern 1996): i. Agar (1985) described the structures associated with the Nabitah suture as north-trending, upright folds affecting a passive margin sedimentary group, which lies unconformably on an older gneissic basement. Progressive deformation culminated in the development of east-verging thrust. ii. Quick and Bosch (1989) and Quick (1991) interpreted the structures related to the Nabitah suture as resulting from sinistral transpression.

10.7.4 The Western Margin of the ANS The high-grade gneissic terrane to the west of the ANS is thought to be pre-Neoproterozoic in age and referred to as the Nile Craton (Rocci 1965), the Eastern Saharan Craton (Bertrand and Caby 1978), the Sahara-Congo Craton (Kröner 1977), or the Central Saharan Ghost Craton (Black and Liegeois 1993). The presence of a pre-Neoproterozoic (though highly reactivated during the Neoproterozoic) continental crust to the west of the ANS is based on U/Pb zircon ages (Wust et al. 1987; Kröner et al. 1987b; Sultan et al. 1990, 1992, 1993; Stern et al. 1994) and Pb, Sr, and Nd isotopic compositions (Harris et al. 1984; Dixon and Golombek 1988; Schandelmeier et al. 1988; Harms et al. 1990; Sultan et al. 1992; Stern et al. 1994). Vail (1983) proposed that the western boundary of the ANS is marked by a deformed sedimentary prism overlying the high-grade rocks of the Nile Craton and can be traced from the Sekerr region in northern Kenya, through the Ingessana region in east-central Sudan, to the Eastern Bayuda Desert in northern Sudan. Abdelsalam and Dawoud (1991) argued that the boundary in central Sudan should be assigned to the Kabus ophiolitic mélange, which lies at 500 km to the west of the Ingessana region. The contact between the ANS and the Nile Craton in northern Sudan was identified by Almond and Ahmed (1987) as the north-trending, sub-vertical Keraf suture (Fig. 10.17). The Keraf, Kabus, and Sekerr sutures show gross similarities in their tectonic setting but have different

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structural styles. The Keraf suture is outlined by north-trending upright folds in the north and north to north-northwest trending, sinistral, and strike-slip faults to the south. The Kabus suture is defined by an east-verging fold and thrust belt, where the stacking order indicates pre-Neoproterozoic crust thrust across an ophiolitic mélange, which was itself thrust across the volcano-sedimentary sequences of the ANS. The Sekerr suture is defined by a west-verging fold and thrust belt, where the imbricated ophiolite and volcano-sedimentary rocks were tectonically emplaced from east to west across the Tanzania Craton margin. This difference in structural style along the western margin of the ANS may be due to an overall southeast-northwest compression, which accompanied an oblique collision between east and west Gondwana (Abdelsalam et al. 1995). This non-orthogonal compression (i.e., maximum compression axis not at right angles to the orogen front) might have been resolved into a north-trending upright fold belt and north to north-northwest trending, sinistral strike-slip faults in the north, and east- or west-verging nappes to the south.

10.8

Post-accretionary Structures

The ANS was affected by regional Neoproterozoic deformations younger than the collision between arcs and distinct from collision between the ANS and Gondwana fragments. These younger deformations are typically localized along north or northwest-trending pure shear or simple shear zones. These structures are referred to as post-accretionary because they post-date the inferred arc–arc sutures in the ANS. These deformational belts are characterized by the development of north-trending, upright, and tight folds developed between *700 and 650 Ma (Stern et al. 1989, 1990; Abdelsalam 1994). Two of these zones are exposed in northeast Sudan (Fig. 10.17): the Hamisana Shear Zone (Stern et al. 1989, 1990; Miller and Dixon 1992) and the Oko Shear Zone (Almond and Ahmed 1987; Abdelsalam 1994). The north to northwest-trending Oko Shear Zone sinistrally offsets the Nakasib suture (Almond and Ahmed 1987; Almond et al. 1989; Abdelsalam 1994). The evolution of the Oko Shear Zone is poorly constrained between 700 and 560 Ma, which is broadly contemporaneous with that of the Hamisana Shear Zone (Abdelsalam and Stern 1993). It developed through two phases of deformation (Abdelsalam 1994). The early phase was characterized by the development of NNW-trending, tight, and upright folds. This was followed by the development of sub-vertical northwest-trending, sinistral strike-slip faults, which were initiated as a conjugate set of ductile shear zones. The most prominent of the northwest-trending strike-slip fault zones in the ANS is the Najd fault system in Arabia and

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Egypt. The *1200 km long, 300 km wide Najd fault system (Fig. 10.17; Moore 1979; Agar 1987) is defined by northwest-trending, sinistral strike-slip faults, which offset the north-trending Al Amar and Nabitah sutures and the northeast-trending Bir Umq and Yanbu sutures (Delfour 1979; Moore 1979; Davies 1984; Stacey and Agar 1985; Agar 1986, 1987; Stoeser and Stacey 1988; Stern 1985; Sultan et al. 1986; El-Gaby et al. 1988). Other northwest-trending strike-slip faults crop out in the Red Sea Hills of the Sudan (Abdelsalam 1994), southern Ethiopia (Berhe 1986, 1990; Bonavia and Chorowicz 1993; Alene and Barker 1993), and Kenya (Key et al. 1989; Mosley 1993). The cumulative displacement along the Najd fault system is *240 km (Agar 1987). The period of activity of the Najd fault system was thought to be 580–530 Ma (Fleck et al. 1980). However, Stacey and Agar (1985) suggested that the Najd fault system was active between *630 and 600 Ma as a dextral system and then as a sinistral system up to *530 Ma. Moore (1979) proposed that the Najd Fault System was the product of deformation of a brittle cover overlying a basement, comprising rigid blocks, which moved laterally relative to each other, while Schmidt et al. (1979), Fleck et al. (1980), Davies (1984) and Agar (1987) proposed that the Najd Fault System is related to collision between the ANS and a rigid indenter to the east of the Al Amar suture. Stern (1985) argued that the Najd fault system could not have been formed by continent–continent collision in eastern Arabia. Stern (1985) suggested an alternative model where the Najd system represents a set of transform faults developed in response to a major episode of extension in the northwestern part of the ANS. Abdelsalam (1994) suggested that the northwest-trending, sinistral strike-slip faults were developed by continuous east-west shortening deformation as zones of strain when the ANS collided with the Nile Craton in the west and the Ar Rayn microplate to the east at *670–610 Ma. Stern (1994) suggested that the northwest-trending faults were associated with escape tectonics due to collision between East and West Gondwana along the MB.

10.9

Post-amalgamation Basins of the NE Arabian Shield

The development and structure of late Neoproterozoic (650– 540 Ma) volcano-sedimentary and molasse basins in the northeastern part of the ANS record periods of uplift, erosion, extension, subsidence, compression, and strike-slip faulting that post-dated the completion of terrane amalgamation in the northern EAO and predated the initiation of a passive margin on the northern flank of the Gondwana supercontinent (Johnson 2003). These post-amalgamation basins are a distinctive component of the geology of the

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northern EAO in the ANS. The basins varying in size, in their present deformed states, from aggregates of basins extending over as much as 72,000 km2 to small isolated basins of 200 km2 (Johnson 2003). They are unconformable on the juvenile volcanic arc rocks that make up the late Neo proterozoic post-amalgamation sedimentary basins in the ANS. The basins are unconformably overlain by lower Paleozoic passive margin siliciclastic rocks, which were deposited on the EAO and the end-Neoproterozoic Gondwana supercontinent that resulted from East and West Gondwana convergence (Stern 1994) and are intruded by large volumes of late to post-orogenic granitoids. EAO post-amalgamation depositional basins are rare outside the ANS, and late Neoproterozoic events elsewhere in the orogen include the different processes of synorogenic ductile shearing and granulite-grade metamorphism in Madagascar (Martelat et al. 2000; de Wit et al. 2001).

10.10

Conclusions

Accretionary orogens form along continental margins where oceanic lithosphere has been subducted. They are major sites of continental growth that take place when materials are accreted to the overriding continental crust, while subduction continues (Cawood et al. 2009). The ANS is a classic example of an accretionary orogen, involving terrane amalgamation and accretion during subduction, leading to the progressive formation of superterranes and proto-continents, and finally achieving continental growth. Terrane studies are important evidence for the migration of parts of the Earth’s crust from one place to another distant place. Construction of terrane maps is a precondition for identifying ancient tectonic configurations that require, for example, knowledge of subduction polarity, convergence directions, etc. Most accreted terranes are autochthonous, i.e., are sedimentary or volcanic tectonic elements developed near to the site of terrane collision, such as accretionary prisms, fore-arc basins, and melange bodies. Terranes transported long distances to the site of accretion are allochthonous, or ‘suspect’ or ‘exotic’ terranes. These are highly varied and include oceanic plateaux, oceanic ridges, microcontinents and continental fragments, island arcs, etc. Where the present configuration of terranes reflects syn- and post-assembly deformation, the terranes are referred to as tectonostratigraphic terranes. Tectonos-tratigraphic terranes may be classified as stratigraphic; disrupted; metamorphic; volcanic arc, depending on their foremost characteristics. They form within oceans or by rift fragmentation of continents. Terranes can be identified by their distinctive geological history, stratigraphic successions, characteristics of their volcanics, and by geochronology, which differ markedly

from those of the surrounding crustal units. The terranes are separated from the surroundings by major faults (thrusts, strike-slip) but not by unconformities. All this evidence points to the origins of the terranes in a separate location or environment. The style of terrane accretion may involve obduction, over thrusting, subduction, underplating or simple suturing. Accretion commonly results in crustal thickening and uplift, and trench migration. The timing of amalgamation or accretion is constrained by post-amalgamation basins and stitching plutons. The Rodinia supercontinent existed at *1.23 Ga and was prceded by the Columbia supercontinent (2.0–1.8 Ga). Rodinia was assembled during the Grenville Orogeny (1.3– 0.9 Ga) and broke-up and was dispersed between 900 and 600 Ma, perhaps due to the activity of a mantle superplume that persisted from 830 to 650 Ma. Stages of breakup of Rodinia can be recognized in crustal arching, bimodal magmatism, rift sedimentary successions, and development of spreading ridges. Oceans that existed at the time of Rodinia breakup were Mozambique, Braziliano and Adamastor oceans. Gondwana and Laurasia were assembled from the dispersed fragments of Rodinia. The formation of the supercontinent Pangaea (300 to *180 Ma) by combination of Gondwana and Laurasia is the best studied and documented example of global scale accretion tectonics. Gondwana formed by amalgamation of East and West halves during closure of the Mozambique Ocean at about 550 Ma. East Gondwana comprised ANS, Madagascar, India, Antarctica and Australia. West Gondwana consisted of the majority of Africa, plus South America and a complex of cratons enclosed in the Pan-African/Brasiliano fold belts. East Gondwana assembly resulted from a complex set of events during 750–530 Ma in two stages: 750–620 Ma amalgamation of ANS arc terranes and collision of Madagascar-India-East Antarctica; followed by 570–530 Ma assembly of the ANS superterrane with Australia and East Antarctica. The West Gondwana assembly began around 900–700 Ma. Final amalgamation of East and West Gondwana (550–530 Ma) was preceded by the establishment of the short-lived Pannotia supercontinent. Three orogenies were key to the assembly of Gondwana: The East African Orogeny (800–650 Ma), the Kuunga Orogeny (c. 550 Ma) and the Brasiliano Orogeny (660–530 Ma). Collectively these have been referred to as the Pan-African Orogeny (PAO). The term ‘Pan-African’ refers to a sequence of tectono-thermal events at 500 *100 Ma within Africa and adjacent Gondwana elements. The term was broadened to include orogenic events of the same time range (950– 450 Ma) on a more global scale. In the intervening years, the Pan-African orogenic cycle has been more narrowly defined both spatially and temporally, such that it is possible to

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recognize individual orogenic events within Gondwana. The PAO comprised a series of deformation belts including the ANS and the Mozambique Belt (MB), the latter extending from East Antarctica through East Africa, up to the ANS. The PAO and Grenville Orogenies are the largest known orogenic systems on Earth. The East African Orogen marks the collision zone of East and West Gondwana. In its southern parts it records an oblique continent-continent collision and high-grade metamorphic reworking of a much older crust. In its northern part, in the ANS, it records stacking of juvenile arc terranes, followed by tectonic escape. The ANS is the result of a 300-million-year period of crustal growth, involving the accretion of juvenile subductionrelated magmatic arcs, periods of sedimentation and volcanism in basins developed on newly post-amalgamated arc terranes, crustal thickening, the emplacement of large amounts of granitic magma, periodic uplift and erosion, extension, and tectonic escape on transcurrent faults Much of the ANS grew by terrane accretion, during orthogonal and oblique convergence. It is the best example of a Neoproterozoic juvenile crustal province, possibly containing traces of older crustal (Rodinia) materials. The ANS extends 3500 km N-S and 1500 km E-W and is divided by sutures. The southern half is dominated by N-S trends. The northern half is more complex, but with a strong NW-SE structural grain. Subduction in the ANS began c. 870 Ma, with arc–arc convergence underway at 780 Ma. Terrane contacts include sutures, sutures reactivated as faults, and later faults. Most terranes were arc-related, though one of the easternmost terranes, Ad Dawadmi, represents an accretionary prism/fore-arc sequence; and the nearby Khida terrane is thought to be an older continental microplate. The number of terranes in the ANS is disputed, with as few as five and as many as fourteen being described. The ANS assembly involved numerous events: postamalgamation sedimentation and volcanism, crustal thickening, granite magmatism, uplift and erosion, tectonic extension and escape, and transcurrent faulting. The first stage of ANS began with its southern part, including areas south of the Nakasib-Bir Umq suture, and from Nakasib suture to the Yanbu-Allaqi-Heiani suture. These arc terranes collide between 830 and 710 Ma. The second stage involved formation of the Midyan—Eastern Desert terrane by collisions in the range 760–730 Ma, to form the ‘western arc terranes’. The third stage witnessed the collision of the western arc terranes with the earlier amalgamated terranes to form the proto-ANS an arc superterrane at 680–640 Ma. The Nabitah Orogeny marks the climax of this stage. The fourth and final stage involved the accretion of the proto-ANS with the Saharan craton along the Keraf suture, as a result of collisions in the east of the ANS of the Ad Dawadmi and Ar Rayn terranes (650–542 Ma).

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Tectonic models of the ANS have evolved. Recognition of ophiolites introduced the application of plate tectonic theory to the evolution of the ANS. Models for the ANS include: (1) single immature island arc; (2) successive accretion of island arcs; (3) repeated rifting of old sialic crust; and (4) microplate accretion. A continental margin model for ANS evolution was popular, involving a pre-Pan-African sialic infrastructure, over which ophiolites and arc volcanics were thrusted (a suprastructure). Following the emplacement of the suprastructure, continuing subduction resulted in Andean-type magmatism. A more recent model of originally entirely ensimatic setting prior to cratonization is better supported by current data. The ophiolite-decorated terrane boundaries within the ANS have encouraged interpretation of these as sutures. However, there is a rarity of Penrose-type pseudostratigraphy suggesting that these may not be true ophiolites. Many have the chemical characteristics of fore-arc oceanic crust. The sutures can be divided into arc–arc sutures (Allaqi-HeianiYanbu; Nakasib; Baraka-Tulu Dimtu; Adola-Moyale plus perhaps some ophiolitic mélange in the Eastern Desert) and arc-continent sutures (Al Amar; Keraf; Sekerr) Post-accretion structures include N-S to NW-SE trending intense strain zones (Hamisana and Oko) that have also been originally interpreted as shear zones. The timing of development of the Hamisana and Oko shear zone is poorly constrained to date somewhere between 700 and 560 Ma. In both zones the early structures are N–S to NNW-trending upright tight folds, followed by NW-SE trending shear zones. Fault and shear zone systems in the ANS are mainly N–S oriented in the southern parts and NW–SE oriented (with a minor NE-trending set) in the northern parts. The NW-trending sinistral Najd Fault System (NFS) is associated with elongated gneiss domes and pull-apart sedimentary basins. The Najd activity has been divided into a 630– 600 Ma stage of dextral shear, followed by a 600–530 Ma sinistral strike-slip shear history with *240 km sinistral displacement estimates. The origins of the NFS remain somewhat controversial with indentation tectonic and continental transform interpretations proposed. Post-amalgamation volcano-sedimentary basins were also developed in the ANS. These basins lie unconformably in the suture zones. They are intruded by large volumes of lateto post-orogenic granitoids.

References AbdelMagid AEM, AbdelRahman EM (1981) The basement geology of SE Sudan. 11th. Colloq. African Geol., Milton Keynes, Abstract, p 7 Abdelsalam MG (1994) The Oko shear zone: post-accretionary deformations in the Arabian-Nubian Shield. J Geol Soc Lond 151:767–776

260 Abdelsalam MG (2010) Quantifying 3D post-accretionary tectonic strain in the Arabian-Nubian Shield: Superimposition of the Oko Shear Zone on the Nakasib Suture, Red Sea Hills, Sudan. J Afr Earth Sci 56:167–178 Abdelsalam MG, Dawoud AS (1991) The Kabus ophiolitic melange, Sudan, and its bearing on the W boundary of the Nubian Shield. J Geol Soc Lond 148:83–92 Abdelsalam MG, Stern RJ (1993) Timing of events along the Nakasib suture and the Oko shear zone, Sudan. In: Thorweihe U, Schandelmeier H (eds) Geoscientific Research in Northeast Africa. Balkema, Rotterdam, pp 99–103 Abdelsalam MG, Stern RJ (1996) Sutures and shear zones in the Arabian-Nubian Shield. J Afr Earth Sci 23:289–310 Abdelsalam MG, Stern RJ, Schandelmeier H, Sultan M (1995) Deformational history of the Keraf Zone in northeast Sudan, revealed by Shuttle Imaging Radar. J Geol 103:475–491 Abdelsalam MG, Stern RJ, Copeland P, Elfaki EM, Elhur B, Ibrahim FM (1998) The Neoproterozoic Keraf Suture in NE Sudan: sinistral transpression along the eastern margin of West Gondwana. J Geol 106:133–147 Abdelsalam MG, Liegeois JP, Stern RJ (2002) The Saharan Metacraton. J Afr Earth Sci 34:119–136 Abdelsalam MG, Abdeen MM, Dowaidar HM, Stern RJ, Abdelghaffar AA (2003) Structural evolution of the Neoproterozoic Western Allaqi-Heiani suture, southeastern Egypt. Precamb Res 124:87– 104 Abd El-Wahed M, Hamimi Z (2021) The Egyptian Nubian Shield Within the Frame of the Arabian–Nubian Shield. In: Hamimi Z, Arai S, Fowler AR, El-Bialy MZ (eds.) The Geology of the Egyptian Nubian Shield. Regional Geology Reviews. Springer, Cham. https://doi.org/10.1007/978-3-030-49771-2_2 Agar RA (1985) Stratigraphy and paleogeography of the Siham group: direct evidence for a late Proterozoic continental micro-plate and active continental margin in the Saudi Arabian Shield. J Geol Soc Lond 142:1205–1220 Agar RA (1986) The Bani Ghayy Group; sedimentation and volcanism in pull-apart grabens of the Najd strike-slip orogen, Saudi Arabian Shield. Precamb Res 31:259–274 Agar RA (1987) The Najd fault system revisited: a two-way strike-slip orogen in the Saudi Arabian Shield. J Struct Geol 9:41–48 Alene M, Barker AJ (1993) Tectono-metamorphic evolution of the Moyale region, south Ethiopia. Precamb Res 62:271–283 Ali KA, Stern RJ, Manton WI, Kimurab J-I, Khamees HA (2009) Geochemistry, Nd isotopes and U-Pb SHRIMP zircon dating of Neoproterozoic volcanic rocks from the Central Eastern Desert of Egypt: New insights into the *750 Ma crust-forming event. Precamb Res 171:1–22 Ali KA, Azer MK, Gahlan HA, Wilde SA, Samuel MD, Stern RJ (2010) Age constraints on the formation of Neoproterozoic ophiolites along the Allaqi-Heiani Suture, South Eastern Desert of Egypt. Gondw Res 18:583–595 Almond DC, Ahmed F (1987) Ductile shear zones on north Red Sea Hills, Sudan and their implication for crustal collision. Geol J 22:175–184 Almond DC, Darbyshire DPF, Ahmed F (1989) Age limit for major shearing episodes in the Nubian Shield of northeast Sudan. J Afr Earth Sci 9:489–496 Al-Saleh AM, Boyle AP, Mussett AE (1998) Metamorphism and 40Ar/39Ar dating of the Halaban ophiolite and associated units: evidence for two-stage orogenesis in the eastern Arabian shield. J Geol Soc Lond 155:165–175 AI Shanti AMS, Gass IG (1983) The upper Proterozoic ophiolite melange zone of the easternmost Arabian Shield. J Geol Soc Lond 140:867–876

A. F. Osman and A.-R. Fowler Al Shanti AMS, Mitchell AHG (1976) Late Precambrian subduction and collision in the Al-Amar-Idsas region, Arabian Shield, Kingdom of Saudi Arabia. Tectonophys 30:41–47 Ayalew T, Bell K, Moore JM, Parrish RR (1990) U-Pb and Rb-Sr geochronology of the west Ethiopian Shield. Geol Soc Am Bull 102:1309–1316 Bakor AR, Gass JG, Neary CR (1976) Jabal al Wask, Northwest Saudi Arabia: An Eocambrian back-arc ophiolite. Earth Planet Sci Lett 30:1–9 Bauernhofer AH, Hauzenberger CA, Wallbrecher E, Muhongo S, Hoinkes G, Mogessie A, Opiyo-Akech N, Tenczer V (2009) Geochemistry of basement rocks from SE Kenya and NE Tanzania: indications for rifting and early Pan-African subduction. Int J Earth Sci 98:1809–1834 Be’eri-Shlevin Y, Eyal M, Eyal Y, Whitehouse MJ, Litvinovsky B (2012) The Sa’al volcano-sedimentary complex (Sinai, Egypt). A latest Mesoproterozoic volcanic arc in the northern ArabianNubian Shield. Geology 40:403–406 Bentor YK (1985) The crustal evolution of the Arabo-Nubian massif with special reference to the Sinai Peninsula. Precamb Res 28:1–74 Berhe SM (1986) Geologic and geochronological constraints on the evolution of the Red Sea-Gulf of Aden and Afar Depression. J Afr Earth Sci 5:101–117 Berhe SM (1990) Ophiolites in Northeast and East Africa: implications for Proterozoic crustal growth. J Geol Soc Lond 147:41–57 Bertrand J-M, Caby R (1978) Geodynamic evolution of the Pan-African orogenic belt: a new interpretation of the Hoggar Shield (Algerian Sahara). Geolog Rundsch 67:357–388 Black R, Liegeois J-P (1993) Cratons, mobile belts, alkaline rocks and continental lithospheric mantle: The Pan-African testimony. J Geol Soc Lond 150:89–98 Blasband B, White SH, Brooijmans P, de Boorder H, Visser W (2000) Late Proterozoic extensional collapse in the Arabian-Nubian Shield. J Geol Soc Lond 157:615–628 Bogdanova SV, Pisarevsky SA, Li ZX (2009) Assembly and breakup of Rodinia (Some Results of IGCP Project 440). Stratig Geol Correl 17(3):259–274 Bonavia FF, Chorowicz J (1993) Neoproterozoic structures in the Mozambique Orogenic belt of southern Ethiopia. Precamb Res 62:307–322 Brito Neves BB, Cordani UG (1991) Tectonic evolution of South America during the Late Proterozoic. Precamb Res 53:23–40 Burke K, Dewey JF, Kidd WSF (1977) World distribution of sutures— the site of former oceans. Tectonophys 40:69–99 Calvez JY, Alsac C, Delfour J, Kemp J, Pellaton C (1983) Geologic Evolution of Western, Central, and Eastern Parts of the Northern Precambrian Shield, Kingdom of Saudi Arabia. Saudi Arabian Deputy Ministry for Mineral Resources Open-File Report BRGM-OF-03-17, 57 p Calvez JY, Alsac C, Delfour J, Kemp J, Pellaton C (1984) Geological evolution of western, central, and eastern parts of the northern Precambrian Shield, Kingdom of Saudi Arabia. King Abdulaziz University, Jiddah, Faculty of Earth Science Bulletin 6:24–48 Camp VE (1984) Island arcs and their role in the evolution of the west Arabian Shield. Geol Soc Am Bull 95:913–921 Cawood PA (2005) Terra Australis Orogen: Rodinia breakup and development of the Pacific and Iapetus margins of Gondwana during the Neoproterozoic and Paleozoic. Earth Sci Rev 69:249–279 Cawood PA, Buchan C (2007) Linking accretionary orogenesis with supercontinent assembly. Earth Sci Rev 82:217–256 Cawood PA, Kröner A, Collins WJ, Kusky TM, Mooney WD, Windley BF (2009) Accretionary orogens through Earth history. In: Cawood PA, Kröner A (eds) Earth Accretionary Systems in Space and Time. Geol Soc Spec Publ 318:1–36

10

Terrane Accretion Within the Arabian-Nubian Shield

Church WR (1988) Ophiolites, sutures, and micro-plates of the Arabian-Nubian Shield: a critical comment. In: El Gaby S, Greiling RO (eds) The Pan-African belt of northeast Africa and adjacent areas. Friedr. Vieweg & Sohn, Weisbaden, pp 289–318 Church WR (1991) Discussion of “Ophiolites in northeast and east Africa: implication for Proterozoic crustal growth”. J Geol Soc Lond 148:600–606 Clift PD, Vannucchi P, Morgan JP (2009) Crustal redistribution, crust– mantle recycling and Phanerozoic evolution of the continental crust. Earth Sci Rev 97:80–104 Coleman RG (1984) Ophiolites and the tectonic evolution of the Arabian Peninsula. Geol Soc Lond Spec Publ 13:359–366 Collins AS, Pisarevsky SA (2005) Amalgamating eastern Gondwana; the evolution of the Circum-Indian Orogens. Earth Sci Rev 71:229– 270 Collins AS, Windley BF (2002) The Tectonic Evolution of Central and Northern Madagascar and Its Place in the Final Assembly of Gondwana. J Geol 110:325–339 Condie KC (2002) Breakup of a Paleoproterozoic Supercontinent. Gondw Res 5:41–43 Coney PJ (1978) Mesozoic-Cenozoic Cordilleran plate tectonics. In: Smith RB, Eaton GP (Eds.) Cenozoic tectonics and regional geophysics of the western Cordillera. Geol Soc Am Mem 152:33–50 Coney PJ, Jones DL, Monger JWH (1980) Cordilleran suspect terranes. Nature 288:328–333 Coolen JJMMM, Priem HNA, Verdurmen EAT, Verschure RH (1982) Possible zircon U-Pb evidence for Pan-African granulite-facies metamorphism in the Mozambique Belt of southern Tanzania. Precamb Res 17:31–40 Cordani UG, D’Agrella-Filho MS, Brito-Neves BB, Trindade RIF (2003) Tearing up Rodinia: the Neoproterozoic palaeogeography of South American cratonic fragments. Terra Nova 15:350–359 Cox GM, Lewis JC, Collins SA, Halverson PG, Jourdan F, Foden J, Nettle D, Kattan F (2012) Ediacaran terrane accretion within the Arabian-Nubian Shield. Gondw Res 21:341–352 Dalziel IWD (1997) Neoproterozoic-paleozoic geography and tectonics: review, hypothesis, environmental speculation. Geol Soc Am Bull 109:16–42 Davies FB (1984) Strain analysis of wrench faults and collision tectonics of the Arabian-Nubian Shield. J Geol 82:37–53 Delfour J (1979) Pan-African orogeny in the northern part of the Arabian Shield. Bull Soc Geol France S7-XXI(4):449–456 Delfour J (1980a) Geologic, tectonic and metallogenic evolution of the northern part of the Precambrian Arabian Shield (Kingdom of Saudi Arabia). Bull. BRGM, sect. II, 1–2:1–19 Delfour J (1980b) Geologic Map of the Ar Rika quadrangle, sheet 22G, Kingdom of Saudi Arabia. Saudi Arabian Deputy Ministry for Mineral Resources Geoscience Map GM 51A, scale 1:250,000, with text, p 34 Delfour J, Dhellemmes R, Elsass P, Vaslet D, Brosse JM, Le Nindre YM, Dottin O (1982) Geologic Map of the Ad Dawadimi quadrangle, sheet 24G, Kingdom of Saudi Arabia. Saudi Arabian Deputy Ministry for Mineral Resources Geoscience Map GM 60A, scale 1:250,000, with text, p. 36 de Wall H, Greiling RO, Sadek MF (2001) Post-collisional shortening in the late Pan-African Hamisana high strain zone, SE Egypt: field and magnetic fabric evidence. Precamb Res 107:179–194 de Wit MJ, Bowring SA, Ashwal LD, Randrianasolo LG, Morel VPI, Rambeloson RA (2001) Age and tectonic evolution of Neoproterozoic ductile shear zones in southwestern Madagascar, with implications for Gondwana studies. Tectonics 20:1–45 Dilek Y, Ahmed PT (2003) Ophiolite in Earth History. Geol Soc Lond Spec Publ 218:685–700 Dixon TH, Golombek MP (1988) Late Precambrian crustal accretion rates in northeast Africa and Arabia. Geology 16:991–994

261 Drury SA, De Souza Filho CR (1998) Neoproterozoic terrane assemblages in Eritrea: review and prospects. J Afr Earth Sci 27:331–348 Duyverman HJ (1984) Late Precambrian granitic and volcanic rocks and their relation to the cratonisation of the Arabian Shield. Faculty Earth Science Bulletin, King Abdulaziz University, Jeddah 6: 50–69 El-Bayoumi RM (1980) Ophiolites and associated rocks of Wadi Ghadir, east of Gebel Zabara, Eastern Desert, Egypt. PhD Thesis, Cairo University, Egypt El-Bayoumi RMA (1983) Ophiolite and mélange complex of Wadi Ghadir area, Eastern Desert, Egypt. Bull Fac Sci King Abdulaziz Univ 6:329–342 El-Gaby S, List F, Tehrani R (1988) Geology, evolution and metallogenesis of the Pan-African belt in Egypt. In: El-Gaby S, Greiling RO (eds) The Pan-African Belt of Northeast Africa and Adjacent Areas. Friedr. Vieweg & Sohn, Weisbaden, pp 17–68 El-Nisr SA (1997) Late precambrian volcanism at Wadi Allaqi, south Eastern Desert, Egypt: evidence for transitional continental arc/margin environment. J Afr Earth Sci 24:301–313 El Ramly MF, Greiling R, Kroner A, Rashwan AA (1984) On the tectonic evolution of the Wadi Hafafit area and environs, Eastern Desert of Egypt. Bull Fac Earth Sci 6:114–126 King Abdulaziz University, Jiddah El-Sharkawy MA, El-Bayoumi RM (1979) The ophiolites of Wadi Ghadir area, Eastern Desert, Egypt. Ann Geol Surv Egypt 9:125– 135 Embleton JCB, Hughes DJ, Klemenic PM, Poole S, Vail JR (1982) A new approach to the stratigraphy and tectonic evolution of the Red Sea Hills, Sudan. Precamb Res 16:A19 Evans DA (1998) True polar wander: a supercontinental legacy. Earth Planet Sci Lett 157:1–8 Evans DAD (2000) Stratigraphic, geochronological, and paleomagnetic constraints upon the Neoproterozoic climate paradox. Am J Sci 300:374–433 Feistmantel O (1876) Notes on the age of some fossil floras of India. Rec Geol Surv India 12:163–166 Fitches WR, Graham RH, Hussein IM, Ries AC, Shackleton RM, Rice RC (1983) The Late Proterozoic ophiolite of Sol Hamed, NE Sudan. Precamb Res 19:385–441 Fleck RJ, Greenwood WR, Hadley DG, Anderson RE, Schmidt DL (1980) Rubidium-Strontium geochronology and plate-tectonic evolution of the southern part of the Arabian Shield. U.S. Geol Surv Prof Pap No. 1131, 38 pp Frisch W, Al-Shanti A (1977) Ophiolite belts and the collision of island arcs in the Arabian Shield. Tectonophys 43:293–306 Frisch W, Pohl W (1986) Petrochemistry of some mafic and ultramafic rocks from the Mozambique Belt, SE-Kenya. Mitteilungen der Österreichischen Geologischen Gesellschaft 78:97–114 Fritz H, Abdelsalam M, Ali KA, Bingen B, Collins AS, Fowler AR, Ghebreab W, Hauzenberger CA, Johnson PR, Kusky TM, Macey P, Muhongo S, Stern RJ, Viola G (2013) Orogen styles in the East African Orogen: a review of the Neoproterozoic to Cambrian tectonic evolution. J Afr Earth Sci 86:65–106 Garson MS, Shalaby IM (1975) Precambrian-Lower Palaeozoic plate tectonics and metagabbros in the Red Sea region. Spec Pap Geol Assoc Can 14:573–596 Garson MS, Shalaby IM (1976) Precambrian—Lower Paleozoic plate tectonics and metallogenesis in the Red Sea region. Geol Assoc Canada Spec Pap 14:573–596 Gass IG (1981) Pan-African (Upper Proterozoic) plate tectonics of the Arabian-Nubian Shield. In: Kröner A (ed) Precambrian Plate Tectonics. Elsevier, Amsterdam, pp 388–405 Genna A, Nehlig P, Le Goff E, Guerrot C, Shanti M (2002) Proterozoic tectonism of the Arabian Shield. Precamb Res 117:21–40

262 Gradstein FM, Ogg JG, Smith AG (2004) A Geologic Time Scale 2004. Cambridge University Press, Cambridge Gray DR, Foster DA, Meert JG, Goscombe BD, Armstrong R, Trouw RAJ, Passchier CW (2008) A Damara orogen perspective on the assembly of southwestern Gondwana. Geol Soc Lond Spec Publ 294:257–278 Greenwood WR, Hadley DG, Anderson RE, Fleck RJ, Schmidt DL (1976) Late Proterozic cratonization in southwestern Saudi Arabia. Phil Trans Roy Soc Lond 280:517–527 Greenwood WR, Hadley DG, Schmidt DL (1973) Tectono-stratigraphic subdivision of Precambrian rocks in the southern part of the Arabian Shield. Geol Soc Am Abstracts Progr 5:643 Greenwood WR, Anderson RE, Fleck RJ, Roberts RJ (1980) Precambrian geologic history and plate tectonic evolution of the Arabian shield: Saudi Arab. Director. Gen Miner Res Bull 24:35 pp Greenwood WR, Stoeser DB, Fleck RJ, Stacey JS (1982) Late Proterozoic island-arc complexes and tectonic belts in the southern part of the Arabian Shield, Kingdom of Saudi Arabia. Saudi Arabian Ministry Petrol. Mineral Resour Open File Rep. USGS-OF-02-8 Greiling RO, Abdeen MM, Dardir AA, El Akhal H, El Ramly MF, Kamal El Din GM, Osman AF, Rashwan AA, Rice AH, Sadek MF (1994) A structural synthesis of the Proteorozoic Arabian-Nubian Shield in Egypt. Geol Rundsch 83:484–501 Grunow AM (1999) Gondwana events and palaeogeography: a palaeomagnetic review. J Afr Earth Sci 28:53–69 Hadley DG, Schmidt DL (1980) Sedimentary rocks and basins of the Arabian Shield and their evolution. In: Cooray PG, Tahoun SA (eds), Evolution and Mineralization of the Arabian–Nubian Shield, Inst Appl Geol, Jiddah, Bull 3, vol 4, Pergamon Press, pp 25–50 Haider AH, Berhe S (1995) Crustal evolution in the Mozambique belt and implications for reconstruction of Gondwana. Centennial Geocongress Abstracts Geological Society South Africa 370:75p Hargrove US, Stern RJ, Kimura J-I, Manton WI, Johnson PR (2006) How juvenile is the Arabian-Nubian Shield? Evidence from Nd isotopes and pre-Neoproterozoic inherited zircon in the Bi’r Umq suture zone, Saudi Arabia. Earth Planet Sci Lett 252:308–326 Harms U, Schandelmeier H, Darbyshire DPF (1990) Pan-African reworked early/middle Proterozoic crust in NE Africa west of the Nile: Sr and Nd isotope evidence. J Geol Soc Lond 147:859–872 Harris NBW, Hawkesworth CJ, Ries AC (1984) Crustal evolution in northeast and East Africa from model Nd ages. Nature 309:773–776 Harris NBW, Gass IG, Hawkesworth FRS, Hawkesworth CJ (1990) A geochemical approach to allochthonous terranes: a Pan-African case study. Phil Trans R Soc London A 331:533–548 Hauzenberger CA, Sommer H, Fritz H, Bauernhofer A, Kröner A, Hoinkes G, Wallbrecher E, Thöni M (2007) SHRIMP U-Pb zircon and Sm–Nd garnet ages from the granulite-facies basement of SE Kenya: evidence for Neoproterozoic polycyclic assembly of the Mozambique Belt. J Geol Soc Lond 164:189–201 Hoffman PF, Kaufman AJ, Halverson GP, Schrag DP (1998) A Neoproterozoic snowball Earth. Science 281:1342–1346 Howell DG (ed) (1985) Tectonostratigraphic Terranes of the Circum-Pacific Region. Houston: Circum-Pacific Council for Energy and Mineral Resources, Earth Sciences Series Number 1, 585 pp Howell DG (1989) Tectonics of Suspect Terranes. Chapman & Hall, London and New York, p 232p Howell DG (1993) Tectonics of suspect terranes—mountain building and continental growth. Topics in Earth Sciences 3. Chapman & Hall. London, 231p Howell DG (1995) Principles of Terrane Analysis, 2nd edn. Chapman & Hall, London Howell DG, Jones DL, Schermer ER (1985) Tectonostratigraphic terranes of the circum-Pacific region. In: Howell DG (ed) Tectonostratigraphic terranes of the circum-Pacific region.

A. F. Osman and A.-R. Fowler Circum-Pacific Council for Energy and Mineral Resources, Houston, pp 3–30 Hussein IM (1977) Geology of the Halaib area of the northern Red Sea Hills, Sudan, with special reference to the Sol Hamed basic complex. M.Phil. thesis, Portsmouth Polytechnic Hussein IM, Kröner A, Durr S (1982) Wadi Onib: a dismemberedPan-African ophiolite in the Red Sea Hills of the Sudan. Precamb Res 16:A52 Hussein IM, Kröner A, St Durr (1984) Wadi Oniba dismembered Pan-African ophiolite in the Red Sea Hills of Sudan. Faculty Earth Science Bulletin, King Abdulaziz University, Jeddah 6:320–327 Irwin WP (1972) Terranes of the western Paleozoic and Triassic belt in the southern Klamath Mountains, California: U.S. Geol. Surv. Prof. Pap. 800-C:C103–C111 Jackson NJ, Ramsay CR (1980) Time-space relationships of Upper Precambrian volcanic and sedimentary units in the Central Arabian Shield. J Geol Soc Lond 137:617–628 Johansson A (2014) From Rodinia to Gondwana with the ‘SAMBA’ model—A distant view from Baltica towards Amazonia and beyond. Precamb Res 244:226–235 Johnson PR (1996) Geochronologic and isotopic data for rocks in the east-central part of the Arabian Shield: stratigraphic and tectonic implications. Saudi Arabian Deputy Ministry for Mineral Resources, Open-file report USGSOF-96-3 Johnson PR (2001) Implications of SHRIMP and microstructural data on the age and kinematics of shearing in the Asir terrane, southern Arabian shield, Saudi Arabia. Gondw Res 4:172–173 Johnson PR (2003) Post-amalgamation basins of the NE Arabian Shield and implications for Neoproterozoic tectonism in the northern East African Orogen. Precam Res 123:321–337 Johnson PR (ed) (2012) The Geology of the Saudi Arabian Shield. Saudi Geological Survey, Jeddah, p 479 Johnson PR (2014) An Expanding Arabian-Nubian Shield Geochronologic and Isotopic Dataset: Defining Limits and Confirming the Tectonic Setting of a Neoproterozoic Accretionary Orogen. The Open Geol J 8:3–33 Johnson PR, Kattan F (2001) Oblique sinistral transpression in the Arabian shield: the timing and kinematics of a Neoproterozoic suture zone. Precamb Res 107:117–138 Johnson PR, Vranas GJ (1984) The origin and development of Late Proterozoic rocks of the Arabian Shield. Saudi Arabian Deputy Ministry for Mineral Resources Open-File Report RF-OF-04-32 Johnson PR, Woldehaimanot B (2003) Development of the Arabian– Nubian shield: perspectives on accretion and deformation in the northern East African Orogen and the assembly of Gondwana. In: Yoshida M, Windley BF, Dasgupta S (eds) Proterozoic East Gondwana: Supercontinent Assembly and Breakup. Geol Soc Lond Spec Publ 206:290–325 Johnson PR, Scheibner E, Smith EA (1987) Basement fragments, accreted tectonostratigraphic terranes, and overlap sequences: elements in the tectonic evolution of the Arabian Shield. In: Leitch EC, Scheibner E (eds) Terrane Accretion and Orogenic Belts. American Geophysical Union, Washington, pp 323–343 Johnson PR, Abdelsalam MG, Stern RJ (2003) The Bi’r Umq-Nakasib Suture Zone in the Arabian-Nubian Shield: A Key to Understanding Crustal Growth in the East African Orogen. Gond Res 6:523–530 Johnson PR, Kattan FH, Al-Saleh AM (2004) Neoproterozoic ophiolites in the Arabian Shield: Field relations and structure. In: Kusky TM (ed) Precambrian Ophiolites and Related Rocks, Elsevier, Developments in Precambrian Geology, 13:129–162 Johnson PR, Andresen A, Collins AS, Fowler AR, Fritz H, Ghebreab W, Kusky T, Stern RJ (2011) Late Cryogenian-Ediacaran history of the Arabian-Nubian Shield: a review of depositional, plutonic, structural, and tectonic events in the closing stages of the northern East African Orogen. J Afr Earth Sci 61:167–232

10

Terrane Accretion Within the Arabian-Nubian Shield

Johnson PR, Halverson GP, Kusky TM, Stern RJ, Pease V (2013) Volcanosedimentary Basins in the Arabian-Nubian Shield: Markers of Repeated Exhumation and Denudation in a Neoproterozoic Accretionary Orogen. Geoscience 3:389–445 Jones DL, Silberling NJ, Gilbert W, Coney P (1982) Character, distribution, and tectonic significance of accretionary terranes in the Central Alaska Range. J Geophys Res 875:3709–3717 Jones DL, Howell DG, Coney PJ, Mayer JWH (1983) Recognition, character, and analysis of tectonostratigraphic terranes in western North America. In: Hashimoto M, Uyeda S (eds) Accretion Tectonics in the Circum-Pacific regions. Terra Scientific Publishing Co., Tokyo, pp 21–35 Kazmin V (1976) Ophiolites in the Ethiopian basement. Ethiopian Institute of Geological Surveys, Addis Ababa, Note 35, 17 pp Kazmin V, Shiferaw A, Tefera M, Berhe SM, Chewaka S (1979) Precambrian structure of Western Ethiopia. Ann Geol Surv Egypt 9:1–18 Kennedy WO (1964) The structural differentiation of Africa in the Pan-African tectonic episode. 8th Ann Rep, Res Inst Afr Geol, Leeds Univ pp 48–49 Key RM, Charsley TJ, Hackman TJ, Wilkinson AF, Rundle CC (1989) Superimposed upper Proterozoic collision controlled orogenies in the Mozambique Orogenic belt of Kenya. Precamb Res 44:197–225 Kirschvink JL, Ripperdan RL, Evans DA (1997) Evidence for large-scale reorganization of Early Cambrian continental masses by inertial interchange true polar wander. Science 277:541–545 Kröner A (1977) The Precambrian geotectonic evolution of Africa: plate accretion vs. plate destruction. Precamb Res 4:163–213 Kröner A (1984) Late Precambrian plate tectonics and orogeny: a need to redefine the term Pan-African. In: Klerkx J, Michot J (eds) African Geology. Tervuren, Belgium, pp 23–28 Kröner A (1985) Ophiolites and the evolution of tectonic boundaries in the late Proterozoic Arabian-Nubian Shield of northeastern Africa and Arabia. Precamb Res 27:277–300 Kröner A, Sassi FP (1996) Evolution of the northern Somali basement: new constraints zircon ages. J Afr Earth Sci 22:1–15 Kröner A, Stern RJ (2004) Pan-African Orogeny. In: Selley RC, Cocks LRM, Plimer IR (eds) Encyclopedia of Geology, 1. Elsevier, Amsterdam, pp 1–12 Kröner A, Greiling R, Reischmann T, Hussein IM, Stern RJ, Durr S, Kruger J, Zimmer M (1987a) Pan-African crustal evolution in the Nubian segment of the northeast Africa. In: Kröner A (ed) Proterozoic lithospheric evolution. American Geophysical Union, Geodynamics Series, Washington, 17, pp 235–257 Kröner A, Stern RJ, Dawoud AS, Compston W, Reischmann T (1987b) The Pan-African continental margin in northeast Africa: evidence from a geochronological study of granulites at Sabaloka, Sudan. Earth Planet Sci Lett 85:91–104 Kröner A, Stern RJ, Linnebacker P, Manton W, Reischmann T, Hussein IM (1991) Evolution of Pan-African island arc assemblages in the south Red Sea Hills, Sudan, and in SW Arabia as exemplified by geochemistry and geochronology. Precamb Res 63:99–118 Kröner A, Todt W, Hussein IM, Mansour M, Rashwan A (1992) Dating of Late Proterozoic ophiolites in Egypt and Sudan using the single grain zircon evaporation technique. Precamb Res 59:15–32 Kröner A, Hegner E, Collins AS, Windley BF, Brewer TS, Razakamanana T, Pidgeon RT (2000) Age and magmatic history of the Antananarivo Block, central Madagascar, as derived from zircon geochronology and Nd isotopic systematics. Am J Sci 300:251–288 Kröner A, Hegner E, Lehmann B, Heinhorst J, Wingate MTD, Liu DY, Ermelov P (2008) Palaeozoic arc magmatism in the Central Asian Orogenic Belt of Kazakhstan: SHRIMP zircon ages and whole-rock Nd isotopic systematics. J Asian Earth Sci 32:118–130 Kusky T, Matsah M (2003) Neoproterozoic Dextral Faulting on the Najd Fault System, Saudi Arabia, preceded sinistral faulting and

263 escape tectonics related to closure of the Mozambique Ocean. In: Yoshida M, Windley BF, Dasgupta S, Powell C (eds) Proterozoic East Gondwana: Supercontinent Assembly and Breakup. Geol Soc, Lond Spec Publ 206:327–361 Kusky TM, Ramadan TM (2002) Structural controls on Neoproterozoic mineralization in the South Eastern Desert, Egypt: an integrated field, Landsat TM, and SIR-C/X SAR approach. J Afr Earth Sci 35:107–121 Kusky TM, Abdelsalam M, Tucker R, Stern RJ (2003) Evolution of the East African and related Orogens, and the assembly of Gondwana. Special issue Precamb Res 123:81–85 Laval M, Le Bel L (1986) Felsic plutonism in the Al Amar-Idsas area, Kingdom of Saudi Arabia. J Afr Earth Sci 4:87–98 Lenoir JL, Kuster D, Liegeois JP, Utke A, Haider A, Matheis G (1994) Origin and regional significance of late Precambrian and early Palaeozoic granitoids in the Pan-African belt of Somalia. Geol Rundsch 83:624–641 Li ZX, Bogdanova SV, Collins AS, Davidson A, De Waele B, Ernst RE, Fitzsimons ICW, Fuck RA, Gladkochub DP, Jacobs J, Karlstrom KE, Lu S, Natapov LM, Pease V, Pisarevsky SA, Thrane K, Vernikovsky V (2008) Assembly, configuration, and break-up history of Rodinia: a synthesis. Precam Res 160:179– 210 Martelat J-E, Lardeaux JM, Nicollet C, Rakotondrazafy R (2000) Strain pattern and late Precambrian deformation history in southern Madagascar. Precamb Res 102:1–20 Medlicott HB, Blandford WT (1879) A manual of geology of India, Calcutta Meert JG (2002) Paleomagnetic evidence for a Paleo-Mesoproterozoic supercontinent, Columbia. Gondw. Res. 5:207–215 Meert JG (2003) A synopsis of events related to the assembly of eastern Gondwana. Tectonophys 362:1–40 Meert JG (2012) What’s in a name? The Columbia (Paleopangea/Nuna) supercontinent. Gondw Res 21:987–993 Meert JG, Lieberman BS (2008) The Neoproterozoic assembly of Gondwana and its relationship to the Ediacaran-Cambrian Radiation. Gondw Res 14:5–21 Meert JG, Torsvik TH (2003) The making and unmaking of a super-continent: Rodinia revisited. Tectonophys 375:261–288 Meert JG, Van der Voo R (1997) The assembly of Gondwana 800–550 Ma. J Geodyn 23:223–236 Meert JG, Van der Voo R, Ayub S (1995) Paleomagnetic investigation of the Neoproterozoic Gagwe lavas and Mbozi Complex, Tanzania and the assembly of Gondwana. Precamb Res 74:225–244 Miller MM, Dixon TH (1992) Late Proterozoic evolution of the north part of the Hamisana zone, northeast Sudan: constraints on Pan-African accretionary tectonics. J Geol Soc Lond 149:743–750 Monger JWH, Souther JG, Gabrielse H (1972) Evolution of the Canadian Cordillera: A plate-tectonic model. Am J Sci 272:577– 602 Möller A, Mezger K, Schenk V (1998) Crustal age domains and the evolution of the continental crust in the Mozambique Belt of Tanzania: combined Sm–Nd, Rb–Sr and Pb–Pb isotopic evidence. J Petrol 39:749–783 Möller A, Mezger K, Schenk V (2000) U-Pb dating of metamorphic minerals: Pan-African metamorphism and prolonged slow cooling of high pressure granulites in Tanzania, East Africa. Precamb Res 104:123–146 Moore JM (1979) Tectonics of the Najd transcurrent fault system, Saudi Arabia. J Geol Soc Lond 136:441–453 Moores EM (1982) Origin and significance of ophiolites. Rev Geophys Space Phys 20:735–760 Moores EM, Twiss RJ (1995) Tectonics. Freeman, New York, p 415 Mosely PN (1993) Geological evolution of the late Proterozoic “Mozambique Belt” of Kenya. Tectonophys 221:223–250

264 Nassief MO, MacDonald R, Gass IG (1984) The Jebel Thurwah Upper Proterozoic ophiolite complex, west Saudi Arabia. J Geol Soc Lond 141:537–546 Nasseef AO, Bakor AR, Hashad AH (1980) Petrography of possible ophiolitic rocks along the Qift-Qusier Road, Eastern Desert, Egypt. Bulletin King Abdelaziz Univ 3:157–168 Nawab ZA (1979) Geology of the AI Amar-ldsas region of the Arabian Shield. Institute Applied Geology Bulletin, King Abdulaziz University, Jeddah 3:29–39 Nicolas A (1989) Structures of ophiolites and dynamics of oceanic lithosphere. Kluwer Academic Publishers, Dordrecht, The Netherlands, p 367 Pallister IS, Stacey JS, Fischer LB, Premo WR (1987) Arabian Shield ophiolites and Late Proterozoic microplate accretion. Geology 15:320–323 Pallister JS, Stacey JS, Fischer LB, Premo WR (1988) Precambrian ophiolites of Arabia: geologic settings, U-Pb geochronology, Pb-isotope characteristics, and implications for continental accretion. Precamb Res 38:1–54 Pisarevsky SA, Wingate MTD, Powell CMcA, Johnson S, Evans DAD (2003) Models of Rodinia assembly and fragmentation. In: Yoshida M, Windley BE, Dasgupta S (eds), Proterozoic East Gondwana: Supercontinent Assembly and Breakup. Geol Soc Lond Spec Publ 206:35–55 Powell CM, Young G (1995) Are Neoproterozoic glacial deposits preserved on the margins of Laurentia related to the fragmentation of two supercontinents?: comment and reply. Geology 23:1053– 1054 Quick JE (1991) Late Proterozoic transpression on the Nabitah fault system—implications for the assembly of the Arabian Shield. Precamb Res 53:119–147 Quick JE, Bosch PS (1989) Tectonic history of the north Nabitah fault zone, Arabian Shield, Kingdom of Saudi Arabia. Ministry of Petroleum and Mineral Resources, Jeddah, Technical Report. USGS TR-08-2, 87p Reischmann T (2000) Ophiolites and island arcs in the late Proterozoic Nubian Shield. Ofioliti 25:1–13 Ries AC, Shackleton RM, Graham RH, Fitches WR (1983) Pan-African structures, ophiolites and mélanges in the east Desert of Egypt: a traverse at 26° north. J Geol Soc Lond 140:75–95 Ries AC, Vearncombe JR, Price RC, Shackleton RM (1992) Geochronology and geochemistry of the rocks associated with a late Proterozoic ophiolite in West Pokot, northwest Kenya. J Afr Earth Sci 14:25–36 Robinson FA, Foden JD, Collins AS, Payne JL (2014) Arabian Shield magmatic cycles and their relationship with Gondwana assembly: Insights from zircon U-Pb and Hf isotopes. Earth Planet Sci Lett 408:207–225 Robinson FA, Foden JD, Collins A (2015a) Geochemical and isotopic constraints on island arc, synorogenic, post-orogenic and anorogenic granitoids in the Arabian Shield, Saudi Arabia. Lithos 220– 223:97–115 Robinson FA, Foden JD, Collins AS (2015b) Zircon geochemical and geochronological constraints on contaminated and enriched mantle sources beneath the Arabian Shield, Saudi Arabia. J Geol 123:463– 489 Rocci G (1965) Essai d’interprétation des mésures géochronologiques. La structure de l’ouest africain. Sci Terre 10:461–479 Rogers JJW, Santosh M (2004) Continents and Supercontinents. Oxford University Press, 289 pp Schandelmeier H, Darbyshire DPF, Harms U, Richter A (1988) The east Saharan Craton: evidence for pre-Pan-African crust in NE Africa west of the Nile. In: El-Gaby S, Greiling RO (eds) The Pan-African belts of Northeast Africa and Adjacent areas. Friedr. Vieweg & Sohn, Weisbaden, pp 69–94

A. F. Osman and A.-R. Fowler Searle MP, Stevens RK (1984) Obduction processes in ancient, modern and future ophiolites. In: Gass IG et al (eds) Ophiolites and Oceanic Lithosphere. Geol, Soc Lond, pp 303–319 Schmidt DL, Hadley DG, Stoeser DB (1979) Late Proterozoic crustal history of the Arabian shield, southern Najd province, Kingdom of Saudi Arabia. King Abdulaziz University, Jiddah. Ins Appl Geol Bull 3(2):41–58 Scotese CR (2009) Late Proterozoic plate tectonics and palaeogeography: a tale of two supercontinents, Rodinia and Pannotia. In: Craig J, Thurow J, Thusu B, Whitham A, Abutarruma Y (eds) Global Neoproterozoic Petroleum Systems: The Emerging Potential in North Africa. Geol Soc Lond Spec Public 326:67–83 Shackleton RM (1979) Precambrian tectonics of northeast Africa. Inst. Appl. Geol. Bull. King Abdulaziz Univ. Jeddah 3:1–6 Shackleton RM (1986) Precambrian collision tectonics in Africa. In: Coward MP, Ries AC (eds) Collision tectonics. Geol Soc Lond Spec Publ 19:329–349 Shackleton RM (1994) Review of late Proterozoic sutures, ophiolitic melanges and tectonics of eastern Egypt and northeast Sudan. Geol Rundsch 83:537–546 Shackleton RM, Ries AC, Graham RH, Fitches WR (1980) Late Precambrian ophiolitic melange in the east Desert of Egypt. Nature 285:472–474 Shanti M (1982) The Jabel Ess ophiolite complex. Precamb Res 16: A55–A56 Shanti M, Roobol MJ (1979) A late Proterozoic ophiolite complex at Jebel Ess in north Saudi Arabia. Nature 279:488–491 Shimron AE (1980) Proterozoic island arc volcanism and sedimentation in Sinai. Precam Res 12:437–458 Silver EA, Reed DL (1988) Backthrusting in accretionary wedges. J Geophys Res Solid Earth 93(B4):3116–3126 Silver EA, Reed D, McCaffrey R, Joyodiwiryo Y (1983) Back arc thrusting in the Eastern Sunda Arc, Indonesia: a consequence of arc-continent collision. J Geophys Res 88(B9):7429–7448 Snoke AW, Barnes CG (2006) The development of tectonic concepts for the Klamath Mountains province, California and Oregon. Geol Soc Am Spec Pap 410:1–29 Squire RJ, Campbell IH, Allen CM, Wilson CJL (2006) Did the Transgondwanan Supermountain trigger the explosive radiation of animals on Earth? Earth Planet Sci Lett 250:116–133 Stacey JS, Agar RA (1985) U-Pb isotopic evidence for the accretion of a continental micro plate in the Zalm region of the Saudi Arabian Shield. J Geol Soc Lond 142:1189–1203 Stacey JS, Hedge CE (1984) Geochronologic and isotopic evidence for early Proterozoic crust in the eastern Arabian shield. Geology 12:310–313 Stacey JS, Stoeser D (1983) Distribution of oceanic and continental leads in the Arabian-Nubian Shield. Contrib Mineral Petrol 84:91– 105 Stacey JS, Doe BR, Roberts RJ, Delevaux MH, Gramlich JW (1980) A lead isotope study of mineralization in the Saudi Arabian Shield. Contrib Min Petrol 74:175–188 Stacey JS, Stoeser DB, Greenwood WR, Fischer LB (1984) U-Pb zircon geochronology and geologic evolution of the Halaban–Al Amar region of the eastern Arabian Shield, Kingdom of Saudi Arabia. J Geol Soc Lond 141:1043–1055 Stein M (2003) Tracing the plume material in the Arabian-Nubian Shield. Precam. Res. 123:223–234 Stein M, Goldstein SL (1996) From plume-head to continental lithosphere in the Arabian-Nubian Shield. Nature 382:773–778 Stern RJ (1985) The Najd fault system, Saudi Arabia and Egypt: a Late Precambrian rift related transform system. Tectonics 4:497–511 Stern RJ (1993) Tectonic evolution of the Late Proterozoic East African Orogen: constraints from crustal evolution in the Arabian-Nubian Shield and the Mozambique Belt. In: Thorweihe U, Scandelmeier H

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Terrane Accretion Within the Arabian-Nubian Shield

(eds) Geoscientific Research in Northeast Africa. Balkema, Rotterdam, pp 73–74 Stern RJ (1994) Arc assembly and continental collision in the Neoproterozoic East African Orogen implications for the consolidation of Gondwanaland. Ann Rev Earth Planet Sci 22:319–351 Stern RJ (2002) Crustal evolution in the East African Orogen: a neodymium isotopic perspective. J Afr Earth Sci 34:109–117 Stern RJ, Abdelsalam MG (1998) Formation of juvenile continental crust in the Arabian-Nubian shield: evidence from granitic rocks of the Nakasib Suture, NE Sudan. Geol Rundsch 87:150–160 Stern RJ, Hedge CE (1985) Geochronologic and isotopic constraints on late Precambrian crustal evolution in the Eastern Desert of Egypt. Am J Sci 285:97–127 Stern RJ, Johnson PR (2010) Continental lithosphere of the Arabian Plate: a geologic, petrologic, and geophysical synthesis. Earth-Sci Rev 101:29–67 Stern RJ, Kröner A (1993) Geochronologic and isotopic constraints on the late Precambrian crustal evolution m northeast Sudan. J Geol 101:555–574 Stern RJ, Scholl DW (2010) Yin and yang of continental crust creation and destruction by plate tectonic processes. Int Geol Rev 52:1–31 Stern RJ, Gottfried D, Hedge CE (1984) Late Precambrian rifting and crustal evolution in the Northeastern Desert of Egypt. Geology 12:168–172 Stern RJ, Kröner A, Manton WI, Reischmann T, Mansour M, Hussein IM (1989) Geochronology of the late Precambrian Hamisana shear zone, Red Sea Hills, Sudan and Egypt. J Geol Soc Lond 146:1017–1030 Stern RJ, Nielsen KC, Best E, Sultan M, Arvidson RE, Kröner A (1990) Orientation of late Precambrian sutures in the Arabian-Nubian Shield. Geology 18:1103–1106 Stern RJ, Kröner A, Reischmann T, Bender R, Dawoud AS (1994) Precambrian basement around Wadi Halfa: a new perspective on the evolution of the Central Saharan Ghost craton. Geol Rundsch 83:564–577 Stern RJ, Ali KA, Liegeois JP, Johnson PR, Kozdroj W, Kattan FH (2010) Distribution and significance of Pre-Neoproterozoic zircons in juvenile Neoproterozoic igneous rocks of the Arabian-Nubian Shield. Am J Sci 310:791–811 Stoeser DB (1986) Distribution and tectonic setting of plutonic rocks of the Arabian Shield. J Afr Earth Sci 4:21–46 Stoeser DB, Camp VE (1985) Pan-African microplate accretion of the Arabian Shield. Geol Soc Am Bull 96:817–826 Stoeser DW, Frost CD (2006) Nd, Pb, Sr and O isotope characterization of Saudi Arabian Shield terranes. Chem Geol 226:163–188 Stoeser DB, Stacey JS (1988) Evolution, U-Pb geochronology, and isotope geology of the Pan-African Nabitah orogenic belt of the Saudi Arabian Shield. In: El-Gaby S, Greiling RO (eds) The Pan-African Belt of NE Africa and Adjacent Areas: Tectonic Evolution and Economic Aspects of a Late Proterozoic Orogen. Friedr. Vieweg & Sohn, Weisbaden, pp 227–228 Stoeser DB, Fleck RJ, Stacey JS (1984) Geochronology and origin of an early tonalite gneiss of the Wadi Tarib batholith and the formation of syntectonic gneiss complexes in the southeast Arabian Shield. Faculty Earth Science Bulletin, King Abdulaziz University, Jeddah 6:351–364 Stoeser DB, Whitehouse MJ, Agar RA, Stacey JS (1991) Pan-African accretion and continental terranes of the Arabian Shield, Saudi Arabia and Yemen. EOS Abstracts 72:299 Sultan M, Arvidson RE, Sturchio NC (1986) Mapping of serpentinites in the Eastern Desert of Egypt by using Landsat thematic mapper data. Geology 14:995–999 Sultan M, Chamberlin KR, Bowring SA, Arvidson RE, Abuzied H, El Kaliouby B (1990) Geochronologic and isotopic evidence for

265 involvement of pre-Pan-African crust in the Nubian Shield, Egypt. Geology 18:761–764 Sultan M, Bickford ME, El-Kaliouby B, Arvidson RE (1992) Common Pb systematics of Precambrian granitic rocks of the Nubian Shield (Egypt) and tectonic implications. Geol Soc Am Bull 104:456–470 Sultan M, Tucker RD, Gharbawi RI, Ragab AI, El Alfy Z (1993) On the location of the boundary between the Nubian Shield and the old African continent: inferences from U-Pb (zircon) and common Pb data. In: Thorweihe U, Schandelmeier H (eds) Geoscientific Research in NE Africa. Balkema, Rotterdam, pp 75–77 Teixeira JBG, Misi A, da Silva M (2007) Supercontinent evolution and the Proterozoic metallogeny of South America. Gondw Res 11:346– 361 Teklay M, Kröner A, Oberhansli R (1993) Reconnaissance Pb-Pb zircon ages from Precambrian rocks in eastern and southern Ethiopia and an attempt to define crustal provinces. In: Thorweihe U, Schandelmeier H (eds) Geoscientific Research in northeast Africa. Balkema, Rotterdam, pp 133–138 Tenczer V, Hauzenberger CA, Fritz H, Whitehouse MJ, Mogessie A, Wallbrecher E, Muhongo S, Hoinkes G (2006) Anorthosites in the Eastern Granulites of Tanzania—new SIMS zircon U-Pb age data, petrography and geochemistry. Precamb Res 148:85–114 Tetreault JL, Buiter SJH (2012) Geodynamic models of terrane accretion: Testing the fate of island arcs, oceanic plateaus, and continental fragments in subduction zones. J Geophys Res 117: B08403. https://doi.org/10.1029/2012JB009316 Torsvik TH, Cocks LRM (2013) New global palaeogeographical reconstructions for the Early Palaeozoic and their generation. Geol Soc 38(1):5–24. London, Memoirs Torsvik TH, Carter LM, Ashwal LD, Bhushan SK, Pandit MK, Jamteit B (2001a) Rodinia refined or obscured: palaeomagnetism of the Malani igneous suite NW India. Precamb Res 108:319–333 Torsvik TH, Ashwal LD, Tucker RD, Eide EA (2001b) Geochronology and palaeomagnetism of the Seychelles microcontinent: the India link. Precamb Res 110:47–59 Trompette R (1997) Neoproterozoic (*600 Ma) aggregation of Western Gondwana: a tectonic scenario. Precamb Res 82:101–112 Vail JR (1983) Pan-African crustal accretion in northeast Africa. J Afr Earth Sci 1:285–294 Vail JR (1985) Pan-African (late Precambrian) tectonic terranes and reconstruction of the Arabian-Nubian Shield. Geology 13:839–842 Vail JR (1988) Tectonics and evolution of the Proterozoic basement of northeast Africa. In: El-Gaby S, Greiling RO (eds) The Pan-African belts of northeast Africa and Adjacent areas. Friedr. Vieweg & Sohn, Weisbaden, pp 185–226 Vail JR, Almond DC, Hughes DJ, Klemenic PM, Poole S, Nour SEM, Embleton JCB (1984) Geology of the Wadi Oko Khor Hayet area, Red Sea Hills, Sudan. Bulletin of the Geology and Mineral Resources Department of Sudan 34:1–20 Van Hinsbergen DJJ, Buiter SJH, Torsvik TH, Gaina C, Webb SJ (2011) The formation and evolution of Africa from the Archaean to Present: introduction. In: Van Hinsbergen DJJ, Buiter SJH, Torsvik TH, Gaina C, Webb SJ (eds), The Formation and Evolution of Africa: A Synopsis of 3.8 Ga of Earth History. Geol Soc Lond Spec Publ 357:1–8 Vearncombe JR (1983) A Proposed Continental Margin in the Precambrian of Western Kenya. Geol Rundsch 72:663–670 Veevers JJ (2004) Gondwanaland from 650–500 Ma assembly through 320 Ma merger in Pangea to 185–100 breakup: supercontinental tectonics via stratigraphy and radiometric dating. Earth-Sci Rev 68:1–132 Warden AJ, Horkel AD (1984) The Geological Evolution of the NE-Branch of the Mozambique Belt (Kenya, Somalia, Ethiopia). Mitt Osterr Geol Ges 77:161–184

266 Warden AJ, Kasmin V, Kiesel W, Pohl W (1982) Some Geological Data of the Mafic-Ultramafic Complex at Tulu Dimitri, Ethiopia and their Genetic Significance. - Sitzungsberichte. Osterreichische Akademie der Wissenschaften Mathematisch Naturwissenschaftliche Abt. 1, 191, (1–4), Vienna Whitehouse MJ, Stoeser DB, Stacey JS (2001a) The Khida terrane geochronological and isotopic evidence for Paleoproterozoic and Archean crust in the eastern Arabian Shield of Saudi Arabia. Gondw Res 4:200–202 Whitehouse MJ, Windley BF, Stoeser DB, Al-Khirbash S, Ba-Bttat MAO, Haider A (2001b) Precambrian basement character of Yemen and correlations with Saudi Araba and Somalia. Precamb Res 105:357–369 Windley BF, Whitehouse MJ, Ba-Bttat MAO (1996) Early Precambrian gneiss terranes and Pan-African island arcs in Yemen: Crustal accretion of the eastern Arabian Shield. Geology 24:131–134 Woldemichael BW, Kimura J-I, Dunkley DJ, Tani K, Ohira H (2010) SHRIMP U-Pb zircon geochronology and Sr–Nd isotopic systematic of the Neoproterozoic Ghimbi-Nedjo mafic to intermediate intrusions of Western Ethiopia: a record of passive margin magmatism at 855 Ma? Int J Earth Sci 99:1773–1790 Wust HJ, Todt W, Kröner A (1987) Conventional and single grain zircon ages for meta-sediments and granitic clasts from the east

A. F. Osman and A.-R. Fowler Desert of Egypt: evidence for active continental margin evolution in Pan-African times. Terra Cognita Abstracts 7:333–334 Yibas B, Reimold WU, Armstrong R, Koeberl C, Anhaeusser CR, Phillips D (2002) The tectonostratigraphy, granitoid geochronology and geological evolution of the Precambrian of southern Ethiopia. J Afr Earth Sci 34:57–84 Yoshida M (1995) Assembly of East Gondwana during the Grenvillian and its rejuvenation during the Pan-African period. Geol Soc India, Mem 34:25–45 Yoshida M, Upreti BN (2006) Neoproterozoic India within East Gondwana: Constraints from recent geochronologic data from Himalaya. Gondw Res 10:349–356 Yoshida M, Santosh M, Rajesh HM (2003) Role of Pan-African events in the Circum-East Antarctic Orogen of East Gondwana: a critical review. In: Yoshida M, Windley BF, Dasgupta S (eds) Proterozoic East Gondwana: Supercontinent Assembly and Breakup. Geol. Soc., London, Special Publ. 206:57–75 Zhao G, Cawood PA, Wilde SA, Sun M (2002) Review of global 2.1– 1.8 Ga orogens: implications for a pre-Rodinia supercontinent. Earth-Sci Rev 59:125–162 Zhao G, Sun M, Wilde SA, Li S (2004) A Paleo-Mesoproterozoic supercontinent: assembly, growth and breakup. Earth-Sci Rev 67:91–123

Tonian/Cryogenian Island Arc Metavolcanics of the Arabian-Nubian Shield

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Moustafa E. Gharib, Ayman E. Maurice, and Hussam A. Selim

Abstract

The Tonian/Cryogenian island arc volcanic rocks have been recorded in the different terranes of the Arabian-Nubian Shield (ANS). They comprise mafic and felsic lavas and their equivalent pyroclastic rocks, which were mostly metamorphosed in greenschist facies. Geochemically, they cover the compositional range from basalt to rhyolite. These metavolcanic rocks belong to the tholeiitic and calc-alkaline magma series. The spider diagrams of these island arc metavolcanics display Nb-Ta and Zr–Hf troughs similar to modern arc magmas, and their REE patterns vary from almost flat to LREE-enriched, where the degree of LREE enrichment is generally increasing from tholeiitic to calc-alkaline rocks. The low-K tholeiitic mafic volcanic rocks were erupted in primitive intra-oceanic island arcs with thin crust, whereas the calc-alkaline mafic volcanic rocks were formed in more mature arcs with thicker crust. The mafic volcanic rocks were generated by partial melting of variably depleted mantle sources, with more significant contribution from the subducted slab in the calc-alkaline magmas produced in mature arcs. The highly silicic volcanics were largely generated by partial melting of amphibolitic lower arc crust. Age-wise, the island arc metavolcanic rocks in the ANS are classified into two groups: >815 Ma and 62% SiO2) plots in the field of dacite and rhyolite (Fig. 11.4 a, b). On the basis of K2O versus SiO2 diagram constructed by Pecerillo and Taylor (1976), the metavolcanic rocks belong mostly to low- to medium-K suites (Fig. 11.4c). This diagram also shows that the andesitic volcanics are more enriched in potassium than basalt and basaltic andesite varieties and few samples of andesite plot at the boundary between medium and high potassium fields. On AFM diagram (Fig. 11.4d), these metavolcanics have a mixed tholeiitic/calc-alkaline characters for the mafic varieties and predominantly calc-alkaline characters for the felsic types. Generally, the MORB-normalized spider diagrams of the metavolcanic rocks from the Eastern Desert are characterized by Nb troughs, similar to the subduction-related magmas (Pearce 1983; Wilson 1989). MORB-normalized trace element patterns of the tholeiitic mafic volcanics of basalt and basaltic andesite display slight to moderate enrichment in LILE (e.g. K, Rb, Ba), negative Ta-Nb anomalies and HFSE (e.g. P, Zr, Ce, Y, Yb) values comparable to or lower than that of the MORB (Fig. 11.5a). The spider diagrams of calc-alkaline basalt, basaltic andesite and andesite are more or less similar to those of the tholeiitic volcanics (Fig. 11.5c). The more enrichment in Rb and Ba and depletion in Sr in the calc-alkaline andesite relative to basalt and basaltic andesite are attributed to fractional crystallization. Samples of the felsic volcanics have Sr, Nb, P and Ti troughs and enriched in K, Rb, Ba and Th relative to the

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Fig. 11.4 a Classification of the metavolcanic rocks in the Eastern Desert of Egypt using total alkali versus silica diagram (Le Bas et al. 1986). The line separating subalkaline and alkaline fields is after Irvine and Baragar (1971). b SiO2 versus Zr/TiO2 diagram (Winchester and Floyd 1977). c Classification of the metavolcanic rocks using K2O versus SiO2 diagram (Pecerillo and Taylor 1976). The fields of older metavolcanics (OMV), younger metavolcanics (YMV) and Dokhan volcanics (DV) of Egypt are after Eliwa et al. (2006). d AFM diagram for the studied rocks. Boundary between calc-alkaline and tholeiitic field after Irvine and Baragar (1971). Data Sources Stern (1981), Gharib et al. (2003), Gharib and Ahmed (2012), Maurice et al. (2012), Abdel-Karim et al. (2019), Faisal et al. (2020)

MORB (Fig. 11.5e). Rhyolite is generally more enriched in K, Rb, Ba and Th and depleted in Ti and P relative to dacite which is attributed to fractional crystallization. The chondrite-normalized REE patterns of the tholeiitic basalt and basaltic andesite are characterized by nearly flat pattern with somewhat depletion LREE (Fig. 11.5b), a distinctive feature of island arc tholeiitic series (Gill 1970, 1981). They have La/Ybn ranging from 0.46 to 0.99, slightly negative to positive Eu anomalies (Eu/Eu* = 0.90–1.15) and unfractionated HREE (Dy/Lun = 0.83–1.21). These patterns are similar to those of basalts of Neoproterozoic primitive intra-oceanic arcs (Gharib and Ahmed 2012; Maurice et al.

2012). Compared to those of the south Eastern Desert, the tholeiitic basalt and basaltic andesite of the central Eastern Desert are more enriched in LREE, which is most probably attributed to more effective role of subduced slab in the genesis of the island arc volcanics of the central Eastern Desert. On the other hand, the REE patterns of the calc-alkaline basalt and basaltic andesite (Fig. 11.5d) are enriched in LREE relative to HREE (La/Ybn = 1.83–3.72), with variable Eu anomalies (Eu/Eu* = 0.87–1.18), and have fractionated LREE (La/Smn = 1.34–2.12) and less fractionated HREE (Dy/Lun = 1.0–1.2). The REE patterns of the calc-alkaline andesites (Fig. 11.5d) display enrichment in

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Tonian/Cryogenian Island Arc Metavolcanics …

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Fig. 11.5 MORB-normalized trace element a, c, e and chondrite-normalized REE b, d, f patterns of Neoproterozoic island arc metavolcanic rocks of south and central Eastern Desert, Egypt. a and b tholeiitic basalt and basaltic andesite. c and d calc-alkaline basalt, basaltic andesite and andesite. e and f felsic volcanic rocks (dacite and rhyolite). MORB and chondrite normalization values after Pearce (1983) and Sun and McDonough (1989), respectively. Data Sources Stern (1981), Gharib and Ahmed (2012), Maurice et al. (2012), Abdel-Karim et al. (2019), Faisal et al. (2020)

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LREE (La/Ybn = 2.23–5.18), somewhat fractionated LREE (La/Smn = 1.6–2.66), less fractionated or almost flat HREE (Dy/Lun = 0.9–1.27) and slightly negative Eu anomalies (0.77–0.98). The felsic volcanics have almost flat REE patterns (La/Ybn = 0.93–1.24; La/Smn = 0.99–1.26; Dy/Lun = 0.8–1.04, Fig. 11.5f) with negative Eu anomalies (Eu/Eu* = 0.56–0.8). These REE patterns are similar to those of Wadi Ranga rhyolites (Gharib and Ahmed 2012; Maurice et al. 2012). HREE are flat indicating that garnet did not control elemental partitioning during melting or fractionation.

11.2.3.2 Tectonomagmatic Discrimination Plots This section assesses the possible fields of the Tonian/Cryogenian island arc volcanic rocks on some of the traditional tectonomagmatic discrimination diagrams. On the SiO2 versus Nb discrimination diagram (Fig. 11.6a), most of the Eastern Desert metavolcanic rocks plotted in the field of volcanic arc magmas with few samples plot in the overlap field between volcanic arc and within plate fields. Similar result was obtained from the application of Rb versus Y + Nb diagram (Fig. 11.6b) for the intermediate and felsic volcanic rocks. Although this diagram was established for granitic rocks (Pearce et al. 1984), the mafic volcanic rocks also plot in the volcanic arc field (Fig. 11.6b). On the TiO2– Zr diagram, all basalt and basaltic andesite spread over the volcanic arc and MORB fields (Fig. 11.6c). Applying the Zr–Th–Nb discrimination diagram (Fig. 11.6d), the tholeiitic basalt and basaltic andesites are largely plotting in the island arc tholeiites (IAT) field, whereas the calc-alkaline basalts and basaltic andesite fall in calc-alkaline basalt (CAB) field. On Th versus Nb diagram of (Saccani 2015, Fig. 11.6e) tholeiitic basalt and basaltic andesite from the south Eastern Desert plot in the field of IAT, tholeiitic basalt and basaltic andesite from the central Eastern Desert straddle the boundary between IAT and BABB fields. The calc-alkaline basalts and andesite plot mainly in the CAB suite. 11.2.3.3 Petrogenesis (i) Tholeiitic Metavolcanics The compositional characteristics of magma generated in oceanic subduction setting are essentially controlled by the relative contribution of three sources—mantle wedge, subducted slab and early—formed arc crust and by crustal processes such as fractional crystallization, crustal contamination and magma mixing (e.g. Wehrmann et al. 2014; Straub et al. 2015; Turner and Langmuir 2015). The concentration of HFSE in the rocks reflects their concentration in the mantle source as they behave in conservative way during subduction. The slab contribution to the mantle wedge can be evaluated by the LILE and LREE as they are non-conservative elements (Pearce 2008; Pearce and Peate 1995).

M. E. Gharib et al.

X/Yb versus Nb/Yb plots (Pearce 1983; Pearce and Peate 1995) are used to evaluate the mantle and slab contribution to the magma produced the tholeiitic metavolcanics, where X is the conservative or non-conservative element. On Zr/Yb and Sm/Yb versus Nb/Yb diagrams (Fig. 11.7a, b), the tholeiitic metavolcanic rocks plot within the mantle array close to N-MORB reflecting derivation from depleted mantle source, which was more depleted than the N-MORB as indicated by lower Nb/Yb values (Pearce and Peate 1995). The tholeiitic basalts and basaltic andesites have Nb/Th (0.05–0.8) values which are significantly lower than that of N-MORB (19, Ohta et al. 1996; Sun and McDonough 1989), supporting derivation from depleted mantle source. The spider diagrams of the tholeiitic mafic volcanic rocks (Fig. 11.5a) generally show depletion in HFSE relative to N-MORB supporting generation from depleted source. The tholeiitic nature of the metavolcanics as well as their LREE-depleted (especially in the tholeiitic basalt of the south Eastern Desert), or nearly flat REE patterns, strong Nb depletion and low K2O and LILE contents suggest that they were developed in an immature island arc setting with no effective role for the subduction components. The low Mg# (34–61, average = 47) of the tholeiitic metavolcanics implies that the magmas produced these volcanics were not primitive, i.e. experienced fractional crystallization prior to eruption. (ii) Calc-Alkaline Metavolcanics The MORB-normalized trace element spider diagrams of the calk-alkaline basaltic andesite and andesite (Fig. 11.5c) have enrichment in the large-ion lithophile elements (LILE) including Rb, Ba and K over high-field strength elements (HFSE) such as Nb, Hf, Zr and Ti. In addition, fluid-mobile elements, e.g. Ba, are notably enriched. The relative enrichment of Ba indicates involvement of fluids released from the subducted slab because Ba is the most mobile element in such fluids (Xiaoming et al. 2007 and references therein). The calc-alkaline nature and low to medium K2O content of the calc-alkaine basaltic andesite and andesites in addition to their LREE and LILE enrichment reflect significant subduction component and generation in mature island arc setting. On La/Yb and Th/Yb versus Nb/Yb diagrams (Fig. 11.7c, d), the calc-alkaline metavolcanic rocks show more enrichment in La and Th relative the tholeiitic varieties, which indicate more significant contribution from the subducted slab, either as fluids or melts. Hydrous mantle melting can produce primary basaltic andesite and andesite magmas (Grove et al. 2012). So, it is proposed that the parental magma of calc-alkaline basaltic andesite and andesites were formed by low-degree hydrous partial melting of mantle peridotite.

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Fig. 11.6 Tectonic setting discrimination diagrams; a SiO2 versus Nb diagram (Pearce and Gale 1977) to discriminate between volcanic arc and within-plate magmas. b Rb versus Y + Nb (after Pearce et al. 1984; Pearce 1996) for metavolcanics of the Eastern Desert. VAG = volcanic arc granite; ORG = oceanic ridge granite; Syn-COLG = syncollision granite and WPG = within-plate granite. c TiO2–Zr diagram (Pearce 1982) to discriminate between volcanic arc basalts, MORB and within-plate basalts. d Th, Hf and Nb discrimination diagram for basalt after Wood (1980). e NbN–ThN diagram (Saccani 2015) for tectonic setting discrimination of basalts. Nb and Th are normalized to N-MORB values of Sun and McDonough (1989). Data Sources Stern (1981), Gharib and Ahmed (2012), Maurice et al. (2012), Abdel-Karim et al. (2019), Faisal et al. (2020)

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Fig. 11.7 a Zr/Yb versus Nb/Yb plot for the volcanic rocks. b Sm/Yb versus Nb/Yb plot for the volcanic rocks. c La/Yb versus Nb/Yb plot for the volcanic rocks. d Th/Yb versus Nb/Yb plot for the volcanic rocks. Mantle arrays after Green (2006). e Al2O3/(FeOt+MgO)–3CaO–5 (K2O/Na2O) ternary source discrimination diagram (after Laurent et al. 2014) for the felsic volcanics. f K2O versus SiO2 diagram (after Gerdes et al. 2000). Data Sources Stern (1981), Gharib and Ahmed (2012), Maurice et al. (2012), Abdel-Karim et al. (2019), Faisal et al. (2020)

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(iii) Felsic Metavolcanics The low-K nature of the felsic metavolcanics with SiO2 > 63 wt% ruled out the possibility of genesis by fractional crystallization of the medium-K basaltic andesite. On the other hand, the low-K affinity and flat REE patterns of these felsic volcanics are similar to those of the tholeiitic metavolcanic rocks suggesting a possible genetic relationship. The field relation between the felsic metavolcanics and the older tholeiitic and calc-alkaline metavolcanics and the available age dating deny derivation of these felsic volcanic rocks through fractional crystallization of these older mafic metavolcanics. Also, the SiO2 contents of the felsic volcanics (>63 wt%) are too high to represent magma that directly derived by partial melting of mantle source, as the mantle melting cannot yield melts more silicic in composition than andesite (Lloyd et al. 1985; Baker et al. 1995; Abdel-Karim et al. 2019). In addition, the high SiO2, Al2O3 and low MgO contents of these rocks are consistent with partial melting of crustal source. The high Y/Nb ratios (10–27) further support their generation from crustal source (Eby 1992). The origin of silicic magmas may be explained by partial melting of low-K amphibolitic lower arc crust (e.g. Brophy 2008; Leat et al. 2006; Tamura et al. 2009). The chondrite-normalized Dy/Lu values for the felsic volcanics are generally ‘4 wt% MgO and ˂4 wt% LOI) revealed that a maximum depth to the Moho increased from south to north Eastern Desert where the estimated thickness is about 6 km for Abu Hamamid, 7.5 km for Darhib area and 8 km for Wadi Ranga area in the south Eastern Desert, and 14 km at Dabbah area, and about 15 km for Wadi Kariem area in the central Eastern Desert. These values confirm the primitive or immature nature of the arc where these volcanics were generated. On the other hand, the median Sr/Y value (Chapman et al. 2015), for samples having SiO2 = 55–70 wt% and 1–6 wt% MgO, and average (La/Yb)n (Profeta and Ducea 2015), for samples having 55– 70 wt% SiO2,˂4 wt% MgO, of calc-alkaline metavolcanics yield arc crustal thickness of 22.0 and 23.9 km, respectively, at Abu Hamamid area in the south Eastern Desert, and 33.6 and 32.5 km, respectively, at Dabbah area in the central Eastern Desert. Thus, the estimated arc crust thickness values support generation of calc-alkaine metavolcanics during more mature stage of arc evolution and indicate that the arc crustal thickness increased from south to north. The medium to high-K calc-alkaline nature and LREE-enriched patterns of the calc-alkaline metavolcanics reflect significant subduction signature and eruption during mature stage of volcanic arc. This means that the tholeiitic metavolcanics represent the first stage of arc volcanicity in Egypt and formed in primitive or immature island arc stage (with a thin

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Fig. 11.8 a La/Yb versus Th/Yb diagram (Condie 1989). b Th/Yb versus Ta/Yb diagram (after Pearce 1982). Data of fields of Izu arc; Tonga arc and Lesser Antilles arc from Xia and Li (2019 and references therein). c Primitive mantle-normalized incompatible trace element spider diagrams for the tholeiitic metavolcanics in the south Eastern Desert. d Primitive mantle-normalized incompatible trace element spider diagrams for calc-alkaline metavolcanics. e Primitive mantle-normalized incompatible trace element spider diagrams for tholeiitic metavolcanics in the central Eastern Desert. Values of primitive mantle after Sun and McDonough (1989). The pattern of Izu arc, Lesser Antilles arc and back arc basin basalts from Xia and Li (2019). Data Sources Stern (1981), Gharib and Ahmed (2012), Maurice et al. (2012), Abdel-Karim et al. (2019), Faisal et al. (2020)

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arc crust), while the calc-alkaline metavolcanics formed later in more mature island arc stages. Applying the median Sr/Y value (Chapman et al. 2015) for the dacitic felsic metavolcanics yields arc crustal thickness of 9.5 km at Abu Hamamid area and 12.46 km at Wadi Ranga area in the south Eastern Desert, and 20.2 km at Hamama area in the central Eastern Desert. In addition, the Rb versus Sr diagram of Condie (1973) is used to determine the crustal thickness at which the felsic metavolcanics of the Eastern Desert were erupted. This diagram (not shown) revealed that felsic metavolcanics were formed under crustal thickness between 15 and 20 km, which is broadly comparable to the arc crustal thickness estimated from the median Sr/Y value.

11.3

Island Arc Volcanism in the Arabian Shield

Metamorphosed island arc volcanic rocks are reported from northwestern and southwestern Saudi Arabia (e.g. Reischmann et al. 1983; Ali et al. 2010c; Volesky et al. 2017). The greenschist facies arc metavolcanics of Ghawjah Formation at Wadi Sawawin area, Northwestern Saudi Arabia, are composed of basalts and porphyritic andesites with subordinate dacites (Ali et al. 2010c). Plagioclase has altered cores, whereas clinopyroxene is altered to actinolite and chlorite. In the amygdaloidal varieties, the amygdales are filled with chlorite, calcite and quartz. Geochemically, the Ghawjah arc metavolcanics range in composition from basalts to dacites (Fig. 11.9a, b), which are classified as low- to medium-K rocks (Fig. 11.9c). Ali et al. (2010c) proposed that they are tholeiitic and calc-alkaline; however, in terms of their Zr and Y contents, they broadly have tholeiitic and transitional affinity (Fig. 11.9d). Their primitive mantle-normalized spider diagrams are characterized by Nb-Ta troughs (Fig. 11.10a, c, e), which are more pronounced in the calc-alkaline and dacitic rocks (Ali et al. 2010c). The chondrite-normalized REE patterns are generally LREE-enriched (Fig. 11.10b, d, f), with comparable La/Ybn values in basalts (0.80–3.1), andesites (mostly 1.2–3.6) and dacites (2.7–4.0). Primitive volcanic samples do not contain abundant phenocrysts and are distinguished by high Mg# (up to 67) and Ni (up to 175 ppm) and Cr (up to 537 ppm) contents. As to their setting, Ali et al. (2010c) suggested that the Ghawjah metavolcanics were erupted in an intra-oceanic arc setting. The metavolcanic rocks (schists) hosting the Shaib al Tair and Rabathan volcanogenic massive sulphide deposits, Wadi Bidah Mineral District, Southwestern Saudi Arabia, are considered to be formed in oceanic island arc setting similar to Kermadec arc (Volesky et al. 2017 and references therein). These rocks are basaltic to felsic in composition and have been sheared by post-mineralization deformation

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(Volesky et al. 2003). They have been regionally metamorphosed to greenschist facies. The Wadi Bidah basaltic metavolcanic rocks show tholeiitic affinity (Volesky et al. 2017 and references therein). The felsic metavolcaniclatic rocks are the host of volcanogenic massive sulphide deposits (Volesky et al. 2003). U–Pb zircon dating of Ghawjah andesites, northwestern Saudi Arabia, yielded a weighted mean 206Pb/238U age of 763 ± 25 Ma (Ali et al. 2010c), which considered the eruption age of these volcanics. The mean Nd model age of these volcanics is 0.71 Ga, which is comparable to the U– Pb crystallization age. The age of Ghawjah intra-oceanic arc volcanics is similar to that reported for the Neoproterozoic island arc volcanics of the central Eastern Desert of Egypt (* 750 Ma, Ali et al. 2009). On the other hand, the oceanic arc felsic metavolcaniclastic rocks hosting the volcanogenic massive sulphide deposits of Wadi Bidah, southwestern Saudi Arabia, have older age (815–855 Ma, Volesky et al. 2017 and references therein), which is comparable to that of the older arc rocks of the southern Nubian Shield (see below). In addition, the arc rhyolitic volcanics of Al-Lith area, southwestern Saudi Arabia, dated at 847 ± 34 Ma (Kröner et al. 1983). It is worth to mention that Kröner et al. (1991) proposed that the island arc volcanic rocks of Al-Lith area (Reischmann et al. 1983), southwestern Arabian Shield and the comparable volcanics of southeastern Sudan (see below) constitute parts of the same island arc terrane, which were separated by opening of the Red Sea. The Ghawjah oceanic arc volcanic rocks have strongly positive eNd(t) values (+5.4 to +8.2, average = + 6.8, Ali et al. 2010c), reflecting derivation from depleted mantle peridotite source. Primitive samples indicate eruption without significant fractionation, implying short residence time or thin arc crust. The enrichment of these rocks in LILE was attributed to the contribution of subducted slab either as fluids or subducted sediments (Ali et al. 2010c). Because they have similar REE and HFSE ranges, the mafic and felsic arc metavolcanics of Wadi Bidah in southwestern Saudi Arabia are not considered cogenetic, where the felsic rocks are likely to be formed by partial melting of lower mafic crust in the presence of amphibole (Volesky et al. 2017).

11.4

Island Arc Volcanism in the Southern Nubian Shield

11.4.1 Island Arc Metavolcanics of Sudan The island arc metavolcanics of Gebeit Mine (NE Sudan) include porphyritic mafic to felsic lava flows. The mafic/ intermediate lavas are plagioclase- and/or pyroxene-phyric,

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Fig. 11.9 Classification and magmatic affinity of the Neoproterozoic island arc volcanics of Saudi Arabia, Sudan, Ethiopia and Eritrea. a Na2O + K2O versus SiO2 diagram (after Le Bas et al. 1986) with the dividing curve between subalkaline and alkaline rocks after Irvine and Baragar (1971); b SiO2 versus Zr/TiO2 diagram (after Winchester and Floyd 1977); c K2O versus SiO2 diagram (after Rickwood 1989); d Y versus Zr diagram (after Barrett and MacLean 1994). Data Sources Saudi Arabia (Ali et al. 2010c), Sudan (El-Nadi 1989; Kröner et al. 1991; Reischmann and Kröner 1994; Ibinoof et al. 2016), Ethiopia (Tadesse et al. 1999; Alene et al. 2000; Tadesse and Allen 2004; Sifeta et al. 2005), Eritrea (Woldehaimanot 2000; Teklay et al. 2002a, b)

whereas the felsic volcanics contain alkali feldspar and/or quartz phenocrysts (El-Nadi 1989). Later on, Reischmann and Kröner (1994) classified the island arc volcanics of Gebeit area into four groups: clinopyroxene-phyric basalts containing clinopyroxene and plagioclase phenocrysts, hornblende-phyric basalts, plagioclase-phyric andesites and aphyric basalts. On the other hand, the metavolcanic rocks of Khor Ashat and southeast of Tokar areas (southern Sudan) are essentially rhyolites (Kröner et al. 1991). Geochemically, the arc volcanic rocks of Sudan cover the range from basalt to rhyolite (Fig. 11.9a, b). In terms of their K2O contents, they are classified as low- to high-K and shoshonitic rocks (Fig. 11.9c). On the Y versus Zr diagram, they have tholeiitic

to calc-alkaline nature (Fig. 11.9d). On primitive mantlenormalized trace element patterns (Fig. 11.11a, c, e), the Sudanese island arc volcanics show Nb–Ta troughs characteristic of subduction-related magmas. Their chondritenormalized REE patterns (Fig. 11.11b, d, f) are enriched in LREE relative to HREE, with comparable La/Ybn values in basalts (2.0–7.5) and andesites (1.9–6.1) and more fractionated patterns in felsic volcanics (La/Ybn = 7.3–14.3). Available dates of the island arc metavolcanics of Sudan fall essentially in two groups: *710–720 Ma and >830 Ma. Basic to acidic arc metavolcanics from NE Sudan were dated at 712–720 Ma (El-Nadi 1989 and references therein). In addition, Sm/Nd data of Gebeit island arc basalts (NE

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Fig. 11.10 Primitive mantle-normalized trace element a, c, e and chondrite-normalized REE b, d, f patterns of Neoproterozoic island arc volcanic rocks from the Arabian Shield. Primitive mantle and chondrite normalization values after Sun and McDonough (1989). Data Source Ali et al. (2010c)

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Fig. 11.11 Primitive mantle-normalized trace element a, c, e and chondrite-normalized REE b, d, f patterns of Neoproterozoic island arc volcanic rocks of Sudan, Nubian Shield. Primitive mantle and chondrite normalization values after Sun and McDonough (1989). Data Sources El-Nadi (1989), Reischmann and Kröner (1994), Ibinoof et al. (2016)w

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Sudan) yielded and isochron age of 832 ± 26 Ma, which is considered the best estimate for the age of eruption (Reischmann and Kröner 1994). However, Sm/Nd whole-rock dating of the Abutulu calc-alkaline arc/back arc meta-andesite produced 768.31 ± 123.36 Ma (Ibinoof et al. 2016). The acid metavolcanic rocks of southeast of Tokar area (near Eritrean border) gave 207Pb/206Pb age of 840– 855 Ma (Kröner et al. 1991). The arc metavolcanics and associated intrusive rocks of Erkowit-Khor Ashat and southeast of Tokar areas are considered one of the oldest arc terranes in the entire ANS (840–870 Ma, Kröner et al. 1991). Based on U–Pb zircon dating of the Aquaba Pass island arc rhyolite (NE Sudan), Reischmann et al. (1992) proposed that the timing of the island arc volcanic eruption can be constrained between *887 and 868 Ma. On the other hand, Ibinoof et al. (2016) suggested that arc magmatism in southern Sudan was active at around 778 ± 90 Ma. El-Nadi (1989) suggested mature island arc setting for the largely calc-alkaline arc metavolcanic rocks of NE Sudan. These metavolcanics were interpreted to be formed by partial melting of modified mantle peridotite source followed by extensive fractional crystallization. The eNd(t) values of clinopyroxene- and hornblende-phyric basalts (+6.74 ± 0.19) and aphyric basalts and plagioclase-phyric andesites (+6.7 to +8.3) are consistent with different depleted mantle sources and oceanic island arc origin of these volcanics (Reischmann and Kröner 1994). On the other hand, the volcanic rocks of Abutulu unit of southern Sudan have lower eNd(t) value of +5.8 ± 2.5 (Ibinoof et al. 2016).

11.4.2 Island Arc Metavolcanics of Ethiopia The metavolcanic rocks of Tsaliet Group in Werri area, northern Ethiopia, are composed of porphyritic lava flows and pyroclastic rocks, which essentially have basaltic and andesitic composition with rare dacitic samples (Sifeta et al. 2005). On the other hand, the metavolcanic rocks of Mai Kenetal-Negash area, Tigrai, northern Ethiopia, comprise two units: basic/intermediate metavolcanics and acidic metavolcanic/metavolcaniclastics units (Alene et al. 2000). In the basic/intermediate metavolcanics, porphyritic and locally amygdaloidal textures are still recognizable, and relics of pyroxene are preserved (Alene et al. 2000). In the Tuludimtu orogenic belt, western Ethiopia, the mafic to felsic metavolcanic rocks of Dengi (and Sirkole?) Domain are considered to be developed in a convergent setting (Allen and Tadesse 2003; Tadesse and Allen 2004). In Axum area (northern Ethiopia), low-grade island arc metavolcanic rocks are exposed in Adi Nebrid and Adwa belts (Tadesse et al. 1999). In the Adi Nebrid belt, the island arc metavolcanics are basic to intermediate lava flows and pyroclastics with volcaniclastic rocks, whereas in Adwa

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block they are represented by porphyritic basalt and basaltic andesites, pyroclastics and volcaniclastics. The felsic volcanic rocks are subordinate in the two belts. According to the available geochemical data, the arc volcanic rocks of Ethiopia broadly have basaltic to andesitic composition (Fig. 11.9a and b), with subordinate felsic rocks. They are largely classified as low- to high-K rocks (Fig. 11.9c), where the higher K2O contents in some mafic volcanics may reflect modification of magmatic abundances due to hydrothermal alteration. On the Y versus Zr diagram, they broadly have tholeiitic to transitional affinity (Fig. 11.9 d). Similar to modern arc magmas, the spider diagrams of the Ethiopian island arc metavolcanic rocks (Fig. 11.12a, c, e) are characterized by Nb-Ta or Nb with Zr-Hf troughs. The chondrite-normalized REE patterns (Fig. 11.12b, d, f) of these volcanics show enrichment in LREE, with comparable and relatively wide ranges of La/Ybn values in basalts (2.1– 8.7), andesites (2.0–9.5) and dacite/rhyolite (2.2–8.2). The geochemical data of the essentially basaltic and andesitic lava flows and tuffs of Dengi Domain of Tuludimtu orogenic belt are consistent with the development in a volcanic arc setting (Tadesse and Allen 2004). The Werri area metavolcanic rocks are subalkaline (tholeiitic and calcalkaline) and similar to volcanic rocks erupted in developed island arcs (Sifeta et al. 2005). Similarly, the metavolcanic rocks of Mai Kenetal-Negash area are predominantly calc-alkaline and akin to volcanic rocks produced during the mature stage of island arc evolution (Alene et al. 2000). Moreover, Tadesse et al. (1999) proposed that the island arc volcanics of Axum area were erupted at different stages of island arc evolution. The low-K tholeiitic mafic/intermediate volcanics of Adi Nebrid belt represent Neoproterozoic volcanism in an immature intra-oceanic island arc, whereas the calc-alkaline essentially mafic/intermediate metavolcanics of Adwa belt to the east were generated at evolved stages of an island arc. In northern Ethiopia, the metavolcanic rocks of Werri area have mean TDM model age of 960 Ma and initial eNd(800) range from +3.8 to + 4.9 (Sifeta et al. 2005), suggesting derivation from juvenile Neoproterozoic mantle, whereas the magmas of the metavolcanic rocks of Mai Kenetal-Negash area were most probably generated from depleted source (Alene et al. 2000).

11.4.3 Island Arc Metavolcanics of Eritrea The arc metavolcanic rocks of northern Eritrea comprise calc-alkaline basic and felsic lava flows and crystal and lapilli tuffs (Teklay et al. 2002a). In the Adobha belt of northern Eritrea, the Himbol volcanic rocks of Nakfa terrane and the Hager terrane to the west are considered to belong to a calc-alkaline island arc setting (Woldehaimanot 2000). The

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Fig. 11.12 Primitive mantle-normalized trace element a, c, e and chondrite-normalized REE b, d, f patterns of Neoproterozoic island arc volcanic rocks of Ethiopia, southern Nubian Shield. Primitive mantle and chondrite normalization values after Sun and McDonough (1989). Data Sources Alene et al. (2000), Sifeta et al. (2005)

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Fig. 11.13 Primitive mantle-normalized trace element a, c, e and chondrite-normalized REE b, d, f patterns of Neoproterozoic island arc volcanic rocks of Eritrea, southern Nubian Shield. Primitive mantle and chondrite normalization values after Sun and McDonough (1989). Data Sources Woldehaimanot (2000), Teklay et al. (2002a, b)

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Himbol volcanic rocks are composed of mafic to felsic metavolcanics and dacitic to rhyolitic tuffs (Woldehaimanot 2000). On the other hand, the arc metavolcanic rocks of central Eritrea have low-K tholeiitic nature and range in composition from basalt to rhyolite (Teklay et al. 2002b). Mafic to felsic island arc metavolcanic rocks are also recorded in east-central Eritrea (Andersson et al. 2006). These metavolcanic rocks, as well as other meta-igneous rocks, show typical oceanic arc geochemical characteristics ranging from primitive to more evolved. Collectively, the island arc volcanic rocks of Eritrea range in composition from basalt to rhyolite (Fig. 11.9a, b) and are largely classified as low- to medium-K rocks (Fig. 11.9c). As to magmatic affinity, they are tholeiitic to calc-alkaline rocks (Fig. 11.9d). The spider diagrams of these metavolcanic rocks (Fig. 11.13a, c, e), with Nb and Zr-Hf troughs, are similar to those of modern island arc magmas. The Eritrean island arc metavolcanics have LREE-enriched chondrite-normalized REE patterns (Fig. 11.13b, d, f), with broadly comparable La/Ybn values (mostly 2.2–4.2 in basalts, 2.5–5.3 in andesites and 1.7–6.7 in dacite/rhyolite). SHRIMP U–Pb zircon dating of arc metarhyolite from Nakfa area, northern Eritrea, yielded a concordant age of 854 ± 3 Ma (Teklay et al. 2002a). Dating of subductionrelated magmatic rocks from northern and central Eritrea indicates that major arc magmatism of Eritrea began at *850 Ma (Teklay et al. 2002a & b; Teklay 2006). This age is consistent with or overlaps that of the earlier (>*830 Ma) island arc rocks of Sudan. The low-K tholeiitic nature and high initial eNd(t) values (up to +9.0) of the arc metavolcanic rocks of central Eritrea are consistent with depleted mantle source for the parental magmas of these metavolcanics (Teklay et al. 2002b), whereas the mantle source of the calc-alkaline arc basic-felsic Nakfa metavolcanic rocks of northern Eritrea was modified by plume prior to subduction as indicated by lower eNd(t) values (+4.8 to +5.7, Teklay et al. 2002a). Moreover, the island arc low-K tholeiitic rocks of central Eritrea have higher initial eNd(t) values (+5.3 to +9.0) than associating boninites (+4.1 to +5.1), reflecting an increase in contribution of subduction components, such as subducted sediments and/or altered oceanic crust, to the mantle source of boninitic volcanics (Teklay et al. 2002b). The variations in source mantle fertility of the Eritrean island arc volcanics imply heterogeneity in the composition of the Neoproterozoic mantle.

11.5

Conclusions

The Tonian/Cryogenian island arc volcanic rocks comprise mafic and felsic lavas and their equivalent pyroclastic rocks, which were mostly metamorphosed in greenschist facies. They cover the compositional range from basalt to rhyolite,

and their geochemical characteristics are similar to modern tholeiitic and calc-alkaline arc magmas, which erupted in primitive and mature arcs, respectively. The mafic island arc volcanics of ANS were derived from variably depleted mantle sources modified by subduction components contributed as subducted slab-derived fluids and/or subducted sediments, where the role of subduction components was more pronounced in the calc-alkaline volcanics. However, geochemical evidence indicated source enrichment by mantle plumes prior subduction in some southern Nubian Shield arcs. On the other hand, arc felsic volcanics are assumed to be formed by partial melting of lower mafic arc crust. Age-wise, the Tonian/Cryogenian island arcs of ANS are classified into two groups:>815 Ma and 0.89); enriched in Ni, Cr, and Co; and depleted in Al2O3 and CaO. These features are most consistent with the general consensus that the ANS ophiolites all formed in supra-subduction zone settings. However, authors continue to disagree on the detailed setting, i.e., fore-arc or back-arc spreading centers. We contend that the evidence strongly favors the fore-arc setting for the majority of the ANS ophiolites. Specifically, in serpentinites, the relics of Cr-spinel have Cr# mostly >60, and the relics of olivines and orthopyroxene are Mg-rich, both suggesting that the serpentinites are residual to high degrees of melt extraction, as found most commonly (though not exclusively) in modern fore-arc peridotites. The ophiolitic sequences of the ANS are worthy targets for mineral exploration. They host a variety of important mineral deposits, including chromite, magnesite, talc, platinum-group elements, base metals (Cu–Ni–Co), and gold. There is a spatial and genetic relationship between carbonatized ultramafics, subsequent granite intrusions, and gold mineralization. Keywords

















Arabian-Nubian Shield Ophiolites Supra-subduction zone Spinel Fore-arc setting Alteration Mineralization

© The Author(s), under exclusive license to Springer Nature Switzerland AG 2021 Z. Hamimi et al. (eds.), The Geology of the Arabian-Nubian Shield, Regional Geology Reviews, https://doi.org/10.1007/978-3-030-72995-0_12

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12.1

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Introduction

The Arabian-Nubian Shield (ANS) may be the largest tract of juvenile continental crust of Neoproterozoic age on Earth (Stern 2002; Patchett and Chase 2002; Hargrove et al. 2006). It extends from Jordan and southern Israel in the north to Eritrea and Ethiopia in the south and from Egypt in the west to Saudi Arabia in the east (Fig. 12.1). Prior to Red Sea-Gulf of Aden opening, the Arabian and Nubian Shields were conjoined as a continuous tract of juvenile Neoproterozoic crust. The ANS is the northern continuation of the Mozambique belt, and together, they have been referred to as the East African Orogen (Stern 1994), formed by collision between East and West Gondwana during the Pan-African orogeny (Stern 1994; Dilek and Ahmed 2003). The Pan-African orogeny was a series of major Neoproterozoic orogenic events through which continental, island arc, and oceanic terranes were brought together to form the crystalline basement of the Gondwana supercontinent. It is the largest known system of orogenies on Earth and produced more continental crust than any other (Rino et al. 2008). The ANS is an excellent natural laboratory for studying the process of major crustal accretion. It is a collage of Neoproterozoic juvenile arcs, younger sedimentary and volcanic basins, voluminous granitoid intrusions, and enclaves of pre-Neoproterozoic crust that crop out in the western Arabian Plate and the northeastern African Plate (Kröner et al. 1991; Reischmann and Kröner 1994; Stern 1994; Kusky et al. 2003; Johnson and Woldehaimanot 2003; Abdallah et al. 2019). The stabilization of the ANS occurred before the development of an extensive peneplain in mid-Cambrian times (*520 Ma) and was exhumed in the Neogene as a consequence of Red Sea rifting and flank uplift. The latest tectonic event in the ANS (starting *30 Ma ago) was the formation of the Red Sea and related structures such as the Gulf of Suez and the Gulf of Aqaba. The abundance of ophiolites and ophiolitic mélanges is a distinctive part of the ANS, and these terranes are key components of the shield, providing important clues about its origin and hosting much of its mineralization. They represent fragments of upper mantle and ancient oceanic crust that were obducted and tectonically emplaced onto older continental margins during the closure of the Mozambique Ocean (e.g., Stoeser and Camp 1985; Abdelsalam and Stern 1996; Shackleton 1996; Stern et al. 2004; Azer and Stern 2007; Ali et al. 2010; Abuamarah et al. 2020). In the ANS, ophiolites are concentrated along strongly deformed linear belts that represent the sutures of closure of the oceanic basins and typically juxtapose juvenile terranes and older continental terranes (e.g., Stoeser and Camp 1985; Johnson et al. 1987; Abdelsalam and Stern 1996). These ophiolitic sutures represent relics of ancient

subduction zones, and the tectonic events associated with suturing inevitably left the ophiolite sequences in the ANS dismembered, deformed, altered, and metamorphosed. The authors of this contribution have focused their work mostly on the ultramafic mantle sections of the ANS ophiolites, and so this chapter, after providing an overview of the Neoproterozoic ophiolites of the ANS, will then concentrate mostly on their mantle sections and on the conspicuous alteration and mineral potential of these ultramafic units. Alteration results from circulation of hydrothermal fluids and can—depending on protolith, fluid composition, fluid/rock ratio, and temperature—produce various assemblages including serpentinite, listvenite, rodingite, magnesite, and talc-carbonate rocks. Carbonate alteration resulted in veins and dikes of calcite, dolomite, ankerite, magnesite and breunnerite. The migration of these CO2-bearing solutions also resulted in diffuse and pervasive carbonation of a wide range of the Neoproterozoic rocks of the ANS. The ophiolites of the ANS host important mineral deposits, such as gold, chromite, magnesite, and talc (e.g., Klemm et al. 2001; Kusky and Ramadan 2002; Azer 2013), associated in many cases with alteration of their ultramafic sections.

12.2

Tectonic History of the Arabian-Nubian Shield

The abundance of ophiolites, ophiolitic mélanges, and island arc assemblages in the ANS testifies to the plate-tectonic processes that led to the assembly of the shield. Several models for the structure and evolution of the ANS have been proposed (e.g., El-Gaby et al. 1988; Stern 1994; Abdelsalam and Stern 1996; Blasband et al. 2000; Johnson et al. 2011), but a common feature of these models is an early stage of accretion of island arcs and oceanic terranes (e.g., Vail 1985; Stoeser and Camp 1985; Harris et al. 1990; Abdelsalam and Stern 1996; Johnson and Kattan 2001). Each of these models is informed by evidence gained from study of the ANS ophiolites and, in turn, offers explanations for their presence, their supra-subduction zone (SSZ) character, their metamorphism, and their fragmentation. The geological history of the ANS was divided by Bentor (1985) into four main phases: (1) oceanic ophiolite phase (*1000 Ma), recorded by typical members of the ophiolite assemblage; (2) island arc phase (*950–650 Ma), recorded by tonalites, trondhjemites, gabbros and island arc-related volcano-sedimentary successions; (3) post-collisional batholithic phase (*640–590 Ma), recorded by subaerial, medium- to high-K calc-alkaline andesite to rhyolite, plutonic equivalents, and volcaniclastic sediments; and (4) anorogenic within-plate alkaline phase (*590–550 Ma), recorded by alkaline, high-level granites and syenites and their volcanic equivalents (e.g., the Katharina Province).

12

Neoproterozoic Ophiolites of the Arabian-Nubian Shield

299

Fig. 12.1 Simplified map of the Arabian-Nubian Shield (Johnson and Woldehaimanot 2003). Basement outcrops are white; Phanerozoic cover is shown in yellow; structural trends are highlighted; ophiolitic rocks are shown in black; and gneissic rocks are shown in stipple

El-Gaby et al. (1988) suggested a classic and influential model for the evolution of the Pan-African belt in the northern part of the ANS, especially the Eastern Desert and Sinai (Fig. 12.2). In this model, they depicted an ensimatic island arc accreted onto an attenuated and reactivated older continental margin. The old continent was fringed by an island arc about 800–700 Ma ago (Fig. 12.2a). Somewhat later, probably around 700 Ma ago, the island arc was swept against the old continent, thereby thrusting ophiolites and the island arc volcanic and volcaniclastic sequences onto the margin of the old continent (Fig. 12.2b). The Pan-African belt then acquired a Cordilleran character (Fig. 12.2c). Intrusion of subduction-related and mantle-derived magmas would have induced softening and remobilization of the early Neoproterozoic continental crust or “infrastructure” around 655 Ma ago. The presence

of inverted density gradients between the overthrust sheets and the underlying remobilized granitoids of the infrastructure could then have caused Rayleigh–Taylor instability and the formation of large-scale undulations (Fig. 12.2d) that are reflected on the surface as geanticlines cored by buoyant diapirs of granitoid material. Calc-alkaline magmas erupted along the geanticlines. The cratonized basement was intermittently intruded by sub-alkaline to peralkaline silicic rocks after the end of the Pan-African orogeny around 570 Ma ago, continuing most probably into the Cenozoic. Abdelsalam and Stern (1996) also modified the Bentor (1985) history (Fig. 12.3), again dividing the formation of the ANS into four major tectonomagmatic episodes between about 900 and 550 Ma. The initial rifting of supercontinent Rodinia (*900–850 Ma, not shown in Fig. 12.3) was

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Fig. 12.2 Cartoon illustrating the tectonic evolution of the late Neoproterozoic rocks of the northwest ANS (El-Gaby et al. 1988): a island arc stage, b overthrusting of ophiolites and island arc volcanics and volcaniclastics over the old continental margin, c Cordilleran stage, and d buoyancy-driven modification of crustal structure in the late Cordilleran stage

followed by seafloor spreading to open the Mozambique Ocean, attended by formation and accretion of fringing arc and back-arc basins to form the juvenile material of the ANS. Terrane accretion resulted in the formation of arc-arc sutures between *800 and 700 Ma (Fig. 12.3a). Closure of the Mozambique Ocean (*750–650 Ma) led to collision of the ANS with rifted block fragments of East and West Gondwanaland along north-trending arc-continental sutures. The continuation of convergence of East and West Gondwana

resulted in crustal shortening in the ANS, localized along north-trending linear belts (Fig. 12.3b). The shortening deformation culminated with the formation of major northwest-trending sinistral strike-slip fault systems (with minor conjugate northeast-trending dextral faults, not shown) at *640–540 Ma (Fig. 12.3c). Johnson et al. (2011) adopted a similar four-stage scheme for ANS development between about 870 and 550 Ma (Fig. 12.4), introducing the ideas of escape tectonics and orogenic collapse during the final stage.

Fig. 12.3 Cartoon outlining the tectonic evolution of deformation belts in the ANS: a development of the ANS as a composite terrane made-up of accreted arcs; b closing of the Mozambique Ocean to produce arc-continental sutures and north-trending shortening zones; c the culmination of the north-trending shortening zones in the formation of northwest-trending sinistral strike-slip systems (Abdelsalam and Stern 1996)

12 Neoproterozoic Ophiolites of the Arabian-Nubian Shield 301

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Fig. 12.4 Schematic illustration of stages in the development of ANS showing its setting in the supercontinent cycle bracketed by the break-up of Rodinia and the assembly of Gondwana (Johnson et al. 2011)

12.3

Components of the ANS Ophiolites

Ophiolites are remnants of oceanic crust; complementary studies of modern ocean crust and ancient examples in ophiolite sequences have led to a unified understanding of the typical sequences of rock types and their typical thicknesses in both normal and atypical ocean basins (Dilek and Newcomb 2003). Thrust contacts are documented at the base of some, but not all, ANS ophiolites. The metamorphic soles of ANS ophiolites have been studied along the western part of the Allaqi-Heiani ophiolite belt in the South Eastern Desert of Egypt, where El Naby and Frisch (1999) inferred temperatures of up to 700 °C and pressures up to 0.8 GPa. However, thrusts at the base of some other ophiolites show no obvious thermal effects. Reconstruction of the original sequences and structures of the ANS ophiolites shows, at least approximately, all the characteristics of oceanic crust, and thus are remnants of Neoproterozoic seafloor (Stern et al. 2004). All the ANS ophiolites, however, have been dismembered to varying extents; some are found only as blocks within tectonic mélanges. Locally, in the best-preserved examples, one

finds the complete “Penrose” ophiolitic sequence, including a lower “mantle” unit of serpentinized ultramafic rocks and an upper “crustal” unit of layered and isotropic gabbros, sheeted dikes, pillow lavas, and pelagic deep-sea sedimentary rocks. Typical ophiolite sequences found in the ANS are listed in Table 12.1. Several ANS ophiolites have sheeted dike complexes (such as Ghadir, Onib, and Ess), but in a number of other cases, dike complexes are missing. A reconstructed pseudostratigraphy through a complete ANS ophiolite section is given in Fig. 12.5. The base of most ANS ophiolites consists of serpentinized peridotite, often expressing strongly deformed or transposed compositional layering, forming peridotite tectonite. The protoliths of the serpentinites were dominantly harzburgite and dunite. Rarely, the lowest unit in some ANS ophiolites may contain minor lherzolite, generally interpreted as fertile mantle that escaped major melt extraction (although refertilization may also be responsible). The harzburgite is generally interpreted to be the depleted mantle from which overlying mafic rocks were derived, and the deformation is related to the extensional strain across the ridge axis being accommodated in shear zones within the

12

Neoproterozoic Ophiolites of the Arabian-Nubian Shield

Table 12.1 List of complete “Penrose-type” ophiolites in the Arabian-Nubian Shield (updated after Stern et al. 2004)

303

Name

Country

References

Gebel Gerf

Egypt

Zimmer et al. (1995)

Wadi Ghadir

El Bayoumi (1983)

Abu Dahr

Gahlan et al. (2015)

Fawakhir

El-Sayed et al. (1999)

Sol Hamed

Nasr and Beniamin (2001), Abu-Alam and Hamdy (2014)

Oshib-Ariab belt

Sudan

AbdelmRahman (1993)

Arbaat

Abdelsalam and Stern (1993)

Atmur-Delgo

Harms et al. (1994), Schandelmeier et al. (1994)

Sol Hamed

Fitches et al. (1983), Nasr and Beniamin (2001)

Wadi Onib Bir Tuluhah

Hussein et al. (2004) Saudi Arabia

Pallister et al. (1988), Abuamarah et al. (2020)

Darb Zubaydah

Quick (1990)

Wadi Khadra

Quick (1991)

Halaban

Al-Saleh et al. (1998), Abuamarah (2019b)

Ess

Pallister et al. (1988), Gahlan et al. (2020a)

Al ‘Ays (Wask)

Bakor et al. (1976), Gahlan et al. (2020b)

Tharwah

Nassief et al. (1984)

Bir Umq

Shanti (1983), Abuamarah (2019a)

Tathlith

Pallister et al. (1988)

harzburgite. Resting above the harzburgite are layered ultramafic cumulates including dunite and pyroxenite with pods of chromitite (Gahlan et al. 2012, 2020a, b). These units are interpreted as early cumulates or products of melt-rock reaction formed by primitive magmas extracted from the underlying harzburgites that cool and react as they ascend. The base of the crust is conventionally placed between the depleted harzburgite below and the layered ultramafic cumulates above (Gahlan et al. 2015), although other authors would place it higher, at the first appearance of abundant plagioclase in the rocks. The layered ultramafic cumulates grade upward into a thick unit of strongly layered gabbro. Individual layers within this unit may include pyroxenite and anorthosite. The layered gabbro is succeeded upward by isotropic gabbro, which may (despite its name) express some minor or cryptic layering. The next unit in the complete ophiolite sequences is typically a sheeted dike complex. The transition zone from gabbro to dikes often shows complex intrusive relations, with each lithology apparently able to intrude the other. Distinctive pods or dikes of felsic rock called plagiogranite tend to occur near the gabbro-dike transition; these often contain zircon and are popular targets, when present, for U-Pb geochronology studies. In ideal cases, each diabase dike intrudes into the center of the previously intruded dike, forming a sequence of dikes that have chilled margins developed only on one side. The sheeted dikes represent magma conduits that fed basaltic flows at the surface. These

flows are typically pillowed, with lobes and tubes of basalt forming bulbous shapes diagnostic of underwater basaltic volcanism. Some ophiolites of the ANS are overlain by deep-sea pelagic sediments.

12.4

Geochronology

Because the ophiolites record the key events of the seafloor spreading and terrane accretion stages in the assembly of the ANS, the absolute ages of emplacement and metamorphism of the ophiolites are essential elements of the overall tectonic history of the shield. The available radiometric ages for the ANS ophiolitic rocks are listed in Table 12.2. The ophiolitic rocks of the ANS have ages ranging from 670 to 870 Ma. The ANS ophiolites have been dated using U-Pb zircon techniques (Stacey et al. 1984; Pallister et al. 1988; Ali et al. 2010) and Pb–Pb zircon evaporation techniques (Kröner et al. 1992; Zimmer et al. 1995; Loizenbauer et al. 2001) on zircons separated from gabbros and plagiogranites. Other ages have been determined using Sm–Nd mineral and whole-rock isochron techniques (Claesson et al. 1984; Zimmer et al. 1995; Worku, 1996). Ali et al. (2010) suggested two stages for the evolution of the Ess-Yanbu-Onib-Sol Hamed-Gerf-Allaqi-Heiani (YOSHGAH) ophiolite belt (*810–780 Ma and *750–730 Ma), which extends from the Arabian Shield across the Red Sea into the Nubian Shield sector of the ANS. They concluded

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M. G. Shahien et al.

Fig. 12.5 Schematic lithostratigraphic section of the ANS ophiolites (not to scale)

that accretion between the Gabgaba-Gebeit-Hijaz terranes to the south and the South Eastern Desert-Midyan terranes to the north occurred as early as 730 Ma and no later than

709 ± 4 Ma. In the Ali et al. (2010) chronology, the igneous ages of the ANS ophiolites are slightly older than the island arc stage (*770 to *670 Ma; Stern et al. 1991; Ali et al.

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Neoproterozoic Ophiolites of the Arabian-Nubian Shield

Table 12.2 Age determinations of Arabian-Nubian Shield ophiolites

305

Saudi Arabia Locality

Material, rock type

Method

Age (Ma)

References

Bir Umq

Diorite

U-Pb

838 ± 10 Ma

Pallister et al. (1988)

Ess

Al-Wask

Plagiogranite

U-Pb

782 ± 5 Ma

Pallister et al. (1988)

Plagiogranite

U-Pb

764 ± 3 Ma

Pallister et al. (1988)

Basalt

Rb-Sr whole rock

831 ± 47 Ma

Dunlop et al. (1986)

Pyroxene, gabbro

Sm–Nd

828 ± 47 Ma

Dunlop et al. (1986)

Whole rock and mineral

Sm–Nd

782 ± 38 Ma

Claesson et al. (1984)

Zircon, isotropic gabbro

U-Pb

706 ± 11 Ma

Pallister et al. (1988)

Gabbro and trondhjemite

U-Pb

740–770 Ma

Pallister et al. (1988)

Thurwah

Gabbro

U-Pb

870 ± 11 Ma

Pallister et al. (1988)

Halaban

Gabbro

U-Pb

694 ± 8 Ma

Stacey et al. (1984)

Metamorphic sole

40

680 Ma

Al-Saleh et al. (1998)

Halaban

Metagabbro

40

679–681 Ma

Al-Saleh et al. (Al-Saleh et al. 1998)

Bir Tuluhah

Plagiogranite

U-Pb

823 ± 11 Ma

Pallister et al. (1987)

Ar/39Ar Ar/39Ar

847 ± 13 Ma Jabal Thurwah

Gabbro

U-Pb

870 ± 11 Ma

Pallister et al. (1988)

Al-Amar

Gabbro

U-Pb

698 Ma

Pallister et al. (1988)

Locality

Material, rock type

Method

Age (Ma)

References

Wadi Ghadir

Zircon, plagiograinte

207

746 ± 19

Kröner et al. (1992)

Abu Swayel

Zircon, gabbro

207

Eastern Desert

Pb/206Pb Pb/206Pb

729 ± 17

Zircon, diorite

207

Pb

736 ± 11

Zircon, layered gabbro

207

Pb/206Pb

741 ± 21

Whole rock, Gabbro

Sm–Nd

720 ± 9

CPX, PL

Sm–Nd

770 ± 52

G. Harga-zarga-Heinai

Whole rock, basalt

Sm–Nd

834 ± 19

Meatiq

Zircon, ophiolite cover

207

804 ± 7

Wadi Allaqi

Zircon, gabbro

206

730 ± 6

Ali et al. (2010)

Jabal Shilman

metagabbro

Pb–Pb

729 ± 17

Kröner et al. (1992)

metadiorite

Pb–Pb

736 ± 11

G. Gerf

2010), whereas the younger accretion-related ages from the ophiolites overlap the island arc stage. There have been a number of more focused studies of the ages of particular ophiolite sequences in the Eastern Desert

Pb/

206

Pb/206Pb Pb/238U

Zimmer et al. (1995)

Loizenbauer et al. (2001)

of Egypt. Kröner et al. (1992) obtained single zircon Pb–Pb evaporation ages of 729 ± 17 Ma and 736 ± 11 Ma for metagabbro and metadiorite associated with serpentinites of the Jabal Shilman area. The plagiogranite of the ophiolite

306

sequence at W. Ghadir yielded a 207Pb/206Pb single zircon evaporation age of 746 ± 19 Ma (Kröner et al. 1992). This age is identical to a 207Pb/206Pb age of 741 ± 21 Ma for zircons from the layered gabbro unit of the ophiolite sequence at Gebel Gerf, in the South Eastern Desert (Kröner et al. 1992); however, Zimmer et al. (1995) obtained a mean Sm/Nd age of 720 ± 9 Ma from 13 gabbro samples from the G. Gerf ophiolite. Interestingly, Sm/Nd isochrons constructed from clinopyroxene and plagioclase separates from these same gabbros yielded a rather higher mean age (770 ± 52 Ma). Zimmer et al. (1995) adopted a two-stage model for Pb isotope evolution and calculated the time of ophiolite formation as 771 ± 53 Ma. The reported Sm/Nd age of 834 ± 19 Ma (Zimmer et al. 1995) for basalt/diabase sheeted dikes and pillowed basalts of the G. Harga-Heinai ophiolite sequence in the South Eastern Desert must be taken with some caution as these rocks represent the topmost part of the sequence. The ophiolite cover of the Gebel Meatiq core complex yielded a 207 Pb/206Pb single zircon evaporation age of 804 ± 7 Ma (Loizenbauer et al. 2001). The ophiolitic layered gabbro of Wadi Abu Fas, a part of the Wadi Allaqi ophiolite in the South Eastern Desert, gave a concordia age of 730 ± 6 Ma (Ali et al. 2010). The emplacement of this ophiolitic gabbro is also constrained by two younger intrusive bodies: a cross-cutting gabbro yielded a concordia age of 697 ± 5 Ma and a cross-cutting quartz-diorite yielded a concordia age of 709 ± 4 Ma (Ali et al. 2010). There are a number of geochronological constraints that define approximate ages for the Neoproterozoic ophiolitic rocks on the Saudi Arabian side of the ANS. Pallister et al. (1987) report U–Pb model ages of 823 ± 11 Ma and 847 ± 14 Ma for zircons separated from plagiogranite dikes that cut the serpentinite south of Bir Tuluhah. These two ages provide a minimum age for the ultramafic protoliths of the serpentinite and taken to establish a lower limit to the age of the ophiolite. The ages of the Bir Umq ophiolite and associated intrusions range between 838 and 764 Ma (Dunlop et al. 1986; Pallister et al. 1988). A U-Pb zircon model age of 838 ± 10 Ma was obtained from diorite in the ophiolite close to the Bir Umq fault (Pallister et al. 1988). Ophiolitic metagabbro of Bir Umq yield a composite Sm– Nd isochron age of 828 ± 47 Ma, while metabasalt yields a Rb-Sr whole-rock isochron of 831 ± 47 Ma (Dunlop et al. 1986). Plagiogranite cutting serpentinized and carbonated peridotite of Bir Umq gave single-point zircon model ages of 764 ± 3 Ma and 782 ± 5 Ma that are interpreted as minimum ages for ophiolite emplacement (Pallister et al. 1988). Dilek and Ahmed (2003) presented a systematic overview of the ages and occurrence of Proterozoic ophiolites within the Arabian Shield in Saudi Arabia. They concluded that the ophiolites within the Yanbu and Bir Umq suture zones in the west are among the oldest (870–740 Ma) in the Shield; these

M. G. Shahien et al.

are also the ophiolites with chemical characteristics most like modern fore-arc oceanic crust. The ophiolites of the Hulayfah-Ruwah suture zone in the central Arabian Shield are coeval with and/or slightly younger (*843–821 Ma) than the ophiolites in the west and probably developed in a rifted ensimatic arc system that evolved as a volcanic archipelago near the Afif continental plate. Younger ophiolites (c. 694 Ma) of the Halaban and Al-Amar suture zones in the eastern Arabian Shield were incorporated into a subduction-accretion complex that evolved at the Andean-type active margin along the eastern edge of the Afif continental plate. The Halaban suture zone ophiolites represent fore-arc oceanic crust, whereas the Al-Amar suture zone ophiolites are scraped-off fragments of Mozambique ocean floor, seamounts, and/or ocean island(s). The ophiolites of the Nabitah-Hamdah fault zone within the Asir terrane are the youngest (*627 Ma) in the Shield, post-collisional in origin, display mid-ocean ridge basalt chemical affinity, and developed in an intracontinental para-rift zone.

12.5

Protolith and Geodynamic Setting

Ophiolite suites around the world and through geologic time, while sharing many basic defining features, also reveal significant variations in their internal structure, geochemical characteristics, emplacement mechanisms, and tectonic environments (Dilek et al. 2008; Dilek and Furnes 2014). They represent remnants of ancient oceanic lithosphere created in a variety of tectonic settings, including nearly all parts of the Wilson cycle from rift–drift and seafloor spreading stages to subduction initiation and terminal closure (Dilek and Furnes 2011, 2014; Furnes et al. 2014). In the end, their preservation requires that they are tectonically emplaced onto continental margins, usually upon the closing of an ocean basin. The ophiolitic rocks of the ANS in particular have been studied for a long time, because (relying on uniformitarian analogy to younger ophiolites and samples from active plate boundaries) one can use their characteristics to reconstruct the geodynamic evolution of the Pan-African belt. However, available information about the ANS ophiolites varies in quality and quantity; only a few of the sequences are mapped and characterized in sufficient detail to fully determine their tectonic setting and so considerable controversy remains. Published studies over the years have assigned the ANS ophiolites to a diversity of tectonic settings, including both open mid-ocean ridge and supra-subduction zone settings (e.g., Zimmer et al. 1995; Dilek and Ahmed 2003; Ahmed and Hariri 2008; Ahmed and Habtoor 2015; Gahlan et al. 2015, 2018; Obeid et al. 2016; Azer et al. 2019; Abuamarah 2019a, b and many others). Along with the growing

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Neoproterozoic Ophiolites of the Arabian-Nubian Shield

realization that most of the world’s recognized ophiolites are associated with subduction, though, more recent studies have shifted toward a consensus that the ANS ophiolites were generated in supra-subduction zone tectonic settings (e.g., Nassief et al. 1984; Pallister et al. 1988; Azer and stern 2007; Abd El-Rahman et al. 2009; Abuamarah et al. 2020; Gahlan et al. 2020a, b and many others). The abundance of immature and volcaniclastic sediments deposited on top of the ophiolites suggests formation at an intra-oceanic convergent margin, but two distinct kinds of oceanic lithosphere are created at such margins; these are most conveniently divided into fore-arc and back-arc settings. We note that there is some confusion about the concept of fore-arc spreading and want to be clear how the term is used in the ophiolite community. In mature subduction systems with well-defined volcanic fronts, the area between the volcanic front and the trench—the fore-arc—is typically amagmatic. However, during subduction initiation, the area that will later develop into such a fore-arc is in fact generally a major locus of seafloor spreading and an ophiolite nursery (Stern et al. 2012). Debate continues about the assignment of the ANS ophiolites to episodes of back-arc basin opening or to fore-arc spreading connected with subduction initiation. This issue is critical to the use of the ophiolites to define the tectonic history of the shield because fore-arc ophiolites mark episodes when new subduction zones form (Stern 2004), often associated with major plate reorganizations. In contrast, back-arc basins can form at any time in the evolution of a convergent plate margin. For a number of years, efforts to assign the tectonic setting of the ANS ophiolites focused mostly on the trace element compositions of the lavas and most such studies favored a back-arc setting (e.g., Ledru and Auge 1984; Nassief et al. 1984; Ahmed and Hariri 2008; El-Sayed et al. 1999; Nehlig et al. 2002; Farahat et al. 2004; Abd El-Rahman et al. 2009). These studies rarely, if ever, considered evidence from the associated serpentinized peridotites. However, interpreting the tectonic setting of Neoproterozoic ophiolitic rocks on the basis of major and trace element compositions of metavolcanic rocks encounters considerable difficulties due to the effects of fractional crystallization and alteration. Even when these problems are minimized, it can be very difficult to distinguish fore-arc and back-arc lavas on the basis of chemical compositions (Azer and Stern 2007). Several authors recognized that the geochemical characteristics of the ANS pillow lavas are transitional between those of island arcs and MORB and used this observation to support the back-arc formation scenario for the ANS ophiolites. However, the transitional compositional features of ophiolitic lavas simply reflect the hydrous nature of the magmatism and the effect of adding components from the subduction zone into the overlying mantle wedge, which can occur in both the fore-arc and the back-arc

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(Azer and Stern 2007). More recently, a number of studies have begun to focus on evidence from the mantle sections of the ANS ophiolites, arguing that the signatures recorded by chemical compositions of fresh relics of mafic minerals offer an unequivocal basis for distinguishing fore-arc and back-arc settings. These studies have uniformly, in recent years, favored fore-arc settings for all of the studied ANS ophiolites (Azer and Stern 2007; Khalil and Azer 2007; Ali et al. 2020b; Abdel-Karim et al. 2020a; Abuamarah et al. 2020; Gahlan et al. 2020a, b and many others). This assignment is supported by the observation of magmas with boninitic affinities in the ANS (Wolde et al. 1993; Yibas et al. 2003; Katz et al. 2004; Abdel Aal et al. 2003; Saleh 2006; Teklay 2006); in the Phanerozoic, boninites appear to be restricted to and absolutely diagnostic of fore-arc settings (e.g., Murton 1989; Johnson and Fryer 1990; Bédard 1999; Beccaluva et al. 2004). Even in highly serpentinized peridotite, the chemistry of relict minerals (olivine, pyroxene, and chromian spinel) offers a useful tool to distinguish the tectonic environment of the ultramafic protolith (e.g., Dick and Bullen 1984; Barnes and Roeder 2001; Ohara et al. 2002; Arif and Jan 2006; Uysal et al. 2012; Khalil et al. 2014; Obeid et al. 2016). In many cases, spinel-group minerals are the only phases that retain most of their original chemistry in highly serpentinized peridotites. Hence, particularly in ultramafic rocks so penetratively serpentinized that no relics of primary silicate minerals remain, the chemical compositions of the cores of unaltered accessory chromite have been widely recognized as a potentially important petrogenetic indicator (e.g., Dick and Bullen 1984; Arai 1992, 1994; Zhou et al. 1996; Barnes and Roeder 2001; Sobolev and Logvinova 2005; Arif and Jan 2006). Spinel from mid-ocean ridge and back-arc basin peridotites generally has Cr# (molar 100  Cr/[Cr + Al]) less than 50, whereas spinel in fore-arc harzburgites generally has higher Cr# (60–80) and spinel from boninites typically has Cr# of 70–90 (Barnes and Roeder 2001; Ohara et al. 2002). The ranges and averages of Cr# and Mg# in chromian spinel from various serpentinized peridotite localities in the ANS are summarized in Table 12.3. Spinel affected by alteration and extreme enrichment or depletion in Cr2O3, Al2O3, and FeO(t) are excluded. From this table, it is clear that the Cr# of spinel in the ANS ophiolitic peridotites are mostly > 60 and similar to those of modern fore-arc peridotites. The data underlying the average in Table 12.3 are plotted individually in Cr# versus Mg# space in Fig. 12.6, where is can be seen that the data cluster above Cr# = 60 and essentially overlap the compositional range defined by modern fore-arc peridotites. Fresh relics of primary olivine, even in highly serpentinized peridotites, can provide insights into the tectonic setting in which the ultramafic protoliths formed, including

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Table 12.3 Microprobe analyses of fresh accessory chromites from serpentinites in Eastern Desert of Egypt and Saudi Arabia Eastern Desert of Egypt Area El-Rubshi

Fawakhir

Cr#

Average

Mg#

Average

No. of spots

0.73–0.85

0.81

0.17–0.43

0.32

7

Khudeir (1995)

0.77–0.78

0.78

0.36–0.40

0.38

5

Khalil (2000)

0.53

0.53

0.6

0.6

1

Ahmed et al. (2001)

0.75

0.75

0.25–0.28

0.26

2

Khudeir (1995)

0.66–0.80

0.75

0.32–0.50

0.34

25

0.71–0.87

0.76

0.46–0.69

0.58

9

Wadi Sodmein

0.59

0.59

0.3

0.3

1

Gabal El-Degheimi

0.62–0.79

0.7

0.25–0.66

0.47

16

Bir Al-Edied

0.70–0.80

0.74

0.50–0.72

0.67

8

Um Saneyat

0.47–0.79

0.62

0.14–0.66

0.42

12

Muweilha

Um Khariga

Wadi Arayis

References

0.78–0.84

0.81

0.14–0.23

0.18

2

0.47–0.57

0.5

0.45–0.63

0.52

21

0.64–0.85

0.69

0.26–0.59

0.42

14

0.64–0.66

0.65

0.58

0.58

2

0.88–0.91

0.89

0.29–0.45

0.37

15

0.64–0.71

0.68

0.52–0.55

0.53

2

0.65–0.78

0.73

0.16–0.70

0.48

8

0.65–0.67

0.66

0.46–0.63

0.58

23

Hamdy et al. (2013b) Abdel-Karim and El-Shafei (2018) El Bahariya (2003) Azer (2014) Azer and Khalil (2005) Khalil (2000) El Bahariya (2003) El Dien et al. (2016) Khalil (2000) El Bahariya (2003) Abdel-Karim et al. (2020a) El Bahariya (2003) Khalil and Azer (2007) Ali et al. (2020a)

0.68–0.77

0.72

0.23–0.45

0.38

19

Hamdy and Lebda (2007)

0.60–0.79

0.68

0.23–0.58

0.47

26

Obeid et al. (2016)

Gerf

0.84–0.87

0.85

0.56–0.68

0.63

4

Abdel-Karim et al. (2016)

G. Kurbiyay

0.81–0.86

0.83

0.26–0.35

0.31

5

Abdel-Karim et al. (2020a)

Wadi Ambaut

0.62–0.69

0.66

0.43–0.55

0.48

18

Abdel-Karim et al. (2020b)

Samadi

0.69–72

0.7

0.51–0.55

0.53

27

Ali et al. (2020a)

Um Khasila

0.47–0.72

0.65

0.5–0.7

0.56

19

Um Salem

0.66–0.8

0.72

0.16–0.34

0.26

35

Abu Siayil

0.84–0.86

0.85

0.44–0.47

0.46

9

Ali et al. (2020b)

W. Hammariya

0.83–0.85

0.85

0.61–0.66

0.63

4

Abdel-Karim et al. (2014)

G. Umm Halham

0.61–0.86

0.72

0.49–0.81

0.68

11

G. Garf

0.84–0.87

0.85

0.56–0.68

0.63

4

Abdel-Karim and El-Shafei (2018)

Wadi Atalla

0.48–0.75

0.62

0.18–0.69

0.53

18

Hamdy and El Dien (2017)

0.57–0.71

0.66

0.44–0.62

0.51

28

El Dien et al. (2016)

Abdel-Karim et al. (2014)

Gebel El-Maiyit

0.67–0.80

0.75

0.54–0.70

0.6

24

El Dien et al. (2016)

Wadi Dober

0.56–0.74

0.67

0.12–0.21

0.18

12

Hamdy et al. (2018)

Wadi Alam

0.55–0.69

0.61

0.37–0.54

0.51

15

0.55–0.71

0.62

0.48–0.62

0.56

138

Hamdy and El Dien (2017) El Dien et al. (2019)

Gebel Malo Grim

0.67–0.81

0.77

0.2–0.47

0.276

27

Hamdy and Lebda (2007)

Sol Hamed

0.68–0.81

0.759

0.2–0.5

0.3

20

Abu-Alam and Hamdy (2014)

No. of spots

References

Saudi Arabia Area

Cr#

Average

Mg#

Average

Jabal Ess

0.59–0.77

0.68

0.30–0.63

0.53

40

Gahlan et al. (2020a)

Al-Wask

0.59–0.68

0.64

0.31–0.60

0.48

30

Gahlan et al. (2020b)

Bir Umq

0.60–0.72

0.68

0.23–0.65

0.48

45

Abuamarah et al. (2019a)

Bir Tuluhah

0.58–0.72

0.65

0.54–0.73

0.65

50

Abuamarah et al. (2020)

Halaban

0.60–0.75

0.66

0.35–0.58

0.48

47

Abuamarah et al. (2019b)

Wadi Al Hwanet

0.50–0.62

0.58

0.49–0.60

0.55

33

Ali et al. (2020b)

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Neoproterozoic Ophiolites of the Arabian-Nubian Shield

the distinction between supra-subduction and mid-ocean ridge settings (Parkinson and Pearce 1998). Magma genesis in different tectonic settings leads to different extents of melting and residual olivines become increasingly magnesian as melting progresses; the larger the extent of melting, the higher the Mg# of the residual olivine. Hence, the uniformly Mg-rich nature of olivine relics in the ANS serpentinites (Fo > 89) (Ledru and Auge 1984; Nassief et al. 1984; Khudeir 1995; Stern et al. 2004; Khalil and Azer 2007; Abuamarah 2019a, b; Abuamarah et al. 2020; Gahlan et al. 2020a, b) indicates extensive melt extraction. This is a common feature of fore-arc peridotites (Arai 1994; Coish and Gardner 2004). In the fore-arc environment, the mantle has often been previously somewhat depleted; it is then flushed with a large quantity of water that reduces its high solidus temperature and promotes still more melting and depletion. By the time such peridotites are emplaced at the top of the mantle section, they are significantly more depleted than peridotites that ascend and melt under comparatively dry conditions (Pearce et al. 2000), suggesting that they are residual after extensive melting. The compositions of coexisting olivine and spinel, taken together, are indeed consistent with the trend of partial melt extraction and plot at the highly depleted end of the olivine-spinel mantle array, in the field of fore-arc peridotites (Fig. 12.7a). Pyroxene, when present in a residual peridotite and preserved through alteration, can give more clues about petrogenesis. For example, there is a systematic decrease in Al content in residual mantle pyroxene with increasing degree of partial melting (e.g., Gasparik 1987; Dick and Natland 1996). There are many published chemical analyses of fresh relics of pyroxenes (orthopyroxene and clinopyroxene) in the serpentinized peridotites of the ANS (e.g., Abdel Aal et al. 2003; Khalil et al. 2014; Obeid et al. 2016; Abuamarah 2019a, b; Abdel-Karim et al. 2020a; Ali et al. 2020b; Gahlan et al. 2020a; and many others). Orthopyroxene in these rocks is mainly enstatite; clinopyroxene may be diopside or augite, according to the orthopyroxene nomenclature of Morimoto et al. (1988). Both orthopyroxene and clinopyroxene relics in ANS serpentinites have high Mg# (>0.9) and low TiO2, Al2O3, and Cr2O3. Mg# of orthopyroxene is very similar to Mg# of coexisting olivine (Abuamarah 2019a, b; Gahlan et al. 2020a, b), suggesting that these two phases equilibrated under high-temperature mantle conditions (Uysal et al. 2016). The high Mg# of pyroxenes and low contents of Al2O3 and Cr2O3 are again entirely consistent with observations from modern fore-arc peridotites. In fact, the available clinopyroxene compositions in the ANS ultramafic rocks are similar to clinopyroxene from boninite (Fig. 12.7 b), a distinctive rock type characteristic of intra-oceanic fore-arcs (e.g., Bédard 1999; Beccaluva et al. 2004; Dilek et al. 2008; Dilek and Thy 2009).

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A number of geodynamic models have been proposed in recent years that attempt to explain the preponderance of fore-arc ophiolites in the ANS and in the context of the overall evolution of the Pan-African orogeny (e.g., Azer and Stern 2007; Abu-Alam and Hamdy 2014; El Dien et al. 2016; Khedr and Arai 2016; Ali et al. 2020b; Abdel-Karim et al. 2020a). In general, the ophiolite ages within the ANS indicate an eastward progression of continental growth and ophiolite obduction in the East African Orogen through late Neoproterozoic time (Dilek and Ahmed 2003). Azer and Stern (2007) proposed a model for the fore-arc SSZ ophiolites of the northwest corner of the ANS, as summarized in Fig. 12.8. As fore-arc spreading is a short-lived process during the initiation of a new subduction zone, this model made it clear that the Neoproterozoic history of the ANS, especially in the Eastern Desert of Egypt, consisted of a series of subduction initiation events. In many cases, the ophiolites are then thrust over arc-related volcano-sedimentary successions that may preserve some evidence of the mature stage of subduction. Abdel-Karim et al. (2020a) proposed a new model (Fig. 12.9) for the fore-arc SSZ ophiolites of Egypt. In the first stage, during the opening of the Mozambique Ocean between East and West Gondwana, peridotites were produced as the residues of (relatively) anhydrous mid-ocean ridge-type melting (Fig. 12.9a). Subsequently, as the margins of the Mozambique Ocean foundered and underwent subduction initiation, the previously formed depleted peridotites were drawn into a proto-fore-arc spreading center (Fig. 12.9b). Channelization of fluid flow at this stage may leave some dry and infertile peridotites that preserve the original geochemical fingerprint of mid-ocean ridge melt depletion (Uysal et al. 2014). As the arc system matures toward true subduction, the new slab liberates a high flux of fluid (Fig. 12.9c) that drives high-degree partial melting of depleted peridotite, producing boninitic melts from which the high-Cr chromitites noted in some ANS ophiolites could be precipitated.

12.6

Alteration and Metamorphism

The ultramafic members of the ANS ophiolites are generally highly altered, and the alteration processes are interesting as tracers of fluid sources and flow paths and because of their relationship to mineralization. However, it is generally difficult to determine when and where this alteration occurred. Alteration and metamorphism of the ANS ultramafic sections may have occurred on the ocean floor, below the oceanic crust, during obduction and tectonic emplacement, or even upon recent uplift and exposure. Alteration and metamorphism have largely changed the mineralogical make

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Fig. 12.6 Cr# versus Mg# diagram for chromian spinels (after Stern et al. 2004). The field boundaries are from Dick and Bullen (1984), Bloomer et al. (1995), and Ohara et al. (2002). a Fresh accessory spinels in the Egyptian serpentinized peridotites (the data are from: Azer and Stern 2007; Azer et al. 2013, 2019; Khalil et al. 2014; Azer 2014; Gahlan et al. 2015, 2018; Obeid et al. 2016; Boskabadi et al. 2017; Abdel-Karim et al. 2020b; Ali et al. 2020a, 2020b; the figure is adapted from Azer and Asimow, in press), and b fresh accessory spinels in serpentinized peridotites from Saudi Arabia (the data are from: Chevremont and Johan 1982; Le Metour et al. 1982; Abuamarah 2019a, b; Abuamarah et al. 2020; Gahlan et al. 2020a, b)

up of these rocks and may have changed their bulk chemistry as well, but skilled petrographers can still recognize many original igneous textures. Furthermore, the agreement among classifications based on relict mineral chemistry, on inferred original mineralogy, and on measured bulk chemistry indicate that some aspects of the bulk chemical

composition are robustly preserved. Metamorphism and subsequent alteration of the ANS ophiolites are manifested by several distinct processes and their characteristic products. Their mantle sections experienced serpentinization, carbonatization, and listvenitization. Note that metamorphism and each of these alteration processes may or may not

12

Neoproterozoic Ophiolites of the Arabian-Nubian Shield

311

Fig. 12.7 a Cr# of spinel versus Fo content of coexisting olivine from serpentinized peridotites of the ANS (Arai 1992). PM: Primitive mantle, OSMA: Olivine-spinel mantle array (Arai 1994). b TiO2–Na2O–SiO2/100 diagram for fresh relics of clinopyroxene in serpentinized peridotites of the ANS (Beccaluva et al. 1989). WOPB = within-ocean plate basalts; MORB = mid-ocean ridge basalts; IAT = island arc tholeiites; BON + BAA = boninites + basaltic andesites and andesites from intra-oceanic fore-arcs

be contemporaneous and may be independent or coupled. The crustal section records albitization, uralitization, rodingitization, and chloritization. The mantle section and the volcanic carapace are often more extensively altered than the intervening gabbros (Gahlan et al. 2015; Obeid et al. 2016). The most notable alteration process, affecting all the ultramafic rocks in the ANS ophiolites, is serpentinization, a continuous process of replacement of anhydrous primary silicate minerals by hydrous phyllosilicates (e.g., Evans 1977;

O’Hanley 1996; Douville et al. 2002; Allen and Seyfried 2003; Mével 2003). Serpentinization of the ophiolitic ultramafic rocks is pervasive and is often accompanied by a first cycle of carbonatization reflecting the activity of CO2 in the serpentinizing fluids (e.g., Stern and Gwinn 1990; Boskabadi et al. 2017; Azer et al. 2019). Infiltration of further CO2bearing fluids during deformation is shown by the localization of further alteration of serpentinite along shear zones and fault planes. Such localized features include recrystallization of the

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Fig. 12.8 Cartoon showing the tectonic setting of an Egyptian ophiolite, in the northwestern part of the ANS, in the fore-arc environment above a subduction zone (after Azer and Stern 2007)

serpentine minerals and formation of listvenite, talc-carbonate rocks, and magnesite deposits (Gahlan et al. 2018). Certainly, the prevalence of carbonate alteration of ANS ophiolitic ultramafics suggests a tremendous flux of CO2 in mantle-derived fluids, probably during middle and late Neoproterozoic time (Newton and Stern 1990; Stern and Gwinn 1990). The origins and compositions of CO2-bearing fluids that interacted with the ANS ophiolites are controversial. Suggestions include (i) mantle-derived CO2-bearing fluids (e.g., Boskabadi et al. 2017; Hamdy and El Dien 2017a), (ii) both H2O-rich and CO2-rich fluids released from various layers of a subducting slab (e.g., Hamdy et al. 2013a); (iii) mixtures of mantle-derived fluids and remobilized sedimentary carbonate (Stern and Gwinn 1990); and (iv) Paleozoic to recent hydrothermal fluids infiltrating during and after exhumation (˃100 °C) (e.g., Hamdy and Lebda 2007). Just as the fluid source for carbonatization, listvenitization, and rodingization are controversial, so is the timing of these processes. Gahlan et al. (2018) distinguished two episodes of metasomatic stages (listvenitization) associated with the Eastern Desert ophiolites. The first stage of listvenitization was associated with the formation of the oceanic lithosphere and the second stage occurred during obduction of the ophiolitic sequences. Likewise, Azer et al. (2019) distinguished two stages of carbonation in the ANS ophiolites. The first stage formed magnesite masses during deep-seated metasomatism and serpentinization, whereas the

second stage emplaced carbonate veins after serpentinization, during obduction of the ophiolite. Serpentinization can be isochemical, which requires a significant volume increase and causes fracturing, or allochemical, which can in principle proceed at constant volume as fluids transport Mg2+, Fe2+, Ca2+, and Si2+ ions out of the system. Johannes (1969, 1970) provided benchmark experiments of equilibria in the system MgO–SiO2–H2O–CO2 from 0.2 to 1.0 GPa pressure that demonstrate that serpentine coexists only with CO2-poor fluid phases. Serpentine minerals may form by hydration of primary mafic phases or by recrystallization of earlier serpentine phases during prograde metamorphism, or in some cases by retrograde hydrothermal alteration of ultramafic rocks metamorphosed under dry conditions (Deer et al. 1992). Antigorite is the most common prograde reaction product, while lizardite is the most common retrograde reaction product. Published studies of ANS ophiolites indicate that the serpentinites are dominated by antigorite with subordinate lizardite and chrysotile. The simplest interpretation would be that the protoliths were serpentinized during heating and burial (Moody 1976; Deer et al. 1992). The alternative model of retrograde formation of chrysotile and lizardite followed by a second cycle of prograde recrystallization to antigorite (Azer and Khalil 2005) is more complex but difficult to rule out. The dominance of antigorite over chrysotile and lizardite in the ANS serpentinites could be explained either by elevated temperature,

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Neoproterozoic Ophiolites of the Arabian-Nubian Shield

313

Fig. 12.9 Schematic illustration showing a opening of Mozambique Ocean between East and West Gondwana; b subduction initiation or incipient arc stage, in which partial melting occurred in the mantle wedge and proto-fore-arc spreading starts to form over the subduction zone; c the true subduction or mature-arc stage, associated with high fluxes of slab-derived fluids and high-degree partial melting of depleted mantle to produce boninitic melts, from which the high-Cr chromitites were precipitated (Abdel-Karim et al. 2020a)

above 400−600 °C (Evans 2010), or by elevated pressure; increases in both variables favor antigorite. The serpentinization of peridotites can occur in both low-temperature (  250 °C) and high-temperature (  250 °

C) environments (Evans 2010). At low temperature, olivine relics tend to retain their original composition, while high-temperature serpentinization facilitates interdiffusion of Mg and Fe, resulting in low-Mg# olivine relics coexisting with

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serpentine. Furthermore, experimental studies have demonstrated that pyroxenes react faster than olivine at temperature above 250–300 °C, but olivine reacts faster than pyroxene at temperature < 250 °C (e.g., Martin and Fyfe 1970; Janecky and Seyfried 1986; Allen and Seyfried 2003). In the ANS, olivine appears to have reacted most readily, followed by orthopyroxene and clinopyroxene. Also, fresh olivine relics have high Mg#, typically greater than 89. Both features imply low-temperature hydration. These observations, therefore, suggest that antigorite was largely stabilized by elevated pressure. Upon prograde metamorphism, the reverse of serpentinization, i.e., dehydration, can occur, producing metamorphic olivine (e.g., Arai 1975; Frost 1991; Evans 1977; Nozaka 2003, 2010; Evans 2010). Such metamorphic olivine typically lacks chemical zonation or evidence of ductile deformation such as undulatory extinction (e.g., Mercier and Nicolas 1975, Khalil and Azer, 2007). Metamorphic olivine has been reported in the serpentinized peridotites from the ANS ophiolites (e.g., Gahlan 2006; Khalil and Azer 2007; Gahlan et al. 2020a; Ali et al. 2020a). In fact, two types of metamorphic olivine are observed in the ANS ophiolites: olivine with low MgO contents (Fo < 90; Azer and Khalil, 2005) and olivine with high MgO (Fo > 90; Gahlan et al. 2020a, Ali et al. 2020a). The low-Mg# metamorphic olivine is typically associated with and attributed to the thermal influence of nearby granitoid intrusions (e.g., Khalil and Azer 2007; Gahlan and Arai 2009; Ahmed et al. 2012), whereas high-Mg# metamorphic olivine is typically attributed to low-pressure, high-temperature regional metamorphism (Gahlan et al. 2015, 2020a). Rodingite is a massive light-colored rock composed of Ca-rich minerals. It is produced by infiltration of Ca-bearing solutions into mafic rocks adjacent to serpentinized ultramafic bodies (Best 2003). Determination of the physical and chemical conditions of rodingite formation, concomitant with serpentinization, can provide detailed information on the effects of ancient seafloor hydrothermal processes and on the tectonic history of an ophiolite. Only a few studies have published data about rodingites (Takla et al. 1992; Abdel-Karim 2000; Surour 2019). Although the ANS ophiolites have been intensively studied over the last two decades, the associated rodingites have not yet received much attention. Under the influence of post-magmatic and/or metamorphic processes, primary chromian spinel starts to develop alteration products such as ferritchromite and Cr-magnetite (e.g., Farahat 2008; Saumur and Hattori 2013; Khalil et al. 2014; Gahlan et al. 2015; Obeid et al. 2016; Abdel-Karim and El-Shafei 2018; El Dien et al. 2019; Ali et al. 2020a; Abuamarah 2019a, b; Gahlan et al. 2020a, b; Abuamarah et al. 2020 and many others). The ferritchromites in serpentinites of the ANS are enriched in total iron and strongly depleted in Al2O3 and MgO, reflecting loss in Al2O3 and

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Cr2O3 and increase in both FeO and Fe2O3 due to alteration and metamorphism. Low-temperature serpentinization produces ferritchromite zones with Cr-magnetite outer zones, as a result of fluid ingress along cracks and around the grain boundaries of Cr-spinel. These two secondary phases are usually attributed to the effects of low- to medium-grade metamorphism, up to lower-amphibolite facies (Thalhammer et al. 1990; McElduff and Stumpfl 1991). In the ANS ophiolites, the alteration of chromian spinel to ferritchromite may have started during the late magmatic stage, but it is thought to have mainly occurred during serpentinization and tectonism (Khudeir et al. 1992; Khalil and Azer 2007). The very low Fe3+ contents observed in fresh chromian spinel in ANS serpentinites (Azer 2014, Abuamarah et al. 2020) indicate relatively low oxygen fugacity conditions in their primary sources (Murck and Campbell 1986). However, high Fe3+ in the ferritchromite and Cr-magnetite rims suggest oxidizing conditions during metamorphism (Anzil et al. 2012). Oxidizing conditions favor the reaction of chromian spinel with serpentine to produce chlorite, ferritchromite, and Cr-magnetite (Mellini et al. 2005; González-Jiménez et al. 2009). Therefore, the development of ferritchromite mantles and outer Cr-magnetite rims around chromian spinel cores (sometimes associated with Cr-rich chlorite or “kämmererite”) indicates their formation during prograde alteration under oxidizing conditions (González-Jiménez et al. 2009). This alteration should have taken place during lower-amphibolite facies metamorphism (Suita and Strieder 1996). The minimum temperature of formation of ferritchromite is *500 °C (Mellini et al. 2005). Kämmererite is rarely observed in some serpentinites of the ANS; it is not typically an major product of serpentinization. Rather, its presence in some samples (Azer and Stern 2007; Gahlan et al. 2018, 2020b; Abuamarah 2019a) likely reflects the replacement of chromian spinel during later alteration or regional metamorphism. This is supported both by the significant Cr content of kämmererite and by its close petrographic association with primary relics of Cr-spinel. During alteration of chromian spinel, most Cr and Fe enter into ferritchromite, whereas Al and Mg are released to the surrounding silicate minerals. The excess Al reacts with serpentine minerals to produce chlorite, and buffering to high-Cr activity by adjacent ferritchromite leads to kämmererite (Azer and Stern 2007). The occurrence of kämmererite around primary relics of chromian spinel reflects a further, lower-temperature episode of replacement of chromian spinel, below 300 °C according to the chlorite thermometry results (Gahlan et al. 2018; Abuamarah 2019a). The metamorphic grade implied by the mineral assemblages in serpentinized ultramafic rocks of the ANS range from greenschist facies (olivine–opx–serpentine–ferritchromitemagnetite) to amphibolite facies (olivine–opx–anthophyllite– talc–tremolite–antigorite–ferritchromite–magnetite) (e.g.,

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El-Sayed et al. 1999; Azer and Khalil 2005; Farahat 2008; Khedr and Arai 2013; Gahlan et al. 2015, 2020a, b; Obeid et al. 2016; Abuamarah et al. 2020). The minimum temperature of ferritchromite formation has been reported as *500 °C, which would make it an indicator mineral for loweramphibolite facies metamorphism (Suita and Strieder 1996; Mellini et al. 2005). Yet Evans and Frost (1975) and Suita and Streider (1996) report the transformation of chromian spinel to ferritchromite in greenschist facies rocks. The observation of sharp compositional boundaries between ferritchromite rims and fresh chrome spinel cores (e.g., Gahlan et al. 2015, 2018, 2020a, b; Ali et al. 2020a; Abdel-Karim et al. 2020a, b; Abuamarah et al. 2020 and many others) likely indicates upper greenschist to lower-amphibolite facies metamorphism (Evans and Frost 1975; Frost 1991; Suita and Streider 1996; Barnes and Roeder 2001; Arai et al. 2006). Farahat (2008) attributed the formation of chromian spinel cores followed by ferritchromite and Cr-magnetite rims to transitional greenschistamphibolite to lower-amphibolite facies metamorphism.

12.7

Examples of ANS Ophiolites

12.7.1 Saudi Arabian Ophiolites: Jabal Al-Wask The Arabian Shield, the eastern half of the ANS, features numerous Neoproterozoic ophiolite suites, including Jabal Ess, Jabal Al-Wask, Jabal Tharwah, Bir Umq, Bir Tuluhah, Halaban, Jabal al Uwayjah, Jabal Tays, and many others. They outcrop along large faults and major shear zones and demarcate the sutures between terranes of different lithostratigraphy (Fig. 12.10). We illustrate the typical features of these complexes by discussing in detail one Arabian Shield ophiolite sequence, at Jabal Al-Wask. The Al-Wask ophiolite is one of the largest and best-preserved examples of a complete ophiolite sequence in the Arabian shield (e.g., Al-Shanti 1982; Pallister et al. 1988; Johnson et al. 2004; Gahlan et al. 2020b). It outcrops in the central part of the Yanbu suture, a NE-oriented belt that extends from the northwestern Arabian Shield into northeast Africa (e.g., Abdelsalam and Stern 1996; Johnson and Woldehaimanot 2003; Ali et al. 2010). It lies between latitudes 24° 45′ and 25° 30′ N and longitudes 37° 30′ and 38° 15′ E (Fig. 12.11). The Al-Wask ophiolite has undergone multiple phases of alteration, deformation, and metamorphism up to at least greenschist facies. Pervasive shearing during ophiolite emplacement and extensive serpentinization, carbonatization, and silicification resulted in the development of magnesite and listvenite along shear zones. The Al-Wask ophiolite was recently studied in detail by Gahlan et al. (2020b) and a brief summary of this study is provided in the following paragraphs. The Al-Wask ophiolite exhibits most of the lithologic sequence that defines a

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classic Penrose-type ophiolite suite, although sheeted dikes are absent (Fig. 12.12). It consists of a mantle section of serpentinized ultramafic rocks, structurally overlain by a crustal section of gabbros and pillow lavas. The whole ophiolite sequence is capped by pelagic sedimentary cover and tectonically emplaced over a metamorphosed island arc volcano-sedimentary succession. The Al-Wask ultramafic rocks are strongly serpentinized, deformed (mylonitized, in some cases), metamorphosed, and altered by carbonatization and silicification. Most contacts between the different rock units of Al-Wask ophiolite are structural (faults). An extensive mélange zone accompanies the sole thrust at the base of the sequence. The mélange consists mainly of metagabbro and metavolcanic blocks entrained in a serpentinite matrix. Serpentinized harzburgite is the dominant ultramafic rock type, while serpentinized dunite is also present. Chromitite, though scarce, forms small pods and lenticular layers, generally enveloped by serpentinite, with both massive and disseminated textures. Along shear zones, tectonic lineaments, and fault planes, ultramafic rocks are largely altered to talc-carbonates, magnesite, and listvenite. In places, the Neogene Harrat Lunayyir basalt flowed directly over exposed serpentinite outcrop. Magnesite bodies are commonly hosted by serpentinite, most often at contacts with country rocks and along regional faults. Magnesite forms masses, pockets, and snow-white veins, all with sharp but irregular boundaries. The massive magnesite is cryptocrystalline and composed essentially of anhedral fine-grained magnesite (90–95 vol.%) with minor serpentine minerals, chrome spinel, and iron oxides. Listvenite bodies of various shapes and sizes developed by alteration of ultramafic rocks, particularly along shear zones. Typically reddish-brown in color, fine- to medium-grained, and resistant to erosion, the Al-Wask listvenites form prominent geomorphic ridges. They consist essentially of carbonates and quartz. In the Jabal Al-Wask area, the listvenites are distinguished into carbonate listvenite, silica-carbonate listvenite, and silica listvenite. A normal thickness (up to 2 km; Stern et al. 2004) of metagabbro overlies a sequence of ultramafic cumulates across sheared or faulted contacts. Modally layered gabbros define a very thin zone, exposed only at few localities; the bulk of the crustal section is isotropic gabbro. The metagabbros grade upward from pyroxene-rich to hornblende-rich varieties, with minor diorites throughout. Spilites (interpreted as altered pillow lavas), sparse metasediments, and metatuffs overlie the metagabbro. The primary silicate minerals in the ultramafic mantle section have been largely replaced by serpentine minerals, except for a few relics of fresh olivine and chromian spinel. The serpentinites display bastite and mesh textures after orthopyroxene and olivine, reflecting exclusively harzburgite

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Fig. 12.10 Regional tectonic map of the Arabian Shield showing the distribution of suture zones hosting ophiolite sequences in Saudi Arabia (Nehlig et al. 2002)

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Fig. 12.11 Geologic map showing the outcrop area of the Al-Wask ophiolite and its country and cover rocks, modified after Hadley (1987) and Gahlan et al. (2020b)

and dunite protoliths. Whole-rock compositions have low SiO2, TiO2, CaO, and Al2O3 contents combined with high MgO, Mg#, Ni, Co, and Cr contents. Fresh relics of spinel have high Cr# and low TiO2 contents and primary olvine relics have high Fo contents. All these features point to highly refractory mantle protoliths, as commonly found in supra-subduction zone settings, especially fore-arcs.

12.7.2 Egyptian Ophiolites: Wadi Ghadir Ophiolites and ophiolitic mélanges are distinctive features of the central and southern parts of Eastern Desert of Egypt (Fig. 12.13), but they are completely lacking in the north

Eastern Desert. They occur as tectonized bodies and mélanges of pillow metabasalt, gabbro, and variably altered ultramafic rocks. The presence of ophiolites in Sinai is doubtful (Bentor 1985; El-Gaby et al. 1990); however, small bodies of serpentinite are known in south Sinai, a few tens of meters in length. These have been described as intruding (Madbouly 1991; Moussa 2002) or tectonically emplaced (Abu El-Enen and Makroum 2003) into migmatites and gneisses at Kabr El-Bonaya in the Wadi Kid area. On the other hand, Shimron (1980) and Abu El-Enen and Makroum (2003) inferred that the serpentinites of Kabr El-Bonaya are remnants of a disrupted ophiolite. Abdel Khalek et al. (1994) also described an ophiolitic sequence in the Kid area. El Amawy et al. (2004) and Mehanna et al. (2004) considered

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Fig. 12.12 Schematic lithostratigraphic section (not to scale) of the Al-Wask ophiolite, after Gahlan et al. (2020b)

that serpentinized dunite and peridotite in the Imlig area of south Sinai have ophiolitic characteristics, on the basis of partly serpentinized peridotite within this mafic-ultramafic suite. These interpretations are controversial because the serpentinized harzburgites in Sinai are small (a few tens of meters long) and have no association with pillowed tholeiitic basalts, sheeted diabase dikes, or gabbros. Most researchers discount the presence of ophiolitic rocks in south Sinai. Azer and El-Gharbawy (2011) found that the mafic-ultramafic rocks of the Imlig area represent a layered intrusion, younger than the surrounding calc-alkaline syn-tectonic granodiorite. Mogahed (2020) considered the Kabr El-Bonaya ultramafic body to be an intrusion, younger than the surrounding island arc volcano-sedimentary succession (Wadi Kid Volcanics). The locations in the central and southern Eastern Desert of those ophiolitic bodies from which chemical analyses of chrome spinels are available (Table 12.3) are shown on the regional geologic map of the Eastern Desert of Egypt (Fig. 12.13). We will illustrate the general character of these Egyptian ophiolites with a detailed discussion of the Wadi Ghadir example. The most complete ophiolite section in the Eastern Desert of Egypt is exposed at Wadi Ghadir (El Bayoumi 1983), about 30 km southwest of Mersa Alam (#18 on the map in Fig. 12.13). In this area, ophiolitic rocks occur as allochthonous units within a mélange (e.g., El Sharkawy and El Bayoumi 1979; Takla et al. 1992; Basta et al. 1983; Abd

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El-Rahman et al. 2009) that was previously mapped as a geosynclinal metasedimentary sequence (El Ramly 1972). According to Kröner et al. (1992), the Wadi Ghadir ophiolite exposures mapped by El Sharkawy and El Bayoumi (1979) and El Bayoumi (1983) are the most complete and best-preserved sections through late Proterozoic upper oceanic crust anywhere in the world. The most complete and continuous section through the upper units of the Wadi Ghadir area ophiolite is exposed in the center of the map area shown in Fig. 12.14, where a continuous outcrop a few km long extends from the layered gabbro at the base to the pillow basalts with associated deep-sea sediments at the top. The compiled stratigraphy of the Ghadir ophiolite sequence (El Bayoumi 1983) is represented in Fig. 12.15, with the following elements, starting from the base: serpentinites with chromite lenses, ultramafic cumulates, metagabbro (layered and isotropic), sheeted dikes, pillowed basalt, and marine sediments. Serpentinites, dominantly after harzburgite and to a lesser extent after dunite and lherzolite, are frequently transformed into talc-carbonates, particularly along thrust faults and shear zones. They sometimes enclose chromite lenses and contain sparse relics of enstatite, diopside, and olivine. A thrust fault separates the mantle section from the overlying crustal section. The metagabbro sequence is of tholeiitic composition and begins at the base with cumulate pyroxenite and layered gabbro, followed upward by isotropic gabbro and hornblende gabbro. The hornblende gabbro frequently encloses small dioritic bodies and is cut by dikes of a hornblende-rich rock known as appinite. Sheeted dike swarms, also of tholeiitic composition, are found in-place in the sequence, between the underlying cumulate sequence and the overlying pillow basalts. They are interpreted as feeder conduits that delivered basalt to the sub-aqueous pillow lavas. Although the contacts between individual dikes are hardly discernible, asymmetric chilled margins have been recognized. Chilled margin asymmetry shows that the dikes typically intruded consecutively into one another before cooling was complete. The dikes vary in thickness between 1 and 10 m and form sub-parallel swarms. Some of the dikes extend downward into the cumulate gabbro unit, whereas others extend upwards through the overlying pillow lavas. Low-K tholeiitic basalts, several hundred meters thick, are sometimes pillowed; the pillows always indicate a consistent facing direction, right-side up with respect to the overall stratigraphy. Thinly bedded deep-water facies metasedimentary rocks are quite rare but some sections reach continuous thicknesses of *100–200 m. The dominant lithology is pelites (from pelagic clay protoliths) intercalated with thin chert and calcareous bands. Recent studies of the Wadi Ghadir ophiolite (Khalil and Azer 2007; Abd El-Rahman et al. 2009; Basta et al. 2011; Azer et al. 2019) agree on the fundamental interpretation of

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Fig. 12.13 Distribution of ophiolitic rocks in Eastern Desert of Egypt. Modified after Shackleton (1994). Localities of ophiolitic bodies from which chemical analyses of chrome spinels are available are numbered on the map

the terrain. The ophiolitic sequence was thrust over an island arc volcano-sedimentary succession. The ophiolitic units are variably modified by early seafloor alteration, followed by metasomatism and metamorphism from greenschist to lower-amphibolite facies, and subsequently altered again, including the development of talc-carbonate and quartz-carbonate rocks, especially along shear zones and fault planes. Despite the penetrative serpentinization of the ultramafic section, some samples contain relics of primary

chromian spinel, olivine, and pyroxenes. On the basis of bastite and mesh textures and normative mineralogy, the protoliths of the serpentinites can be identified as dominantly harzburgites, with virtually no primary diopside. The serpentinized peridotites have elevated whole-rock Mg#, Cr, and Ni; low whole-rock CaO and Al2O3; high NiO and Fo content in relict olivine; and high Cr# of fresh chromian spinel cores. All of these features are characteristic of formation in a fore-arc supra-subduction zone environment.

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Fig. 12.14 Geologic map of the Wadi Ghadir area. Modified after El Bayoumi (1983)

12.8

Mineralization of the ANS Ophiolites

In the ANS, the strong association between ophiolites and mineralization serves to emphasize the importance of better understanding the nature, composition, and evolution of the ophiolites. The ANS ophiolites host important mineral deposits, such as chromite, magnesite, talc, iron-nickel laterites, platinum-group elements, Cu–Ni–Co base metals, and gold (e.g., Botros 1993; Klemm et al. 2001; Kusky and Ramadan 2002; Ghoneim 2003; Ahmed and Hariri 2008; Azer 2013; Hamdy et al. 2018). Most prospects and productive mining sites are closely associated with strongly altered zones in the ANS ophiolites, particularly in the carbonatized serpentinites and listvenites (e.g., Osman 1995; Oweiss et al. 2001; Botros 2002, 2004; Al Jahdali et al. 2003; Al Jahdali 2004; Harbi et al. 2006; Zoheir and Lehmann 2011; Abd El-Rahman et al. 2012; Boskabadi et al. 2017; Gahlan et al. 2018). Therefore, the exploration for new

mineral deposits in the ANS is most likely to succeed if focused on the alteration products of ophiolites. In the following paragraphs, we will discuss two of the most important mineral associations in the ANS ophiolites, gold and chromite.

12.8.1 Gold Historically worked gold deposits are associated with ophiolitic rocks in the ANS. Gold mining in the alteration products of ANS ophiolites, especially those of the Eastern Desert of Egypt, extends back to Pharaonic times (Harraz 2000; Klemm and Klemm 2013). Gold is also known to occur in the intensely serpentinized and carbonatized rock of the ancient Hajr mine in the Nabitah fault zone (Hariri 1989). The Fawakhir gold mine shows gold concentrations up to *30 g/ton (Hussein 1990), supporting the suggestion of Buisson and Leblanc (1987) that Neoproterozoic

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Fig. 12.15 Compiled stratigraphy of the Wadi Ghadir ophiolite sequence. Modified after El Bayoumi (1983)

ophiolitic mantle sections are modestly enriched in gold. In detail, a spatial and genetic relationship has been observed between carbonatized ultramafics, subsequent granite intrusions, and gold mineralization. Apparently, the carbonatization concentrates gold by up to 1000 times relative to the original ultramafic rocks; interaction with hydrothermal systems associated with granite intrusions then further concentrates gold (Cox and Singer 1986; Azer 2008, 2013). However, the relation between Au enrichment and the carbonatization process still is controversial (e.g., Azer 2013; Boskabadi et al. 2017; Gahlan et al. 2018). Therefore, we can expect to better understand the former by focused studies of the latter. During serpentinization and alteration, gold enrichment is strongly correlated with enrichment in the fluid-mobile elements B, Li, S, As, Rb, Sr, Sb, Cs, Ba, Pb, and U; it is therefore presumably acting to some degree as a fluid-mobile element itself (e.g., Deschamps et al. 2013; Gahlan et al. 2018). A great deal of

experimental, thermodynamic, and field-based work worldwide has elucidated the temperature, fO2 pH, sulfide activity, and other parameters that regulate the fluid mobility of gold (e.g., Gammons and Williams-Jones, 1997; Lang and Baker, 2001; Pitcairn et al. 2014) and the structural conditions that allow fluids to migrate from scavenging zones to accumulation zones. These general economic geology paradigms provide guidance for exploring the specific context of carbonate alteration of serpentinite followed by felsic intrusion. In the Eastern Desert of Egypt, several studies have reported significant gold concentrations in discrete alteration zones in different ophiolite lithologies (e.g., Botros 1993, 2004; El-Mezayen et al. 1995; Hassaan et al. 1996; Osman 1995; Ramadan 1995, 2002; Ramadan and Kontny 2004; Ramadan et al. 2005; Azer 2013). For example, the El-Sid gold mine is confined to hydrothermal quartz veins originating at the contact between granite and serpentinite and

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extending into the serpentinite through a thick zone of graphite schist (El-Bouseily et al. 1985). El-Sid is a good example of vein-type mineralization hosted in sheared ophiolitic ultramafic rocks (Botros 2004). Abd El-Rahman et al. (2012) noted the dominance of vein-type gold deposits along the eastern side of the NW-trending Najd Fault System and attributed this pattern to the abundance of serpentinite bodies in the arc–fore-arc belt. Boskabadi et al. (2017) were the first to observe that gold is remobilized during the lizardite–antigorite transition. They suggested that Au, As, S, and other fluid-mobile elements were leached from altered serpentinites into passing CO2bearing fluids. Fluid inclusion thermometry and stable isotope compositions of gold deposits in ophiolitic fragments in the central Eastern Desert are similar to those in associated carbonate veins (Boskabadi et al. 2017), suggesting a close connection between carbonation of ultramafic rocks by mantle-derived CO2 and mobilization of Au into the economic deposits of the ANS. A recent study by Ferraris and Lorand (2015) in the Lherz peridotite massif suggests that inclusions in olivine are the main host for Au in ultramafic rocks. Likewise, Boskabadi et al. (2017) showed that up to 80% of the whole-rock Au budget in ANS serpentinite samples occurs as nano-inclusions in olivine, whereas only 20% of the Au is found in Cu–Fe–Ni sulfides. Occurrences of listvenites in the ANS have drawn special attention from exploration geologists in the last three decades because of their worldwide spatial and temporal association with gold deposits in ophiolitic complexes (Buisson and Leblanc 1985, 1986, 1987; Khalaf and Oweiss 1993; Harbi et al. 2003; Al Jahdali et al. 2003; Al Jahdali 2004; Ahmed and Hariri 2008; Zoheir and Lehmann Zoheir and Lehmann 2011; Gahlan et al. 2020c). Sadek et al. (2006) recognized Au and Ag mineralization in the listvenites of the Gebel Al Adraq area. The listvenites at G. Sirsir, Eastern Desert of Egypt, show gold concentrations up to 6584 ng/g as well as Ag and As enrichment (Gahlan et al. 2018). On the western side of the ANS, in Saudi Arabia, Thekair (1975) described some ancient gold workings confined to listvenite ridges developed at the margins of serpentinite masses along the Al-Amar fault. Gold-bearing quartz veins are recognized also in altered carbonated ultramafic hosts at Bir Tawilah and Jabal Ghadarah which (along with other ultramafic outcrops) both lie along the major Bir Tawilah thrust fault (Labbé 1984; Couturier 1986; Billa 1987). The listvenite of Bir Tawilah hosts two types of mineralized veins: low-grade veins with gold values up to 2 g/t and high-grade veins with 11–42 g/t (Couturier 1986; Labbé 1984). The listvenites associated with serpentinites from Jabal Ghadarah have gold that ranges between 0.3 and 4.91 g/t (Labbé 1984; Al Jahdali et al. 2003; Al Jahdali 2004; Harbi et al. 2006). Listvenite in the Gebel Zalm area once again contains anomalous gold values (0.28 g/t) (Harbi

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et al. 2003). Sheared carbonatized serpentinite (i.e., listvenite) is the host rock for gold mineralization in the Hajr gold mine (Hariri and Makkawi 2004). Gahlan et al. (2020c) found that the formation of listvenite at Jabal Ess was accompanied by leaching and removal of some components but the accumulation of CO3, SiO2, REE (especially Eu), Au, Zn, As, Sb, and K in the solid products. Also, they show that listvenitization at Jabal Ess concentrated Au in sub-economic to economic proportions; measured Au concentrations are 4–2569 ng/g in carbonate listvenite, 43–3117 ng/g in silica-carbonate listvenite, and 5–281 ng/g in silica listvenite, compared to the low-grade Au concentration in the host serpentinite (0.5– 1.7 ng/g). The listvenites of Jabal Ess merit further exploration for gold.

12.8.2 Chromitite Chromite deposits are commonly hosted by the serpentinized ultramafic rocks that are widely distributed in the ANS. These chromitite ores bodies have different shapes and sizes, including pods, lenses, thin layers, and schlieren (Al-Shanti and El-Mahdy 1988; Habtoor et al. 2017; Abuamarah et al. 2020). The chromitite pods, both massive and disseminated types, are commonly concentrated in the shallowest parts of the ophiolitic sequences, closest to the petrologic Moho (e.g., Ahmed et al. 2001; Gahlan et al. 2015; Abuamarah et al. 2020). The podiform deposits are commonly concordant to sub-concordant with the fabrics of the host rocks and surrounded by an envelope of dunite that grades outward to harzburgite (Habtoor et al. 2017; Abuamarah et al. 2020). In the field, chromitite bodies form prominent outcrops as a result of their resistance to weathering. They may be isolated or in clusters. Most have relatively sharp contacts with the surrounding serpentinites, but in some cases the boundaries of the chromitites are diffuse. Although none are exceptionally large, the ANS chromitites are mostly of metallurgical grade (Cr2O3 > 40 wt% and Al2O3 < 20 wt%) (e.g., Abu El Ela and Farahat 2010; Habtoor et al. 2017; Azer et al. 2019; Abuamarah et al. 2020). The chromitite bodies are interpreted to have formed just below the contact between the oceanic crust and mantle sections (i.e., the petrologic Moho; Abuamarah et al. 2020). Discussions of the origin and geotectonic setting of chromitite ore bodies in general have been echoed by specific debate about the podiform chromitites of the ANS (e.g., Abu El Ela and Farahat 2010; Habtoor et al. 2017; Khedr and Arai 2017; Abuamarah et al. 2020). The common observation of a dunite envelope surrounding podiform chromitites within the generally harzburgite-dominated depleted mantle section (e.g., Ahmed et al. 2001; Abu El

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Ela and Farahat 2010; Ahmed 2013; Khedr and Arai 2016, 2017; Habtoor et al. 2017; Abuamarah et al. 2020) is one of the main motivators for exploration of wall-rock interaction mechanisms for their formation (Arai and Yurimoto 1994; Zhou et al. 1994, 2005; González-Jiménez et al. 2011, 2014). Like podiform chromitites worldwide, those of the ANS mantle section are well-explained as the result of melt– peridotite interaction and subsequent melt mixing in the SSZ upper mantle (e.g., Lago et al. 1982; Paktunc 1990; Arai and Yurimoto 1994, 1995; Zhou et al. 1994, 1996; Arai 1997; Ahmed et al. 2001; Ahmed 2013; Azer et al. 2019; Abuamarah et al. 2020). Abuamarah et al. (2020) recently offered a thermodynamic model of the process of mixing between unreacted basaltic melt and reacted boninite melt in a dunite conduit and confirmed that mixing raises the saturation temperature of chromite above that of olivine. However, the process is relatively inefficient; at least 8.3 kg of melt is required to crystallize 1 g of monomineralic chromite.

12.9

Discussion

Ophiolites make up a small but tectonically important part of the Neoproterozoic rocks of the ANS. In the complete ophiolitic sequence in the ANS, the top consists of fine-grained oceanic sediments (mudstones, cherts, calcareous oozes, etc.). Below this are pillow basalts, which form when hot lava is extruded into deep water. These rocks are often extensively altered by interaction with seawater. Sheeted dikes occur below the pillow basalts. These dikes represent conduits that fed eruption of the sub-aqueous pillow basalts; they typically intrude one another consecutively before cooling is complete. Underlying the volcanic sequence is a thick layer of its intrusive equivalent, gabbros and related rocks. The upper part of the gabbro is typically not stratified, but the basal part of the gabbro commonly contains cumulate layers, including troctolites and pyroxenites. The base of most ANS ophiolites consists of harzburgite, often forming strongly deformed or transposed compositional layering, known as harzburgite tectonite. However, because of folding and shearing, the boundaries between these major units are often faulted and the majority of the ANS ophiolites lack one or more of their diagnostic lithologies. All the ANS ophiolites are strongly deformed, metamorphosed, and altered by silicification and carbonatization. Low-grade greenschist facies metamorphism predominates, but in places, the rocks reach amphibolite grade. The serpentinites represent the most important and distinctive lithology in the dismembered ophiolites (note, however, that isolated serpentinite bodies cannot be confidently assigned to ophiolitic origin without detailed study). Alteration of serpentinite resulted in the development of talc-carbonates, listvenite, and magnesite.

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Ongoing studies are aimed at refining the interpretation of the ANS ophiolites and eliminating ambiguities. Modern petrologic and geochemical techniques, such as laser-ablation inductively-coupled plasma mass spectrometry, allow in situ measurements of diagnostic trace element contents of individual phases as a complement to more traditional whole-rock techniques. Isotopic studies—especially Nd, Hf, and Os—are needed to define the evolution of the mantle source regions for the ophiolites. Also, more and more specific geochronological constraints on the age of ANS ophiolites are needed. For example, precise determination of the emplacement age of an ophiolite allows biostratigraphers and those studying global environmental changes through the Neoproterozoic to place data from the deep water marine sequences at the top of the section in stratigraphic context. Economic considerations also favor continued study of ANS ophiolites. The high grade of ANS chromites may be rich enough to mine, but their small sizes require extensive mapping to keep locating new pods. Given the observation that gold mineralization often appears to be related to carbonate alteration of ophiolites, we can expect to better understand the concentration of Au through focused studies of carbonate alteration. At present, we have only a preliminary understanding of pervasive carbonate alteration of ANS ophiolitic peridotites; more study is required to better understand the age of this alteration, the origins and fluxes of the fluids, and the thermodynamic pathways of fluid-rock interaction and element mobility. Acknowledgements The collaboration between the authors began with support from the US Agency for International Development. PDA’s ongoing work on Egyptian rocks is supported in part by the National Science Foundation, award 1947616. The authors are grateful to Prof. Zakaria Hamimi for editorial handling, as well as the two reviewers (Prof. Shoji Arai and Dr. Hisham A. Gahlan).

References Abd El-Rahman Y, Helmy HM, Shibata T, Yoshikawa M, Arai S, Tamur A (2012) Mineral chemistry of the Neoproterozoic Alaskan-type Akarem Intrusion with special emphasis on amphibole: Implications for the pluton origin and evolution of subduction-related magma. Lithos 155:410–425 Abd El-Rahman Y, Polat A, Dilek Y, Fryer B, El-Sharkawy M, Sakran S (2009) Geochemistry and tectonic evolution of the Neoproterozoic Wadi Ghadir ophiolite, Eastern Desert, Egypt. Lithos 113:158–178 Abdallah SE, Azer MK, El Shammari AS (2019) Petrological and geochemical evolution of Ediacaran rare-metal bearing A-type granites from Jabal Aja complex, northern Arabian Shield. Acta Geologica Sinica, Saudi Arabia. https://doi.org/10.1111/1755-6724. 13825 Abdel Aal AY, Farahat ES, Hoinken G, El-Mahallawi MM (2003) Ophiolites from the Egyptian Shield: a case for a possible inter-arc origin. Mitt Oesterr Ges 148:81–83

324 Abdel Khalek ML, Abdel Maksoud MA, Abdel Tawab MA, Oweiss KA (1994) Geochemistry of the St. Catherine basement rocks, Sinai, Egypt. Sci J Qatar Univ 14:134–145 Abdel-Karim AM (2000) Chlorite schists and rodingites in the mafic-ultramafic rocks from the central Eastern Desert of Egypt: Petrogenesis and metamorphic history. M.E.R.C. Ain Shams University, Earth Science Series 14:150–170 Abdel-Karim AM, Elwan WI, Helmy H, El-Shafey SA (2014) Spinels, Fe-Ti oxides minerals, apatites and carbonates hosted in the ophiolites of Eastern Desert of Egypt: Mineralogy and chemical aspects. Arab J Geosci 7:693–709 Abdel-Karim AAM, Ali S, Helmy HM, El-Shafei SA (2016) A fore-arc setting of the Gerf ophiolite, Eastern Desert, Egypt: evidence from mineral chemistry and geochemistry of ultramafites. Lithos, 263:52–65 Abdel-Karim AM, El-Shafei SA (2018) Mineralogy and chemical aspects of some ophiolitic metaultramafics, central Eastern Desert, Egypt: Evidences from Cr-spinels, sulphides and gangues. Geol J 53:580–599 Abdel-Karim AM, El-Shafei SA, Azer MK (2020a) The Neoproterozoic ophiolitic ultramafic rocks in Eastern Desert of Egypt: implications for petrogenesis and metasomatic processes. Inter Geol Rev. https://doi.org/10.1080/00206814.2019.1708816 Abdel-Karim AM, El-Shafei SA, Azer MK (2020b) Petrology and geochemistry of ophiolitic metaperidotites from the Eastern Desert of Egypt: Insights into geodynamic evolution and post-melting alteration processes. Act Geol Sin (Final accepted) Abdel Rahman EM (1993) Geochemical and geotectonic controls of the metallogenic evolution of selected ophiolite complexes from the Sudan. Berliner geowissenshaftliche Abhandlungen 145 Abdelsalam MG, Stern RJ (1993) Structure of the late Proterozoic Nakasib suture, Sudan. J Geol Soc London 150:1065–1074 Abdelsalam MG, Stern RJ (1996) Sutures and shear zones in the Arabian-Nubian Shield. J Afr Earth Sci 23(3):289–310 Abu El Ela FF, Farahat ES (2010) Neoproterozoic podiform chromitites in serpentinites of the Abu Meriewa-Hagar Dungash district, Eastern Desert, Egypt: Geotectonic implications and metamorphism. Island Arc 19(1):151–164 Abu El-Enen MM, Makroum FM (2003) Tectonometamorphic evolution of the northeastern Kid Belt, Southeast Sinai, Egypt. Ann Geol Surv Egypt VXXVI: 19–37 Abu-Alam TS, Hamdy MM (2014) Thermodynamic modelling of Sol Hamed serpentinite, South Eastern Desert of Egypt: Implication for fluid interaction in the Arabian-Nubian Shield ophiolites. J Afr Earth Sci 68:1–14 Abuamarah BA (2019a) Geochemistry and fore-arc evolution of upper mantle peridotites in the Cryogenian Bir Umq ophiolite, Arabian Shield, Saudi Arabia. Inter Geol Rev 62(5):630–648 Abuamarah BA (2019b) Petrogenetic evolution of Cryogenian Halaban ophiolite, Saudi Arabia: A fragment of fore-arc oceanic lithosphere mantle. Lithos 356: Abuamarah BA, Asimow PD, Azer MK, Ghrefat H (2020) Suprasubduction-zone origin of the podiform chromitites of the Bir Tuluhah ophiolite, Saudi Arabia, during Neoproterozoic assembly of the Arabian Shield. Lithos 360–361: Ahmed AH (2013) Highly depleted harzburgite–dunite–chromitite complexes from the Neoproterozoic ophiolite, south Eastern Desert, Egypt: A possible recycled upper mantle lithosphere. Precamb Res 233:173–192 Ahmed AH, Arai S, Attia AK (2001) Petrological characteristics of podiform chromitites and associated peridotites of the Pan African ophiolite complexes of Egypt. Mineral Deposit 36:72–84 Ahmed AH, Habtoor A (2015) Heterogeneously depleted Precambrian lithosphere deduced from mantle peridotites and associated

M. G. Shahien et al. chromitite deposits of Al’Ays ophiolite, Northwestern Arabian Shield, Saudi Arabia. Ore Geol Rev 67:279–296 Ahmed AH, Harbi HM, Habtoor AM (2012) Compositional variations and tectonic settings of podiform chromitites and associated ultramafic rocks of the Neoproterozoic ophiolite at Wadi Al Hwanet, northwestern Saudi Arabia. J Asian Earth Sci 56:118–134 Ahmed Z, Hariri MM (2008) Neoproterozoic ophiolites as developed in Saudi Arabia and their oceanic and pericontinental domains. Arab J Sci Eng 33:17–54 Al Jahdali NS (2004) Geology of Jabal Ghadarah area, Bir Tawilah district with special emphasis on listvenite as a potential source for gold in the Kingdom of Saudi Arabia. M.Sc. Thesis, King AbdulAziz University, Jeddah, KSA, 197 p Al Jahdali NS, Harbi H, Eldougdoug A (2003) Gold bearing listvenite in Jabal Al Ghadarah area, Central Arabian Sheild, Kingdom of Saudi Arabia. In: Proceedings of the 8th Arab Conference on Mineral Resources, Sanaá. Republic of Yemen, 13–16 October, 1:26–43 Ali K, Azer M, Gahlan H, Wilde S, Samuel M, Stern R (2010) Age constraints on the formation and emplacement of Neoproterozoic ophiolites along the Allaqi-Heiani Suture, South Eastern Desert of Egypt. Gondwana Res 18(4):583–595 Ali RA, Maurice AE, Pitcairn IK, Ahmed AH, Azer MK, Boskabadi A, Bakhit BR, Shahien MG (2020a) Neoproterozoic and Cretaceous mantle oxidation states: Controls and heterogeneity through time. Lithos 356–357: Ali RA, Pitcairn IK, Maurice AE, Azer MK, Bakhit BR, Shahien MG (2020b) Petrology and geochemistry of ophiolitic ultramafic rocks and chromitites across the Eastern Desert of Egypt: Insights into the composition and nature of a Neoproterozoic mantle and implication for the evolution of SSZ system. Precam Res 337: Allen DE, Seyfried WEJ (2003) Compositional controls on vent fluids from ultramafic-hosted hydrothermal systems at mid-ocean ridges: An experimental study at 400 °C, 500 bars. Geochim Cosmochim Acta 67(8):1531–1542 Al-Saleh AM, Boyle AP, Mussett AE (1998) Metamorphism and 40 Ar/39Ar dating of the Halaban ophiolite and associated units: evidence for two-stage orogenesis in the eastern Arabian shield. J Geol Soc 155(1):165–175 Al-Shanti AM, El-Mahdy OR (1988) Geological Studies and Assessment of Chromite Occurrences in Saudi Arabia: KACST Project No. AT-6-094 Al-Shanti MMS (1982) Geology and mineralization of the Ash Shizm-Jabal Ess area (Doctoral dissertation, Ph. D. thesis. King Abdulaziz University), Jiddah, 291 p Anzil PA, Guereschi AB, Martino RD (2012) Mineral chemistry and geothermometry using relict primary minerals in the La Cocha ultramafic body: A slice of the upper mantle in the Sierra Chica of Córdoba, Sierras Pampeanas, Argentina. J South Amer Earth Sci 40:38–52 Arai S (1975) Contact metamorphosed dunite–harzburgite complex in the Chugoku district, western Japan. Contrib Mineral Petrol 52:1–16 Arai S (1992) Chemistry of chromian spinel in volcanic rocks as a potential guide to magma chemistry. Mineral Mag 56:173–184 Arai S (1994) Compositional variation of olivine–chromian spinel in Mg-rich magmas as a guide to their residual spinel peridotites. J Volcanol Geotherm Res 59:279–293 Arai S (1997) Control of wall-rock composition on the formation of podiform chromitites as a result of magma/peridotite interaction. Resour Geol 47:177–187 Arai S, Shimizu Y, Ismail SA, Ahmed AH (2006) Low-T formation of high-Cr spinel with apparently primary chemical characteristics within podiform chromitite from Rayat, northeastern Iraq. Miner Mag 70:499–508

12

Neoproterozoic Ophiolites of the Arabian-Nubian Shield

Arai S, Yurimoto H (1994) Podiform chromitites of Tari-Misaka ultramafic complex, Southwest Japan, as mantle-melt interaction products. Econ Geol 89:1279–1288 Arai S, Yurimoto H (1995) Possible sub-arc origin of podiform chromitites. Island Arc 4:104–111 Arif M, Jan MQ (2006) Petrotectonic significance of the chemistry of chromite in the ultramafic-mafic complexes of Pakistan. J Asian Earth Sci 27:628–646 Azer MK (2008) Origin of Neoproterozoic Ophiolitic Serpentinites and Their Economic Potentialities, Eastern Desert, Egypt. 33rd International Geological Congress, Oslo-Norway (Abstract) Azer MK (2013) Evolution and economic significance of listvenites associated with Neoproterozoic ophiolites in south Eastern Desert, Egypt. Geologica Acta 11:113–128 Azer MK (2014) Petrological studies of Neoproterozoic serpentinized ultramafics of the Nubian Shield: spinel compositions as evidence of the tectonic evolution of Egyptian ophiolites. Acta Geol Pol 64 (1):123–137 Azer MK, El-Gharbawy RI (2011) The Neoproterozoic layered mafic– ultramafic intrusion of Gabal Imleih, south Sinai, Egypt: implications of post-collisional magmatism in the north Arabian–Nubian Shield. J Afr Earth Sci 60(4), 253-272. Azer MK, Asimow PD (2021) Petrogenetic evolution of the Neoproterozoic igneous rocks of Egypt. In: Hamimi Z, Arai S, Fowler TJ, El Bialy M (ed) The Geology of the Arabian-Nubian Shield. Springer Nature, Cham, Switzerland Azer MK, Gahlan HA, Asimow PD, Mubarak HS, Al-Kahtany KM (2019) Multiple stages of carbonation and element redistribution during formation of ultramafic-hosted magnesite in Neoproterozoic ophiolites of the Arabian-Nubian Shield. Egypt J Geol 127:81–107 Azer MK, Khalil AES (2005) Petrological and mineralogical studies of Pan-African serpentinites at Bir Al-Edeid area. Central Eastern Desert, Egypt. J Afr Earth Sci 43:525–536 Azer MK, Samuel MD, Ali KA, Gahlan HA, Stern RJ, Ren M, Moussa HE (2013) Neoproterozoic ophiolitic peridotites along the Allaqi-Heiani Suture, South Eastern Desert, Egypt. Mineral Petrol 107:829–848 Azer MK, Stern RJ (2007) Neoproterozoic (835–720 Ma) serpentinites in the Eastern Desert, Egypt: Fragments of fore-arc mantle. J Geol 115:457–472 Bakor AR, Gass IG, Neary CR (1976) Jabal al Wask, northwest Saudi Arabia: an Eocambrian back-arc ophiolite. Earth and Planetary Sci Lett 30(1):1–9 Barnes SJ, Roeder PL (2001) The range of spinel composition in terrestrial mafic ultramafic rocks. J Petrol 42:2279–2302 Basta EZ, Church WR, Hafez AMA, Basta FF (1983) Proterozoic ophiolitic mélange and associated rocks of Gebel Ghadir area, Eastern Desert, Egypt. In: Proceedings of the Fifth International Conference on Basement Tectonics, October 16–18, Cairo, pp 115– 124 Basta FF, Maurice AE, Bakhit BR, Ali KA, Manton WI (2011) Neoproterozoic contaminated MORB of Wadi Ghadir ophiolite, NE Africa: Geochemical and Nd and Sr isotopic constraints. J Afr Earth Sci 59(2–3):227–242 Beccaluva L, Coltorti M, Giunta G, Siena F (2004) Tethyan vs. Cordilleran ophiolites: a reappraisal of distinctive tectono-magmatic features of suprasubduction complexes in relation to the subduction mode. Tectonophysics 393(1–4):163–174 Beccaluva L, Macciotta G, Piccardo G, Zeda O (1989) Clinopyroxene composition of ophiolite basalts as petrogenetic indicator. Chem Geol 77(3–4):165–182 Bédard J (1999) Petrogenesis of boninites from the Betts Cove ophiolite, Newfoundland, Canada: identification of subducted source components. J Petrol 40(12):1853–1889

325 Bentor YK (1985) The crustal evolution of the Arabo-Nubian Massif with special reference to Sinai Peninsula. Precamb Res 28:1–74 Best MG (2003) Igneous and metamorphic petrology. Blackwell Science Ltd., Oxford Billa M (1987) Exploration of the Jabal Ghadarah gold deposit, Bir Tawilah district. Saudi Arabian Deputy Ministry for Mineral Resources, Open-File-Report BRGM-OF-06-15 Blasband B, White S, Brooijmans P, De Boorder H, Visser W (2000) Late Proterozoic extensional collapse in the Arabian-Nubian shield. J Geol Soc 157(3):615–628 Bloomer SH, Taylor B, MacLeod CJ, Stern RJ, Fryer P, Hawkins JW, Johnson L (1995) Early arc volcanism and ophiolite problem: A perspective from drilling in the Western Pacific. In: Taylor B, Natland J (eds) Active Margins and Marginal Basins of the Western Pacific, Geophysical Monograph, vol 88. American Geophysical Union. Washington, DC, pp 1–30 Boskabadi A, Pitcairn IK, Broman C, Boyce A, Teagle DAH, Cooper MJ, Azer MK, Mohamed FH, Stern RJ, Majka J (2017) Carbonate alteration of ophiolitic rocks in the Arabian-Nubian Shield of Egypt: sources and compositions of the carbonating fluid and implications for the formation of Au deposits. Inter Geol Rev 59(4):391–419 Botros NS (1993) New prospects for gold mineralization in Egypt. Ann Geol Surv Egypt 19:47–56 Botros NS (2002) Metallogeny of gold in relation to the evolution of the Nubian Shield in Egypt. Ore Geol Rev 19:137–164 Botros NS (2004) A new classification of the gold deposits of Egypt. Ore Geol Rev 25:1–37 Buisson G, Leblanc M (1985) Gold in carbonatized ultramafic rocks from ophiolite complexes. Econ Geol 80(7):2028–2029 Buisson G, Leblanc M (1986) Gold-bearing listvenites (carbonatized ultramafic rocks) from ophiolite complexes. In: Gallagher JM, Ixer RA, Neary CR (eds) Metallogeny of Basic and Ultrabasic Rocks. Institution of Mining and Metallurgy, London, pp 121–132 Buisson G, Leblanc M (1987) Gold in mantle peridotites from Upper Proterozoic ophiolites in Arabia, Mali and Morocco. Econ Geol 82:2091–2097 Chevremont P, Johan Z (1982) Wadi Al Hwanet- Jabal Iss ophiolitic complex. Saudi Arabian Deputy Ministry for Mineral Resources Open-File Report BRGM-OF-02-14, 30 p Claesson S, Pallister JS, Tatsumoto M (1984) Samarium-neodymium data on two late Proterozoic ophiolites of Saudi Arabia and implications for crustal and mantle evolution. Contrib Mineral Petrol 85(3):244–252 Coish RA, Gardner P (2004) Suprasubduction-zone peridotite in the northern USA Appalachians: evidence from mineral composition. Miner Mag 68:699–708 Couturier P (1986) Geology and evaluation of the mining potential of the Bir Tawilah main gold prospect. Ministry of Petroleum and Mineral Resources, Deputy Ministry for Mineral Resources, Jiddah, Kingdom of Saudi Arabia. Open-File Report BRGM-OF-06-21. 21 p Cox DP, Singer DA (1986) Mineral deposit models (Vol 1693). Bulletin: US Government Printing Office Deer W, Howie R, Zussman J (1992) An introduction to the rock forming minerals. Second Edition, Longman Scientific and Technical, London, 696 p Deschamps F, Godard M, Guillot S, Hattori K (2013) Geochemistry of subduction zone serpentinites. A review. Lithos 178:96–127 Dick HJB, Bullen T (1984) Chromian spinel as a petrogenetic indicator in abyssal and Alpine-type peridotites and spatially associated lavas. Contrib Mineral Petrol 86:54–76 Dick HJB, Natland JH (1996) Late-stage melt evolution and transport in the shallow mantle beneath the East Pacific Rise. In

326 Proceedings-Ocean Drilling Program Scientific Results (pp 103– 134). National Science Foundation Dilek Y, Furnes H (2014) Ophiolites and Their Origins. Elements 10 (2):93–100 Dilek Y, Ahmed Z (2003) Proterozoic ophiolites of the Arabian Shield and their significance in Precambrian tectonics. In: Dilek Y, Robinson PT (eds) Ophiolites in Earth History. Geol Soc London Spec Pub 218, pp 685–700 Dilek Y, Furnes H, Shallo M (2008) Geochemistry of the Jurassic Mirdita Ophiolite (Albania) and the MORB to SSZ evolution of a marginal basin oceanic crust. Lithos 100(1–4):174–209 Dilek Y, Newcomb S (2003) Ophiolite concept and the evolution of geological thought. Geol Soc Amer Spec Paper 373:504 Dilek Y, Thy P (2009) Island arc tholeiite to boninitic melt evolution of the Cretaceous Kizildag (Turkey) ophiolite: Model for multi-stage early arc–fore-arc magmatism in Tethyan subduction factories. Lithos 113(1–2):68–87 Dilek Y, Furnes H (2011) Ophiolite genesis and global tectonics: Geochemical and tectonic fingerprinting of ancient oceanic lithosphere. Geol Soc Amer Bull 123(3–4):387–411 Douville E, Charlou JL, Oelkers EH, Bienvenu P, Jove Colon CF, Donval JP, Fouquet Y, Prieur D, Appriou P (2002) The rainbow vent fluids (36°14’ N, MAR): The influence of ultramafic rocks and phase separation on trace metal contents on Mid-Atlantic Ridge hydrothermal fluids. Chem Geol 184:37–48 Dunlop HM, Kemp P, Calvez J-Y (1986) Geochronology and isotope geochemistry of the Bir Umq mafic-ultramafic complex and Arj group volcanic rocks, Mahd adh Dhahab quadrangle, central Arabian Shield. Saudi Arabian Deputy Ministry for Mineral Resources Open-File Report BRGM-OF-07-7, p 38 El Amawy M, Wetait M, Mehanna A, Soliman F, Ghabour Y (2004) Geology and structural evolution of the basement rocks of Imlig area, southwest Sinai, Egypt. Ann Geol Surv Egypt 27:19–33 El-Bahariya GA (2003) Petrology and origin of Pan-African serpentinites with particular reference to chromian spinel composition, Eastern Desert, Egypt. Implications for supra-subduction zone ophiolite. In 3rd Int Conf Geol Afr Assuit Egypt, 2003, vol 1, pp 371–388 El Bayoumi RM (1983) Ophiolites and mélange complex of Wadi Ghadir Area, Eastern Desert, Egypt. Bull Inst Appl Geol King Abdul Aziz Univ Jeddah 6:329–342 El Dien HMG, Arai S, Doucet L-S, Li Z-X, Kil Y, Fougerouse D, Reddy S, Saxey D, Hamdy M (2019) Cr-spinel records metasomatism not petrogenesis of mantle rocks. Nature Comm 10:5103 El Dien HMG, Hamdy M, El-Ela ASA, Abu-Alam T, Hassan A, Kil Y, Mizukami T, Soda Y (2016) Neoproterozoic serpentinites from the Eastern Desert of Egypt: insights into Neoproterozoic mantle geodynamics and processes beneath the Arabian-Nubian Shield. Precam Res 286:213–233 El Naby HA, Frisch W (1999) Metamorphic sole of Wadi Haimur-Abu Swayel ophiolite: Implications on late Proterozoic accretion. In: Wall HD, Greiling RO (eds.) Aspects of Pan-African Tectonics. Heidelberg, Germany, pp 9–14 El Ramly MF (1972) A new geological map for the basement rocks in the Eastern and Southwestern desert of Egypt: Scale 1: 1,000,000. Annals of the Geological Survey of Egypt El Sharkawy MA, El Bayoumi RM (1979) The ophiolites of Wadi Ghadir area, Eastern Desert, Egypt. Ann Geol Surv Egypt 9:125– 135 El-Bouseily AM, El-Dahhar MA, Arslan AI (1985) Ore-microscopic and geochemical characteristics of gold-bearing sulfide minerals, El Sid Gold mine, Eastern Desert, Egypt. Mineral Deposit 20:194–200 El-Gaby S, List FK, Tehrani R (1988) Geology, evolution and metallogenesis of the Pan-African Belt in Egypt. In: El-Gaby S,

M. G. Shahien et al. Greiling RO (eds) The Pan-African Belt of northeast Africa and adjacent areas. Vieweg, Braunschweig, pp 17–68 El-Gaby S, List FK, Tehrani R (1990) The basement complex of the Eastern Desert and Sinai. In: Said R (ed) The Geology of Egypt. Balkema, Rotterdam, pp 175–184 El-Mezayen AM, Hassaan MM, El-Hadad M, Hassanein MM (1995) Petrography, geochemistry and ore microscopy of Abu Marawat metavolcanics and associated gold mineralisation, North Eastern Desert, Egypt. Bull Fac Sci Al-Azhar Univ 6(2):1999–2021 El-Sayed MM, Furnes H, Mohamed FH (1999) Geochemical constraints on the tectonomagmatic evolution of the late Precambrian Fawakhir ophiolite, Central eastern Desert, Egypt. J Afr Earth Sci 29:515–533 Evans BW (1977) Metamorphism of Alpine peridotites and serpentinites. Ann Rev Earth Planet Sci 5:397–447 Evans BW (2010) Lizardite versus antigorite serpentinite: magnetite, hydrogen, and life (?). Geology 38:879–882 Evans WB, Frost BR (1975) Chrome-spinel in progressive metamorphism–a preliminary analysis. Geochim Cosmochim Acta 39:959–972 Farahat ES (2008) Chrome-spinels in serpentinites and talc carbonates of the El Ideid-El-Sodmein District, central Eastern Desert, Egypt: their metamorphism and petrogenetic implications. Chem Erde 68:193–205 Farahat ES, El Mahallawi MM, Hoinkes G, Abdel Aal AY (2004) Continental back-arc basin origin of some ophiolites from the Eastern Desert of Egypt. Mineral Petrol 82:81–104 Ferraris C, Lorand JP (2015) Novodneprite (AuPb3), anyuiite [Au(Pb, Sb)2] and gold micro- and nano-inclusions within plastically deformed mantle-derived olivine from the Lherz peridotite (Pyrenees, France): A HRTEM–AEM–EELS study. Phys Chem Miner 42:143–150 Fitches WR, Graham RH, Hussein IM, Ries AC, Shackleton RM, Price RC (1983) The late Proterozoic ophiolite of Sol Hamed, NE Sudan. Precam Res 19(4):385–411 Frost BR (1991) Stability of oxide minerals in metamorphic rocks. Rev Miner Geoch 25:469–488 Furnes H, De Wit M, Dilek Y (2014) Four billion years of ophiolites reveal secular trends in oceanic crust formation. Geo Front 5 (4):571–603 Gahlan HA (2006) Petrological characteristics of the mantle section in the Proterozoic ophiolites from the Pan-African Belt (Doctoral dissertation, Ph. D. thesis, Kanazawa University, Kanazawa, Japan, p 227) Gahlan H, Arai S (2009) Carbonate-orthopyroxenite lenses from the Neoproterozoic Gerf ophiolite, South Eastern Desert, Egypt: The first record in the Arabian Nubian Shield ophiolites. J Afr Earth Sci 53:70–82 Gahlan H, Azer M, Khalil AES (2015) The Neoproterozoic Abu Dahr ophiolite, South Eastern Desert, Egypt: petrological characteristics and tectonomagmatic evolution. Mineral Petrol 109:611–630 Gahlan H, Arai S, Abu El-Ela F, Tamura A (2012) Origin of Wehrlite Cumulates in the Moho Transition Zone of the Neoproterozoic Ras Salatit Ophiolite, Central Eastern Desert, Egypt: Crustal Wehrlites with Typical Mantle Characteristics. Contrib Mineral Petrol 163 (2):225–241 Gahlan HA, Azer MK, Asimow PD (2018) On the relative timing of listvenite formation and chromian spinel equilibration in serpentinites. Amer Mineral 103:1087–1102 Gahlan HA, Azer MK, Asimow PD, Mubarak HS, Al-Kahtany KM (2020a) Petrological characteristics of the Neoproterozoic Ess ophiolite mantle section, Arabian Shield, Saudi Arabia: a mineral chemistry perspective. Inter J Earth Sci 109:239–251 Gahlan HA, Azer MK, Asimow PD, Al-Kahtany KM (2020b) Genesis and geodynamic evolution of serpentinized ultramafics and the

12

Neoproterozoic Ophiolites of the Arabian-Nubian Shield

associated magnesite deposits, Al-Wask ophiolite, Arabian Shield, Saudi Arabia. Amer J Sci 320:236–279 Gahlan HA, Azer MK, Asimow PD, Al-Kahtany KM (2020c) Petrogenesis of gold-bearing listvenites from the carbonatized mantle section of the Neoproterozoic Ess ophiolite, Western Arabian Shield, Saudi Arabia. Lithos 12: Gammons CH, Williams-Jones AE (1997) Chemical mobility of gold in the porphyry-epithermal environment. Econ Geol 92(1):45–59 Gasparik T (1987) Orthopyroxene thermobarometry in simple and complex systems. Contrib Mineral Petrol 96(3):357–370 Ghoneim MF (2003) Origin of magnesite veins in serpentinites from Mount El Rubshi and Mount El-Maiyit, Eastern Desert, Egypt. Arch Mineral 54:41–63 González-Jiménez J-M, Griffin WL, Proenza JA, Gervilla F, O’Reilly SY, Akbulut M, Pearson NJ, Arai S (2014) Chromitites in ophiolites: how, where, when, why? Part II. A review and new ideas on the crystallization of chromitites. Lithos 189:140–158 González-Jiménez JM, Kerestedjian T, Proenza JA, Gervilla F (2009) Metamorphism on chromite ores from the Dobromirtsi ultramafic Massif, Rhodope Mountains (SE Bulgaria). Geol Acta 7:413–429 González-Jiménez JM, Proenza JA, Gervilla F, Melgarejo JC, Blanco-Moreno JA, Ruiz-Sánchez R, Griffin WL (2011) High-Cr and high-Al chromitites from the Sagua de Tánamo district, MayaríCristal ophiolitic massif (eastern Cuba): Constraints on their origin from mineralogy and geochemistry of chromian spinel and platinum-group elements. Lithos 125(1–2):101–121 Habtoor AM, Ahmed AH, Akizawa N, Harbi H, Arai S (2017) Chemical homogeneity of high-Cr chromitites as indicator for widespread invasion of boninitic melt in mantle peridotite of Bir Tuluha ophiolite, Northern Arabian Shield, Saudi Arabia. Ore Geol Rev 90:243–259 Hadley DG (1987) Geologic map of the Sahl Al Matran quadrangle, sheet 26C, Kingdom of Saudi Arabia. Saudi Arabian Deputy Ministry for Mineral Resources Geoscience Map GM-86, scale 1:250,000 Hamdy MM, Abu El-Ela AM, Hassan AM, Kil Y, El Dien HMG (2013a) Subduction-related cryptic metasomatism in fore-arc to nascent fore-arc Neoproterozoic mantle peridotites beneath the Eastern Desert of Egypt: mineral chemical and geochemical evidences. 15: EGU2013-14056 (abstract) Hamdy MM, Harraz HZ, Aly GA (2013b) Pan-African (intraplate and subduction-related?) metasomatism in the Fawakhir ophiolitic serpentinites, Central Eastern Desert of Egypt: Mineralogical and geochemical evidences. Arab J Geo 6:13–33 Hamdy MM, El Dien HMG (2017) Nature of serpentinization and carbonation of ophiolitic peridotites (Eastern Desert, Egypt): constrains from stable isotopes and whole-rock geochemistry. Arab J Geo 10(19):429 Hamdy MM, El Dien HMG, El-Wahed MAA, Morishita T (2018) Garnierite-bearing serpentinite from the Central Eastern Desert of Egypt: A signature of paleo-weathering in the Arabian Nubian Shield? J Afr Earth Sci 146:95–117 Hamdy MM, Lebda EM (2007) Metamorphism of ultramafic rocks at Gebel Arais and Gebel Malo Grim, Eastern Desert, Egypt: mineralogical and O-H stable isotopes constraints. Egypt J Geol 51:105–124 Harbi H, Eldougdoug A, Al Jahdali NS (2006) Geology and Geochemistry of Jabal Ghadarah Ophiolitic mélange, Zalim Quadrangle, Central Saudi Arabia. J King Abdulaziz University (Earth Sciences) 17:117–153 Harbi H., Hassanen MA, Eldougdoug AA, Sahal MSA (2003) Geological and geochemical studies on some gold occurrences in Zalim Quadrangle, Western Afif terrane, Arabian shield, Kingdom of Saudi Arabia. King Abdul Aziz University. Research Council, Project No. 421/203, 69 p

327 Hargrove U, Stern R, Kimura J-I, Manton W, Johnson P (2006) How juvenile is the Arabian-Nubian Shield? Evidence from Nd isotopes and pre-Neoproterozoic inherited zircon in the Bi’r Umq suture zone, Saudi Arabia. Earth Planet Sci Lett 252(3):308–326 Hariri MM (1989) Gold Mineralization at the Ancient Hajr Mine. Directorate General of Mineral Resources, Jiddah, Kingdom of Saudi Arabia, Jeddah Hariri MM, Makkawi MH (2004) Gold mineralization distributions within rock units at the Hajr gold mine, southwest Saudi Arabia. Arab J Sci Eng 29(2):111–122 Harms U, Darbyshire DPF, Denkler T, Hengst M, Schandelmeier H (1994) Evolution of the Neoproterozoic Delgo suture zone and crustal growth in Northern Sudan: Geochemical and radiogenic isotope constraints. Geol Rundsch 83(3):591–603 Harraz HZ (2000) A genetic model for a mesothermal Au deposit: Evidence from fluid inclusions and stable isotopic studies at El-Sid Gold Mine, Eastern Desert. J Afr Earth Sci 30:267–282 Harris NBW, Gass IG, Hawkesworth CJ (1990) A geochemical approach to allochthonous terranes: a Pan-African case study. Philosophical Transactions of the Royal Society of London. Series A, Mathematical and Physical Sciences, 331(1620):533–548 Hassaan MM, El Mezayen AM, Dardir AA, Hassanein MM (1996) Primary distribution pattern of gold and associated elements in Abu Marawat mine, North Eastern Desert and significance to exploration. Al-Azhar Bulletin of Science 7(1):995–1016 Hussein AA (1990) Mineral deposits of Egypt. In: Said R (ed) The Geology of Egypt, pp 511–566 Hussein IM, Kröner A, Reischmann T (2004) The Wadi Onib mafic-ultramafic complex: a Neoproterozoic supra-subduction zone ophiolite in the northern Red Sea hills of the Sudan. Developments in Precam Geol 13:163–206 Janecky DR, Seyfried WE (1986) Hydrothermal serpentinization of peridotite within the oceanic crust: Experimental investigations of mineralogy and major element chemistry. Geochim Cosmochim Acta 50:1357–1378 Johannes W (1969) An experimental investigation of the system MgO– SiO2–H2O–CO2. Amer J Sci 267:1083–1104 Johannes W (1970) Zur Entstehung von Magnesitvorkommen. Neues Jb Miner Abh 113:274–325 Johnson LE, Fryer P (1990) The first evidence for MORB-like lavas from the outer Mariana fore-arc: geochemistry, petrography and tectonic implications. Earth Planet Sci Lett 100(1–3):304–316 Johnson PR, Andresen A, Collins A, Fowler A, Fritz H, Ghebreab W, Kusky T, Stern RJ (2011) Late Cryogenian-Ediacaran history of the Arabian-Nubian Shield: a review of depositional, plutonic, structural, and tectonic events in the closing stages of the northern East African Orogen. J Afr Earth Sci 61(3):167–232 Johnson PR, Kattan F (2001) Oblique sinistral transpression in the Arabian shield: the timing and kinematics of a Neoproterozoic suture zone. Precam Res 107(1–2):117–138 Johnson PR, Kattan FH, Al-Saleh AM (2004) Neoproterozoic ophiolites in the Arabian Shield: Field relations and structure. Devel Precam Geol 13:129–162 Johnson PR, Scheibner E, Smith EA (1987) Basement fragments, accreted tectonostratigraphic terranes, and overlap sequences: elements in the tectonic evolution of the Arabian Shield. Terrane Accretion and Orogenic Belts 19:323–343 Johnson PR, Woldehaimanot B (2003) Development of the Arabian-Nubian Shield: Perspectives on accretion and deformation in the East African Orogen and the assembly of Gondwana. In: M. Yoshida, B.F. Windley and S. Dasgupta (Editors), Proterozoic East Gondwana: Supercontinent Assembly and Breakup. Geol Soc London Spec Pub: 289–325 Katz O, Beyth M, Miller N, Stern R, Avigad D, Basu A, Anbar A (2004) A Late Neoproterozoic (630 Ma) boninitic suite from

328 southern Israel: implications for the consolidation of Gondwanaland. Earth Planet Sci Lett 218:475–490 Khalaf IM, Oweiss KA (1993) Gold prospection in the environs of Sukari gold mine, central Eastern Desert, Egypt. Annals of the Geological Survey of Egypt 19:97–108 Khalil AES (2000) Petrological and geochemical studies on the serpentinites of Umm Saneyat, El Deigheimi and El Muweilih areas, Qift-Quseir Road, Eastern Desert, Egypt. PhD thesis, Al Azhar University, Cairo, 308 p Khalil AES, Azer MK (2007) Supra-subduction affinity in the Neoproterozoic serpentinites in the Eastern Desert, Egypt: Evidence from mineral composition. J Afr Earth Sci 49:136–152 Khalil AES, Obeid MA, Azer MK (2014) Serpentinized peridotites at the north part of Wadi Allaqi district (Egypt): Implications for the tectono-magmatic evolution of fore-arc crust. Acta Geol Sin 88 (5):1421–1436 Khedr MZ, Arai S (2017) Peridotite-chromitite complexes in the Eastern Desert of Egypt: Insight into Neoproterozoic sub-arc mantle processes. Gond Res 52:59–79 Khedr MZ, Arai S (2013) Origin of Neoproterozoic ophiolitic peridotites in south Eastern Desert, Egypt, constrained from primary mantle mineral chemistry. Mineral Petrol 107(5):807–828 Khedr MZ, Arai S (2016) Petrology of a Neoproterozoic Alaskan-type complex from the Eastern Desert of Egypt: implications for mantle heterogeneity. Lithos 263:15–32 Khudeir AA (1995) El-Genina El-Gharbia and El-Genina El-Sharkia ultramafic-mafic intrusions, Eastern Desert, Egypt: geology, petrology, geochemistry and petrogenesis. Bull Fac Sci Assiut Univ, 177–219 Khudeir AA, El Haddad MA, Leake BE (1992) Compositional variation in chromite from the Eastern Desert, Egypt. Mineralogical Magazine 56(385):567–574 Klemm D, Klemm R (2013) Gold and gold mining in ancient Egypt and Nubia: Berlin, Springer, 649 p Klemm D, Klemm R, Murr A (2001) Gold of the Pharaohs–6000 years of gold mining in Egypt and Nubia. J Afr Earth Sci 33(3–4):643– 659 Kröner A, Linnebacher P, Stern R, Reischmann T, Manton W, Hussein I (1991) Evolution of Pan-African island arc assemblages in the southern Red Sea Hills, Sudan, and in southwestern Arabia as exemplified by geochemistry and geochronology. Precamb Res 53 (1–2):99–118 Kröner A, Todt W, Hussein IM, Mansour M, Rashwan AA (1992) Dating of late Proterozoic ophiolites in Egypt and Sudan using the single grain zircon evaporation technique. Precamb Res 59:15–32 Kusky TM, Abdelsalam M, Tucker R, Stern RJ (2003) Evolution of the East African and related Orogens, and the assembly of Gondwana. Precam Res 123:81–344 Kusky TM, Ramadan TM (2002) Structural controls on Neoproterozoic mineralization in the South Eastern Desert, Egypt: An integrated field, Landsat TM, and SIR-C/X SAR approach. J Afr Earth Sci 35:107–121 Labbé JF (1984) A preliminary survey of the Bir Tawilah gold and tungsten-tin prospect. Saudi Arabian Deputy Ministry for Mineral Resources. Open-File-Report BRGM-OF-04-21 Lago BL, Rabinowicz M, Nicolas A (1982) Podiform chromite ore bodies: a genetic model. J Petrol 23:103–125 Lang JR, Baker T (2001) Intrusion-related gold systems: the present level of understanding. Miner Deposita 36(6):477–489 Le Metour J, Johan V, Tegyey M (1982) Relationships between ultramafic-mafic complexes and volcanosedimentary rocks in the Precambrian Arabian Shield. Deputy Ministry for Mineral Resources Open-File Report BRGM-OF-12–15, p 90 Ledru P, Auge T (1984) The Al Ays ophiolitic complex; petrology and structural evolution. Saudi Arabian Deputy Ministry for Mineral Resources Open-File Report BRGM-OF-04-15

M. G. Shahien et al. Loizenbauer J, Wallbrecher E, Fritz H, Neumayr P, Khudeir AA, Kloetzli U (2001) Structural geology, simple zircon ages and fluid inclusion studies of the Meatiq metamorphic core complex: Implications for Neoproterozoic tectonics in the Eastern Desert of Egypt. Precambrian Res 110:357–383 Madbouly MI (1991) Petrology and geochemistry of some mafic ultramafic rocks of Sinai, Egypt. M.Sc. Thesis, Cairo Univ, 132 pp Martin B, Fyfe WS (1970) Some experimental and theoretical observations on the kinetics of hydration reactions with particular reference to serpentinization. Chem Geol 6:185–202 McElduff B, Stumpfl EF (1991) The chromite deposits of the Troodos complex, Cyprus—evidence for the role of a fluid phase accompanying chromite formation. Miner Deposita 26(4):307–318 Mehanna A, Wetait M, El-Amawy M, Solimani F, Ghabour Y (2004) Petrogenesis and metamorphism of the basement rocks of Imlig area, southwest Sinai, Egypt. Ann Geol Surv Egypt 27:35–59 Mellini M, Rumori C, Viti C (2005) Hydrothermally reset magmatic spinels in retrograde serpentinites: formation of ‘‘ferritchromit’’ rims and chlorite aureoles. Contrib Mineral Petrol 149:266–275 Mercier J-C, Nicolas A (1975) Textures and fabrics of upper-mantle peridotites as illustrated by xenoliths from basalts. J Petrol 16 (2):454–487 Mével C (2003) Serpentinization of abyssal peridotite at mid-oceanridges. Comptes Rendus Geosciences 335:825–852 Mogahed MM (2020) Genesis and Tectonic Implications of Kabr El-Bonaya Ultramafic Rocks, Sinai Peninsula, Egypt: Constraints from Mineralogical and Geochemical Characteristics. Acta Geologica Sinica-English Edition. https://doi.org/10.1111/1755-6724.14406 Moody JB (1976) Serpentinization: A review. Lithos 9:125–138 Morimoto N, Fabries J, Ferguson AK, Ginzburg IV, Ross M, Seifert FA, Zussman J (1988) Nomenclature of pyroxenes. Mineral Mag 52:535–550 Moussa HE (2002) Mineral chemistry and geochemistry of some mafic-ultramafic intrusions in the South Eastern Desert and Sinai, Egypt. Egypt J Geol 46:213–238 Murck BW, Campbell IH (1986) The effects of temperature, oxygen fugacity and melt composition on the behaviour of chromium in basic and ultrabasic melts. Geochim Cosmochim Acta 50(9):1871– 1887 Murton B (1989) Tectonic controls on boninite genesis. Geological Society, London 42(1):347–377 Nasr BB, Beniamin NY (2001) Ophiolite sequence of Gerf-Hasium and Sol Hamed-Shiab areas, South Eastern Desert, Egypt. Ann Geol Surv Egypt 24:63–78 Nassief MO, Macdonald R, Gass IG (1984) The Jebel Thurwah upper Proterozoic ophiolite complex, western Saudi Arabia. Geol Soc London 141:537–546 Nehlig P, Genna A, Asfirane F (2002) A review of the Pan-African evolution of the Arabian Shield. GeoArabia 7:103–124 Newton RC, Stern RJ (1990) A Late Precambrian CO2 “Event”. Geological Society of America, Dallas, p 190 Nozaka T (2003) Compositional heterogeneity of olivine in thermally metamorphosed serpentinite from Southwest Japan. Amer Mineral 88:1377–1384 O’Hanley DS (1996) Serpentinites. Oxford University Press, Oxford, p 277 Obeid MA, Khalil AES, Azer MK (2016) Mineralogy, geochemistry and geotectonic significance of the Neoproterozoic ophiolite of Wadi Arais area, south Eastern Desert, Egypt. Inter Geol Rev 58:687–702 Ohara Y, Stern RJ, Ishii T, Yurimoto H, Yamazaki T (2002) Peridotites from the Mariana Trough: first look at the mantle beneath an active back-arc basin. Contrib Mineral Petrol 143(1):1–18 Osman A (1995) The mode of occurrence of gold-bearing listvenite at El Barramiya gold mine, Eastern desert, Egypt. Middle East

12

Neoproterozoic Ophiolites of the Arabian-Nubian Shield

Research Centre. Ain Shams University. Earth Sciences Series 9:93–103 Oweiss KhA, El Naggar AA, Abdel Razik KA, Moselhy N, Ali AB (2001) Gold exploration at Heiani area, South Eastern Desert, Egypt. Ann Geol Surv Egypt XXIV:435–450 Paktunc AD (1990) Origin of podiform chromite deposits by multistage melting, melt segregation and magma mixing in upper mantle. Ore Geol Rev 5:211–222 Pallister JS, Stacey JS, Fischer LB, Premo WR (1987) Precambrian ophiolites of Arabia: U-Pb geochronology, Pb isotopic characteristics, and implications for continental accretion. Saudi Arabian Deputy Ministry for Mineral Resources Open-File Report USGS-OF-07-10. USGS Open-File Report 88-606 Pallister JS, Stacey JS, Fischer LB, Premo WR (1988) Precambrian ophiolites of Arabia: Geologic settings, U-Pb geochronology, Pb-isotope characteristics, and implications for continental accretion. Precam Res 38:1–54 Parkinson IJ, Pearce JA (1998) Peridotites from the Izu–Bonin– Mariana fore-arc (ODP Leg 125): evidence for mantle melting and melt–mantle interaction in a supra-subduction zone setting. J Petrol 39(9):1577–1618 Patchett PJ, Chase CG (2002) Role of transform continental margins in major crustal growth episodes. Geology 30:39–42 Pearce JA, Barker P, Edwards S, Parkinson I, Leat P (2000) Geochemistry and tectonic significance of peridotites from the South Sandwich arc–basin system, South Atlantic. Contrib Mineral Petrol 139(1):36–53 Pitcairn IK, Craw D, Teagle DA (2014) The gold conveyor belt: large-scale gold mobility in an active orogen. Ore Geol Rev 62:129–142 Quick JE (1990) Geology and origin of the late Proterozoic Darb Zubaydah ophiolite, Kingdom of Saudi Arabia. Geol Soc Am Bull 102(8):1007–1020 Quick JE (1991) Late Proterozoic transpression on the Nabitah fault system—implications for the assembly of the Arabian Shield. Precambr Res 53(1–2):119–147 Ramadan TM (1995) Geological and geochemical studies on some basement rocks at Wadi Hodein area, South Eastern Desert, Egypt. Ph.D., Al Azhar University, Cairo, 184 p Ramadan TM (2002) Exploration for gold-bearing listvenites at Um Khasila area, Central Eastern Desert, Egypt. Egyptian J Remote Sensing and Space Sci 5:63–76 Ramadan TM, Kontny A (2004) Mineralogical and structural characterization of alteration zones detected by orbital remote sensing at Shalatein District, SE Desert, Egypt. J Afr Earth Sci 40(1–2):89–99 Ramadan TM, Sadek MF, Abu El Leil I, Salem SM (2005) Um El Touyur El Fuqani gold mineralization, South Eastern Desert, Egypt: using Landsat ETM + imagery. Ann Geol Surv Egypt 28:263–281 Reischmann T, Kröner A (1994) Late proterozoic island arc volcanics from Gebeit, Red Sea Hills, north-east Sudan. Geol Rundsch 83 (3):547–563 Rino S, Kon Y, Sato W, Maruyama S, Santosh M, Zhao D (2008) The Grenvillian and Pan-African orogens: World’s largest orogenies through geologic time, and their implications on the origin of superplume. Gond Res 14:51–72 Sadek MF, Ramadan TM, El Leil IA, Salem SM (2006) Using remote sensing technique in lithological discrimination and detection of gold-bearing alteration zones at Wadi Defeit area, southeastern desert, Egypt. In Remote sensing for environmental monitoring, GIS applications, and geology VI (Vol. 6366, p. 63660H). International Society for Optics and Photonics Saleh GM (2006) The chromite deposits associated with ophiolite complexes, southeastern Desert, Egypt: Petrological and geochemical characteristics and mineralization. Chin J Geochem 25(4):307– 317

329 Saumur BM, Hattori K (2013) Zoned Cr-spinel and ferritchromite alteration in fore-arc mantle serpentinites of the Rio San Juan complex, Dominican Republic. Mineral Mag 77:117–136 Schandelmeier H, Wipfler E, Küster D, Sultan M, Becker R, Stern RJ, Abdelsalam MG (1994) Atmur-Delgo suture: A Neoproterozoic oceanic basin extending into the interior of northeast Africa. Geology 22(6):563–566 Shackleton RM (1994) Review of late Proterozoic sutures, ophiolitic mélanges and tectonics of eastern Egypt and northeastern Sudan. Geol Rundsch 83:537–546 Shackleton RM (1996) The final collision zone between East and West Gondwana: where is it? J Afr Earth Sci 23(3):271–287 Shanti M (1983) The Jabal Ess ophiolite complex. Bulletin of the Faculty of Earth Sciences, King Abdulaziz University, Jiddah 6:289–317 Shimron AE (1980) Proterozoic island arc volcanism and sedimentation in Sinai. Precambr Res 12:437–458 Sobolev NV, Logvinova AM (2005) Significance of accessory chrome spinels in identifying serpentinite paragenesis. Inter Geol Rev 47:58–64 Stacey JS, Stoeser DB, Greenwood WR, Fischer LB (1984) U-Pb zircon geochronology and geological evolution of the Halaban–AI Amar region of the Eastern Arabian Shield, Kingdom of Saudi Arabia. J Geol Soc London 141:1043–1055 Stern RJ (1994) Arc assembly and continental collision in the Neoproterozoic East African Orogen: implications for the assembly of Gondwanaland. Ann Rev Earth Planet Sci 22:319–351 Stern RJ (2002) Crustal Evolution in the East African Orogen: A Neodymium Isotopic Perspective. J Afr Earth Sci 34:109–117 Stern RJ (2004) Subduction Initiation: Spontaneous and Induced. Earth Planet Sci Lett 226:275–292 Stern RJ, Gwinn CJ (1990) Origin of Late Precambrian Intrusive Carbonates, Eastern Desert of Egypt and Sudan: C, O, and Sr Isotopic Evidence. Precamb Res 46:259–272 Stern RJ, Kröner A, Rashwan AA (1991) A Late Precambrian (*710 Ma) high volcanicity rift in the southern Eastern Desert of Egypt. Geol Rundsch 80:155–170 Stern RJ, Johnson PR, Kröner A, Yibas B (2004) Neoproterozoic ophiolites of the Arabian-Nubian Shield. In: Kusky TM (ed), Precambrian Ophiolites and Related Rocks. In: Developments in Precambrian Geology 13:95–128 Stern RJ, Reagan M, Ishizuka O, Ohara Y, Whattam S (2012) To understand subduction initiation, study forearc crust; to understand forearc crust, study ophiolites. Lithosphere 4:469–483 Stoeser DB, Camp VE (1985) Pan-African microplate accretion of the Arabian Shield. Geol Soc Am Bull 96(7):817–826 Suita MTF, Strieder AJ (1996) Cr-spinel from Brazilian mafic-ultramafic complexes: metamorphic modifications. Int Geol Rev 38:245–267 Surour AA (2019) Mid-ocean ridge vs. fore-arc and subduction settings: Clues from rodingitization of tectonic fragments in the Neoproterozoic ophiolites of the Eastern Desert, Egypt. Lithos 342:18–30 Takla MA, Basta FF, Surour AA (1992) Petrology and mineral chemistry of rodingites associating the pan-African ultramafics of Sikait-Abu Rusheid area, south Eastern Desert, Egypt. Geology of the Arab World 1:491–507 Teklay M (2006) Neoproterozoic arc-back-arc system analog to modern arc-back-arc systems: evidence from tholeiite-boninite association, serpentinite mudflows and across-arc geochemical trends in Eritrea, southern Arabian-Nubian shield. Precam Res 145:81–92 Thalhammer OAR, Prochaska W, Mühlhans HW (1990) Solid inclusions in chrome-spinels and platinum group element concentrations from the Hochgrössen and Kraubath ultramafic massifs (Austria). Contrib Mineral Petrol 105(1):66–80

330 Thekair ME (1975) Carbonate rocks in Al-Amar Asihailiya district of Saudi Arabia. Unpub. M.S. thesis, Jeddah, Saudi Arabia, Inst. Applied Geology, 211 p Uysal İ, Ersoy EY, Karslı O, Dilek Y, Sadıklar MB, Ottley CJ, Tiepolo M, Meisel T (2012) Coexistence of abyssal and ultra-depleted SSZ type mantle peridotites in a Neo-Tethyan Ophiolite in SW Turkey: Constraints from mineral composition, whole-rock geochemistry (major–trace–REE–PGE), and Re–Os isotope systematics. Lithos 132:50–69 Uysal İ, Ersoy EY, Dilek Y, Kapsiotis A, Sarıfakıoğlu E (2016) Multiple episodes of partial melting, depletion, metasomatism and enrichment processes recorded in the heterogeneous upper mantle sequence of the Neotethyan Eldivan ophiolite, Turkey. Lithos 246:228–245 Uysal İ, Şen AD, Ersoy EY, Dilek Y, Saka S, Zaccarini F, Escayola M, Karslı O (2014) Geochemical make-up of oceanic peridotites from NW Turkey and the multi-stage melting history of the Tethyan upper mantle. Mineral Petrol 108(1):49–69 Vail JR (1985) Pan-African (late Precambrian) tectonic terrains and the reconstruction of the Arabian-Nubian Shield. Geology 13(12):839– 842 Wolde B, Asres Z, Desta Z, Gonzalez JJ (1993) Neoproterozoic zirconium-depleted boninite and tholeiitic series rocks from Adola, southern Ethiopia. Precam Res 80:261–279

M. G. Shahien et al. Worku H (1996) Geodynamic development of the Adola Belt (southern Ethiopia) in the Neoproterozoic and its control on gold mineralization. Ph.D. dissertation, Technical University, Berlin, 156 Yibas B, Reimold WU, Anhaeusser CR, Koeberl C (2003) Geochemistry of the mafic rocks of the ophiolitic fold and thrust belts of southern Ethiopia: constraints on the tectonic regime during the Neoproterozoic (900–700 Ma). Precam Res 121:157–183 Zhou MF, Robinson PT, Malpas J (2005) REE and PGE Geochemical Constraints on the Formation of Dunites in the Luobusa Ophiolite, Southern Tibet. J Petrol 46(3):615–639 Zhou M-F, Robinson PT, Bai W-J (1994) Formation of podiform chromitites by melt/rock interaction in the upper mantle. Miner Depos 29:98–101 Zhou M-F, Robinson PT, Malpas J, Li Z (1996) Podiform chromitites in the Luobusa ophiolite (southern Tibet): Implications for melt-rock interaction and chromite segregation in upper mantle. J Petrol 37:3–21 Zimmer M, Kröner A, Jochum KP, Reischmann T, Todt W (1995) The Gabal Gerf complex: a Precambrian N-MORB ophiolite in the Nubian Shield, NE Africa. Chem Geol 123:29–51 Zoheir B, Lehmann B (2011) Listvenite-lode association at Barramiya gold mine, Eastern Desert. Egypt. Ore Geol Rev 39:101–115

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Ghadir Ophiolites, Eastern Desert, Egypt: A Complete Sequence of Oceanic Crust in the Arabian-Nubian Shield Gaafar A. El Bahariya

Abstract

Keywords

Neoproterozoic ophiolites are widely distributed in the Central and Southern Eastern Desert of Egypt. The most important and well-known ophiolites of the Central Eastern Desert (CED) are exposed in Ghadir area. They occur as intact ophiolitic sequence of oceanic crust, blocks and fragments in an ophiolitic mélange and as dismembered ophiolite components. The whole rock association is foliated and metamorphosed to the greenschist facies up to lower amphibolite facies. The intact ophiolite sequence is composed mainly of metagabbro, sheeted dykes, and pillow lavas. Exotic fragments of oceanic lithosphere tectonically incorporated in a foliated volcaniclastic matrix forming tectonic ophiolitic mélange. The Ghadir pillowed metabasalts are classified as high-Ti MORB-type, which are comparable to MORB or BABB ophiolites and appear to be derived from an evolved magma in an intra-oceanic back-arc basin. The intact ophiolite sequence is remnants of oceanic crust formed the floor of the basin and became tectonically obducted during the closure of the back-arc basin in which they formed. The formation of the Ghadir ophiolitic mélange by tectonic processes in a back-arc basin is followed by tectonic emplacement of some dismembered ophiolitic rocks along structural contacts. The whole MORB ophiolites of the Eastern Desert (ED) of Egypt were suggested to be formed in a back-arc basin due to collision of an island arc system with the African continent. They occur in different geological settings, and the compositional variation within these rocks was assigned to geochemical variations in a single tectonic setting of back-arc basin rather than to different tectonic settings.

Neoproterozoic Ghadir Ophiolites Eastern desert Egypt MORB BABB Ophiolitic melange

G. A. El Bahariya (&) Geology Department, Faculty of Science, Tanta University, Tanta, Egypt e-mail: [email protected]




13.1















Introduction

Neoproterozoic ophiolites are abundant in the ArabianNubian Shield (ANS) and range in age from 690 to 890 Ma (Stern et al. 2004), where they commonly occur along suture zones (Stern 1994; 2008). Deciphering the tectonic settings of the ophiolites in the ANS is an important step for understanding the evolution of the Pan-African orogeny. The ophiolite complexes are classified into mid-ocean ridge basalts (MORB) ophiolites and Supra-Subduction zone (SSZ) ophiolites (Pearce et al. 1984; Dilek et al. 2008). The MORB ophiolites are subdivided into E-MORB, P-MORB, N-MORB, and C (Contaminated)-MORB types (Pearce 2008). The Precambrian rocks of the ED of Egypt represent the northern part of the Nubian Shield. Ophiolites of Egypt are commonly dismembered and restricted to the CED and southern (SED) Eastern Desert (Shackleton et al. 1980; Ries et al. 1983; Abdel-Karim et al. 2016; El Bahariya 2019; 2021). Several types of ophiolitic melanges were documented in the CED of Egypt, where the ophiolitic exotic blocks tectonically or/and sedimentary intermixed with a sheared and foliated matrix giving rise to tectonic mélange, olistostrome, and olistostromal mélange (El-Bahariya 2008b, 2012). Recently, El-Bahariya (2018) classified the ophiolites of the ED into: intact MORB-type ophiolites, dismembered ophiolites and arc-associated ophiolites. The ophiolitic rocks include serpentinized peridotites, quartz carbonates, talc carbonates, metagabbros, and pillow metabasalts, together with scarce sheeted dykes, pelagic sediments, and trondhjemite.

© The Author(s), under exclusive license to Springer Nature Switzerland AG 2021 Z. Hamimi et al. (eds.), The Geology of the Arabian-Nubian Shield, Regional Geology Reviews, https://doi.org/10.1007/978-3-030-72995-0_13

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The best preserved exposures of the Neoproterozoic ophiolites occurred in the Wadi Ghadir area, where an ophiolite sequence was described for the first time by El Sharkawy and El Bayoumi (1979). Later, further studies on the Ghadir ophiolites have been done by many workers (e.g., El Bayoumi 1980; Kroner et al. 1992; El-Bahariya 2007; Abd El-Rahman et al. 2009; Basta et al. 2011). In the present chapter, the author presents a new field geological information, together with a compilation of geochemical data. The main objectives are: to clarify the geological setting and mode of occurrences of the Ghadir ophiolites; to put some light on the tectonic setting and petrogenesis of these oceanic crustal rocks; to compare the Ghadir ophiolite with other ophiolites elsewhere in the ED of Egypt; and to present a geodynamic model for the evolution of the ophiolites in Ghadir area.

13.2

Geological Setting

The Wadi Ghadir area is located in the southern part of the CED. It lies approximately 30 km southwest Marsa Alam town and extends between latitudes 24° 37″ to 24° 57″ N and longitudes 34° 39″ to 35° 05″ E (Fig. 13.1). It is covered mainly by a Neoproterozoic ophiolites, ophiolitic melanges and metavolcanics and volcaniclastic metasediments, where all these rock units were intruded by granitic intrusions. Based on the geological setting, field observations and the

classification of ophiolites given by El-Bahariya (2018), the ophiolitic rocks in the Ghadir area occur as: (i) intact ophiolite sequence, (ii) ophiolitic mélange, and (iii) dismembered ophiolites along tectonic contacts rather than melanges.

13.2.1 Ghadir Ophiolite Sequence The Ghadir ophiolite sequence occurs along Wadi El Beda and Wadi Saudi to the northwest of Gabal Dob Nia (Fig. 13.1). This sequence is composed of massive and layered metagabbros, sheeted dykes, and pillow lavas, where the pillow lavas are tectonically overlain by ophiolitic mélange (Fig. 13.2). Serpentinites and metamorphosed ultramafics are not observed at the base of the section, instead they occur as fragments and blocks in melanges or along tectonic contacts. Layered metagabbros showing cumulate texture and individual layers reaches up to 50 cm thick, which are often disrupted and cannot be traced laterally for more than 20 m (Fig. 13.3a). The massive gabbroic rocks have a mottled appearance of rosette structure, which grades downwards into layered gabbro with foliated fabrics. Occasionally, the gabbroic rocks enclose large pods of appinitic gabbro, anorthosite, and trondhjemite. The layered metagabbros of Wadi El Beda followed by sheeted dykes, which in turn they overlain and followed by the pillow lavas (Fig. 13.2). The contacts of the sheeted dykes with the

Fig. 13.1 a Location map of the Ghadir area; b geological map of the Wadi Ghadir area (after El-Bayoumi 1980)

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Fig. 13.2 General view showing the intact ophiolitic sequence of Wadi El Beda (looking NE) overlained by tectonic ophiolitic mélange

metagabbros are commonly sharp and occasionally the dykes invade metagabbros as channel ways (Fig. 13.3b), while their contacts with the pillow lavas are sapped and the pillows become less distinct. They are medium to fine grained, range in thickness from 1 to 10 m and trending in a N–E and NE–SW or NW–SE direction (Fig. 13.3c). Pillow lavas are exposed at the northeastern end of Wadi El Beda and the eastern end of Wadi Saudi. The pillow lavas in Wadi Saudi are less vesicular and extend laterally for about 70 m. The pillowed metabasalts of Wadi El Beda reach about 200 m thick and are tectonically overlain by ophiolitic mélange with prominent shearing along the contact (Fig. 13.2). The pillows are closely packed bulbous or oval in shape and decreasing in size from base to top. They vary in size from 20 cm to 1.5 m in diameter and composed mainly of pillow lavas (Fig. 13.4a, b). Isolated and broken-pillow breccias, and massive individual pillows were observed among the whole pillowed mass (Fig. 13.4c). Pillows possess convex upper surface, microcrystalline core and vesicular margins with a schistose sheath (2–8 cm thick) encloses each pillow (Fig. 13.4d, e). Amygdales increase in abundance from the interior to the margin of the pillows, vary in shape from circular to ellipsoidal, and rarely irregular. Microscopically, metadolerites of sheeted dykes are medium grained, has intergranular texture as well as blastophitic and subophitic textures. They are composed of triangles of microtabular plagioclase crystals with the interstices filled by hornblende, actinolite, chlorite, and opaques (Fig. 13.5a). The pillowed metabasalts are composed mainly

of phenocrysts and micro-phenocrysts of tabular plagioclase and/or rarely prismatic hornblende, actinolite and actinolitic hornblende, all set in an intergranular, intersertal and variolitic groundmass of plagioclase, hornblende, biotite, iron oxides, chlorite, relics of pyroxene and calcite (Fig. 13.5b–e). The plagioclase in the groundmass occurs as laths, whereas hornblende and actinolite occur as fine prismatic crystals. Ovoidal amygdales filled by calcite, chlorite, quartz, and epidote are commonly occur (Fig. 13.5f). The mineral assemblage constitutes these rocks indicates a greenschist facies to lower amphibolite facies metamorphism.

13.2.2 Ghadir Ophiolitic Mélange The Ghadir ophiolitic mélange forms mappable rock unit overlying the ophiolite sequence and is composed of different exotic fragments and bocks of variable sizes and shapes tectonically intermixed with a foliated and sheared volcaniclastic matrix. The Ghadir ophiolitic mélange is concentrated mainly around Wadi El Beda, where it is tectonically overly the pillow lavas (Fig. 13.2). It is comparable to the tectonic mélange of El-Bahariya (2012). The matrix of the mélange is composed mainly of metapryroclastics and volcaniclastic metasediments. It is foliated and is consists of lapilli lithic and crystal metatuffs, lithic and crystal metatuff and ashes, together with volcaniclastic metagreywackes and minor schists. The blocks and fragments are tectonically incorporated in the matrix and include serpentinites, talc

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Fig. 13.3 a Layered metagabbro with sinistral shearing; b metagabbro (MG) merge into sheeted dykes (SD) with sharp contact (Looking E); c sheeted dykes with irregular inclusion of metagabbro (looking E)

carbonates, quartz carbonate, siliceous rocks, pillow lavas, marble, cherts, and granites (Fig. 13.6). Serpentinites occur as sheets, blocks, and fragments ranging from few centimeters to mountain size tectonically set in a volcaniclastic matrix (Fig. 13.6a, b). Blocks of deformed pillow lavas are incorporated within the matrix of the mélange (Fig. 13.6b, c). Quartz carbonate fragments and blocks (Fig. 13.6d, e) together with fragments and disrupted beds of chert (Fig. 13.6f) are also inserted within the matrix of the mélange. The Ghadir tectonic melange is highly tectonized under low-grade greenschist facies metamorphism. This is indicated by the presence of shear zone extended to the west towards the high grade Hafafit migmatitic association (e.g. El Bahariya and Abd El-Wahed 2003; El Bahariya 2008a).

13.2.3 Dismembered Ophiolites To the W and NW of the mélange, dismembered ophiolite rocks of serpentinites and metagabbro occur as tectonic sheets and blocks along structural contacts. The elongated massive serpentinized peridotite sheet of Gabal Lewiwi

trends NE-SW and composed of serpentinized harzburgite, and wehrlite, and partially altered to talc-carbonate rocks (e.g., Basta et al. 2011). Serpentinized ultramafic rocks occur at the junction between Wadi Ghadir and Wadi Lawi are commonly sheared, foliated and in part surrounded by talc carbonate carapace.

13.3

Geochemistry and Tectonic Setting

13.3.1 Geochemical Characteristics Geochemistry and geochemical characteristics of ophiolites in the Ghadir area were investigated by many workers (e.g., Kroner et al. 1992; El-Bahariya 2007; Abd El Rahman et al. 2009; Abdel-Karim and Ahmed 2010; Basta et al. 2011). Gabbros of Ghadir ophiolite are tholeiitic, have slight enrichment of REE, and display a large variation in Zr and Nb. TiO2, Fe2O3, and MgO, but A2O3, CaO, and Na2O are moderately variable (e.g. Abd El Rahman et al. 2009). Chondrite-normalized REE patterns of the gabbros with the highest Mg-number show convex upward, light rare earth

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Fig. 13.4 a, b Fining upward bulbous or oval pillow lavas (looking N); c isolated and fragmented pillow lavas; d pillow lava possess convex upper surface with peripheral zone rich in vesicles; e pillow lava with foliated sheath

element (LREE)-depleted patterns, while those with low Mg-number show flat REE patterns and then slightly enriched in LREE (Fig. 13.7a). The less evolved sample has a

positive Eu anomaly which decreases with an increase in the overall abundance of REE. The gabbroic rocks are slightly enriched in LILE, up to 10  MORB values, and the less

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Fig. 13.5 Photomicrographs showing the petrographic features of metadolerite and metabasalts: a metadolerite with intergranular or blastophitic and subophitic textures; b Porphyritic metabasalt with plagioclase phenocryst; c metabasalt with variolitic texture; d, e metabasalt with intersertal texture and microphenocrysts of plagioclase and hornblende; f amygdaloidal metabasalts with intergranular texture and amygdales filled with calcite

evolved samples are slightly depleted in HFSE and REE compared to MORB. Moreover, values of Sc and Cr are close to MORB values. The dykes display tholeiitic, transitional, and calcalkaline characters (Abd El Rahman et al. 2009). Their

Mg-numbers are generally higher than, those of the pillow lavas and are slightly depleted in HFSE and REE relative to MORB, while some other have HFSE contents close to MORB values. However, samples from the sheeted dykes have higher Ti/V ratios similar to the intra-plate basalts

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Fig. 13.6 Field observations of the Ghadir ophiolitic mélange: a serpentinite sheet (Sp) tectonically emplaced over the volcaniclastic metasediments (VMS) (looking SE); b blocks of serpentinites (Sp) and pillow lavas (PV) in a volcaniclastic matrix (VMS); c Blocks of pillow lavas (PV) in sheared matrix; d, e fragments of quartz carbonate (Qz) in sheared volcaniclastic matrix; f block of deformed chert (Ch) within the mélange matrix

(Basta et al. 2011). Chondrite-normalized REE patterns of tholeiitic dykes have both LREE depleted chondritenormalized patterns and slightly LREE-enriched patterns. REE patterns of some dikes are similar to those of the amygdaloidal pillow lavas, e.g., (Abd El Rahman et al. 2009). Chondrite-normalized REE patterns of transitional

series dikes are enriched in LREE, which are similar to those of the porphyritic pillow lavas, and they also exhibit negative Eu anomalies. Chondrite-normalized REE patterns of Calc-alkaline dikes show strong LREE enrichment. The rocks of the pillowed lavas are commonly of basaltic composition, have subalkaline nature and tholeiitic character

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Fig. 13.7 REE patterns of Ghadir and Gerf ophiolites (from Zimmer et al. 1995; Abd El Rahman et al. 2009)

(e.g., El-Bahariya 2007). They have high contents of total FeO, CaO, Na2O, TiO2, and Zr. TiO2 and low contents of MgO. Zr values range from 0.74 to 3.64 wt% and from 27 to 448 ppm, respectively. The HFS elements display very close normalized values, which are either slightly higher or lower than MORB values. The REE pattern display slight enrichment in LREE (e.g., Abd El Rahman et al. 2009). Commonly, pillowed rocks are enriched in LILE, HFSE, and REE relative to MORB, but some rocks show negative Eu anomalies (Eu/Eu⁎ = 0.61–0.65) (Fig. 13.7b).

13.3.2 Ophiolitic Affinity and Tectonic Setting The tectonic setting of the MORB basalts are characterized by Ti/V ratios ranging from 20 to 50 (e.g., Shervais 1982). On the Ti–V variation diagram, Ghadir pillowed metabasalts have Ti/V ratios in the range of 35–52 and accordingly plot in the high-Ti part of the OFB field (Fig. 13.8a). They are closely comparable with that of the present-day ocean floor basalts, and their Nb/La ppm ratios are similar to that of both the Lau and Manus basins (e.g., El-Bahariya 2007; Abd El Rahman et al. 2009).

13.3.3 Comparison with Other Ophiolites The obtained geochemical data of the ophiolitic metavolcanics and pillowed metabasalts of Ghadir area are compared with the data of other MORB ophiolites in the ED such as pillow lavas from Mubarak and Um Esh ophiolitic melanges, Muweilih pillow lavas (e.g., El Bahariy 2018) and Gerf ophiolitic metabasalts (e.g., Zimmer et al. 1995). The analyzed samples from Ghadir ophiolites have MORB normalized trace element patterns characterized by enrichment of LILE relative to HFSE, thus creating negative Nb and Ta anomalies. Generally, the REE of pillowed Ghadir metabasalts and Gerf ophiolites show flat to slightly enriched LREE pattern, which are similar with MORB or back-arc basin (BAB) magmas (Fig. 13.7b–d). The geochemical data indicate that all pillowed metabasalts from Ghadir ophiolite, Mubarak, and Um Esh ophiolitic melanges, and Muweilih pillow lavas have Ti/V ratios in the range 20–50 suggesting MORB or BAB basalts tectonic setting except that the Ghadir metabasalts are relatively enriched in Ti (El-Bahariya 2007; Abd El Rahman et al. 2009; Basta et al. 2011). This reflecting lower degree of partial melting for MORB basalts and more fractionation for Ghadir metabasalts.

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Fig. 13.8 Tectonic discrimination and variation diagrams of the Ghadir ophiolitic metabasalts and other MORB metabasalts of the ED of Egypt: a Ti/1000 versus V diagram of Shervais (1982); b, c TiO2 versus Zr and MgO compositional variation diagrams for the ophiolitic MORB metabasalts from the ED. Fields of Ghadir and other MORB ophiolites from El-Bahariya (2018)

In terms of major—and trace-element compositions, MORB ophiolitic metabasalts from the CED show a compositional variation regarding their Ti, Mg, and Zr contents (Fig. 13.8b, c). Ghadir metabasalts show considerable enrichment in Fe, Ti, and V and strong depletion in MgO, Cr, and Ni compared to other MORB metabasalts suggesting that the Ghadir metabasalts are the most fractionated rocks. Since the Ghadir metabasalts and the Heiani and Harga Zarga metabasalts show high abundances of Ti, they are classified as high-Ti basalts (e.g., Serri 1981; Zimmer et al. 1995; Basta et al. 2011). Generally, the MORB metabasalts of the CED can be divided into high-Ti group and low-Ti

groups (El-Bahariya 2018). Moreover, the Mg contents increase and Ti decrease from the Ghadir ophiolitic metabasalts in the south to the Muweilih metabasalts in the north. This reflects the effects of low-pressure fractional crystallization and the compositional variation of the MORB rocks within the back-arc basin. Marginal basin systems and most back-arc basin produce basalts essentially similar to depleted mid-ocean ridge basalts in their geochemistry (e.g., Saunders et al. 1980; Pearce et al. 1984). However, all MORB metabasalts, although from different localities, are very similar to N-MORB or BABB and very different from SSZ lavas. Thus, the trace elements such as high field strength

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elements (HFSE) and REE also suggest similarities of Ghadir ophiolites with back-arc basin (BAB) magmas (e.g. Sun and Nesbitt 1982; Saunders and Tarney 1984; El-Bahariya 2007; Abdel-Karim and Ahmed 2010).

13.4

Petrogenesis and Tectonic Evolution

Ophiolite rocks in Ghadir area occur as: intact ophiolitic sequence of oceanic crust, tectonic ophiolitic mélange, and dismembered ophiolite components along tectonic contacts rather than melanges. The intact ophiolite sequence is exposed mainly in Wadi El Beda and is extended to Wadi Saudi, where the ophiolitic melange tectonically overly this ophiolitic sequence. The dismembered ophiolites occur as sheets and masses in Wadi Ghadir. Pillow lavas of Ghadir ophiolite are characterized by ellipsoidal to bulbous forms, which may reflect the effect of hydrostatic pressures and low viscosity of the magma in an ancient back-arc basin (e.g., Vuagnat 1975). There is a systematic increase in pillow vesicularity, volatile content and vesicle size with decreasing depth of water at the site of emplacement at the time of eruption (e.g., Jones 1965; Moore 1979; El-Bahariya 2007). Thus, the high vesicularity of the pillows suggests extrusion at shallow-water depths. However, the high vesicularity in some back-arc basin lavas may indicate high volatile content rather than eruption in shallow water environment (e.g., Harper 1982). The MORB metabasalts from ED of Egypt show different concentrations of Ni, Cr, Mg, and Zr controlled by the early crystallizing mineral phases such as olivine and pyroxene. The compositional variation within MORB rocks appears to reflect the effects of low-pressure fractional crystallization (i.e., whereby TiO2 and Al2O3 increased and MgO decreased). Such fractionation process was controlled mainly by the early crystallization of olivine and Cr-spinels followed by clinopyroxene and, later, by plagioclase. In respect to their Ti contents, the ophiolitic metabasalts with MORB affinity can be divided into very low to low-Ti group and high-Ti group, which appear to have originated from rather primitive magmas (with high concentrations of Ni, Cr, and Mg and more evolved melts, respectively. High-Ti ophiolites are considered to belong to Mid-Ocean Ridges, whereas very low-Ti belong to Supra-Subduction zones (Beccaluva et al. 1983; Beccaluva and Serri 1988; El-Bahariya 2018). The major and trace-element chemistry of Ghadir metabasalts displays N-MORB affinities like other MORB ophiolites in the ED such as Muweilih, Beririq, ophiolitic metabasalts in melanges and Gerf ophiolites except that the ophiolitic rocks of Ghadir and Gerf are relatively highly enriched in Ti and classified as Ti-rich ophiolites. Ghadir metabasalts with high Zr/Nb ratios and LREE-enriched pattern are suggested to be akin to

N-MORB derived from fertile N-MORB source or from a depleted mantle source modified by a subduction component. The high abundances of Ti–Fe oxides in Ghadir metabasalts may indicate crystallization from evolved fractionated liquids, where the metabasalts with low Ti may be formed by 20–25% partial melting, followed by extensive shallow depth fractionation to form high Ti Ghadir metabasalts (e.g., El-Bahariya 2007). However, the MORB ophiolites could be derived from variably depleted asthenospheric mantle sources related via incremental partial melting of shallow MORB-type source. MORB ophiolites occur in incipient oceans, major oceans and most back-arc basins (e.g., Hawkins 1980; Pearce et al. 1984). Most ophiolites originate in back-arc basins in the vicinity of island arcs, and their rocks should therefore carry geochemical characteristics of magma types from both environments (Moores 1982). The LREE enrichment in Ghadir metabasalts, the low initial 87Sr/86Sr ratios, together with the low MgO contents of the Ghadir basalts indicate that these rocks are similar to contaminated MORB or Back-Arc Basin Basalt (BABB), which formed during the initial stage of back-arc development (Basta et al. 2011). The serpentinites of the Ghadir ophiolite appear to be derived from ultramafic rocks formed at a mid-ocean ridge followed by their emplacement in a subduction zone setting (Khalil 2007) or are comparable with depleted to highly depleted forearc harzburgite of suprasubduction zone setting (Abdel-Karim et al. 2020). Accordingly, it is proposed that the ophiolitic rocks of Ghadir area were developed in an oceanic back-arc basin setting (e.g., Beccaluva et al. 2004). This back-arc basin origin of the Ghadir ophiolite was suggested by many workers (Kröner et al. 1987, El-Bahariya 2007; Abd El-Rahman et al. 2012; Basta et al. 2011). Ages of the ANS ophiolites are tightly clustered around 870–690 Ma (Stern et al. 2004). Zircon dating of a plagiogranite from Ghadir area revealed an age of 746 ± 19 M (Kröner et al. 1992). Since all MORB rocks of the ED appear to be formed in a back arc basin, the compositional variations of these rocks can be attributed to source heterogeneity beneath back-arc basins that resulted from variable mixing of fertile mantle, melt depleted mantle, and slab-derived components (e.g., Martinez and Taylor 2003). However, the nature of back arc magmatism is a result of mixing of a MORB-like source and limited enrichment in LILE relative to arc magmatism (Taylor and Martinez 2003). Ghadir metabasalts show similarity to both MORB and back-arc basin basalts (BABB). This together with the spatial and temporal occurrence of volcaniclastic metasediments suggest that the whole ophiolitic rocks have been formed in an intra-oceanic back-arc basin, where the volcaniclastic metasediments were mainly derived from a coeval volcanic arc (e.g., El-Bahariya 2007, 2012). The following likely sequence of events is deduced as follows:

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a. Formation of oceanic crust with MORB compositions was followed with incipient opening of oceanic intra-arc or back-arc basins. This contemporaneous with the evolution of the Neoproterozoic intra-oceanic island arc rocks of the ED above an E-dipping subduction zone (e.g., Abd El Rahman et al. 2012). b. Obduction of the oceanic crust onto the continental margin during basin collapse and closure due to accretion of arc-back-arc terrain to a continental margin. Successive compressional phases led to the disaggregation and fragmentation of ophiolites that tectonically incorporated within the volcaniclastic matrix to form ophiolitic mélange in a back-arc basin. The whole ophiolite components and associated arc volcaniclastic metasediments are now preserved in different geological settings as: intact ophiolite sequence mainly of oceanic crust, tectonic ophiolitic melange with exotic blocks of ophiolites and dismembered ophiolite rocks along tectonic contacts.

13.5

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4. Collision of an island arc system with the African continent resulted in the obduction of ocean floor and BAB crust as well as island arc volcanics and volcaniclastic sediments, which are now found in the Ghadir area as remnants of ocean floor tectonically overlain by ophiolitic tectonic mélange. 5. The whole MORB ophiolites of the ED were suggested to be formed in a back-arc basin. Major and trace element variation of these rocks is mainly attributed to variable degree of partial melting, and fractionation processes at variable depths. The different geological settings of MORB ophiolites in the ED reflect variable palaeogeographic settings and the effect of subsequent tectonic events. The compositional variation within these rocks was assigned to physical–chemical variations in a single tectonic setting (back-arc basin) rather than to different tectonic settings, where different ophiolites may be developed at different stages of marginal basin evolution. 6. Further, geochemical and age dating investigations are required to clarify the compositional variations and the nature of ophiolite components of the Ghadir area as one of the known ophiolites in the Eastern Desert of Egypt.

Concluding Remarks

The following concluding remarks can be drawn as follows:

References 1. The Ghadir ophiolites were considered as the most important and well-known ophiolite exposures in the ED of Egypt. They occur as intact ophiolitic sequence of oceanic crust, blocks, and fragments in ophiolitic mélange and as dismembered ophiolite components. The whole rock association is metamorphosed to the greenschist facies up to lower amphibolite facies. The intact ophiolite sequence is composed of metagabbro, sheeted dykes, and pillow lavas, but lacking serpentinized ultramafics. The ophiolitic mélange comprises exotic ophiolite fragments intermixed with sheared and foliated volcaniclastic matrix. 2. The Ghadir metabasalts are classified into high-Ti MORB-type, which are similar to MORB or BABB ophiolites and appear to be derived from an evolved magma in an intra-oceanic back-arc basin. The pillowed metabasalts, sheeted dykes, and metagabbros are remnants of oceanic crust formed the floor of the basin and became tectonically obducted in a subduction zone environment during the closure of the back-arc basin in which they formed. 3. The successive compressional phases of deformation led to the formation of the Ghadir ophiolitic mélange by tectonic incorporation of ophiolitic components into the volcaniclastic matrix in a back-arc basin. This followed by tectonic emplacement of some dismembered ophiolite components along structural contacts.

Abdel-Karim AM, Ahmed Z (2010) Possible origin of the ophiolites of Eastern Desert, Egypt, from geochemical perspectives. Arab J Sci and Eng 35(1 A):115–143 Abdel-Karim AAM, Ali S, Helmy HM, El-Shafei HM (2016) A fore-arc setting of the Gerf ophiolite, Eastern Desert, Egypt: evidence from mineral chemistry and geochemistry of ultramafites. Lithos 263:52–65 Abdel-Karim AM, El-Shafei S, Azer MK (2020) Petrology and geochemistry of some ophiolitic metaperidotites from the Eastern Desert of Egypt: Insights into geodynamic evolution and post-serpentinization metasomatism. Acta Geol Sinica 94(4):1–23 Abd El-Rahman Y, Polat A, Dilek Y, Fryer BJ, El-Sharkawy M, Sakran S (2009) Geochemistry and tectonic evolution of the Neoproterozoic Wadi Ghadir ophiolite, Eastern Desert, Egypt. Lithos 113:158–178 Abd El-Rahman Y, Polat A, Dilek Y, KuskyTM E-SM, Said A (2012) Cryogenian ophiolite tectonics and metallogeny of the central Eastern Desert of Egypt. Int Geol Rev 54:1870–1884 Basta FF, Maurice AE, Bakhit BR, Ali KA, Manton WI (2011) Neoproterozoic contaminated MORB of Wadi Ghadir ophiolite, NE Africa: geochemical and Nd and Sr isotopic constraints. J Afr Earth Sci 59:227–242 Beccaluva L, Di Girolamo P, Macciotta G, Morra V (1983) Magma affinities and fractionation trends in ophiolites. Ofioliti 8:307–324 Beccaluva L, Serri G (1988) Boninitic and low-Ti subductionrelated lavas from intraoceanic arc–backarc systems and low-Ti ophiolites: a reappraisal of their petrogenesis and original tectonic setting. Tectonophysics 146:291–315 Beccaluva L, Coltortia M, Giuntab G, Sienaa F (2004) Tethyan vs. cordilleran ophiolites: a reappraisal of distinctive tectono-magmatic features of suprasubduction complexes in relation to the subduction mode. Tectonophysics 393:163–174

342 Dilek Y, Furnes H, Shallo M (2008) Geochemistry of the Jurassic Mirdita ophiolite (Albania) and the MORB to SSZ evolution of a marginal basin oceanic crust: Lithos, 100, pp 174–209 El Bahariya GA (2007) Geology, compositional variation and petrogenesis of possible MORB-type ophiolitic massive and pillowed metabasalts from the Pan-African belt, Eastern desert, Egypt. Egyptian J Geol 51:41–59 El Bahariya GA (2008a) Geology and petrology of Neoproterozoic syntectonic anatectic migmatites around Wadi Abu Higlig, Hafafit region, Eastern Desert, Egypt. Egypt J Geol 5225–5254 El Bahariya GA (2008b) Geology, mineral chemistry and petrogenesis of Neoproterozoic metamorphosed ophiolitic ultramafics, Central Eastern Desert, Egypt: Implications for the classification and origin of the ophiolitic mélange. Egypt J Geol 52:55–82 El-Bahariya GA (2012) Classification and origin of the Neoproterozoic ophiolitic mélange in the Central Eastern Desert of Egypt. In: Dilek Y, Festa A, Ogawa Y, Pini GA (eds) Chaos and geodynamics: Melanges, Melange forming processes and their significance in the geologic record Special issue Bmélanges Tectonophysics, Part VI ophiolitic mélange, pp 357–370 El Bahariya GA (2018) Classification of the Neoproterozoic ophiolites of the Central Eastern Desert, Egypt based on field geological characteristics and mode of occurrence. Arab J Geosci 11:313 El Bahariya GA (2019) Geochemistry and Tectonic Setting of Neoproterozoic Rocks from the Arabian-Nubian Shield: Emphasis on the Eastern Desert of Egypt, Applied Geochemistry with Case Studies on Geological Formations, Exploration Techniques and Environmental Issues, Luis Felipe Mazadiego, Eduardo De Miguel Garcia, Fernando Barrio-Parra and Miguel Izquierdo-Díaz, IntechOpen, https://doi.org/10.5772/intechopen.82519 El Bahariya GA (2021) The ophiolite-dominated suprastructure, Eastern Desert, Egypt. In: Hamimi Z, Arai S, Fowler A, El-Bialy MZ (eds), The Geology of the Egyptian Nubian Shield, Regional Geology Reviews book series (RGR), Springer Nature Switzerland, pp 161–182. https://doi.org/10.1007/978-3-030-49771-2 El Bahariya GA, Abd El-Wahed MA (2003) Petrology, mineral chemistry and tectonic evolution of the northern part of Wadi Hafafit area, Eastern Desert of Egypt. In: 3rd international conference on the geology of Africa. Assiut Univ 2:201–231 El Bayoumi RM (1980) Ophiolites and associated rocks of Wadi Ghadir, east of Gabal Zabara, Eastern Desert, Egypt. Unpublished Ph.D. Thesis, Univ. Cairo, 227 p El Sharkawy MA, El Bayoumi RM (1979) The ophiolites of Wadi Ghadir area, Eastern Desert, Egypt. Ann Geol Surv 9:125–135 Harper GD (1982) Inferred high primary volatile contents in lavas erupted in an ancient back-arc basin, California. J Geol 90:187–194 Hawkins JW (1980) Petrology of back-arc basins and island arcs; their possible role in the origin of ophiolites. In: Panayiotou A (ed), Ophiolites-Proceedings of the International Ophiolite Symposium, Cyprus, 244–254 Jones JG (1965) Pillow lavas as depth indicators. American J Sci 267:181–195 Khalil KI (2007) Chromite mineralization in ultramafic rocks of the Wadi Ghadir area, Eastern Desert, Egypt: mineralogical, microchemical and genetic study. Neues Jahrbuch für MineralogieAbhandlungen 183:283–296 Kröner A, Greiling R, Reischmann T, Hussein IM, Stern RJ, Durr S, Kruger J, Zimmer M (1987) Pan-African crustal evolution in the

G. A. El Bahariya Nubian segment of northern Africa. Am Geophy Union Geodynamics Ser 17:235–257 Kröner A, Todt W, Hussein IM, Mansour M, Rashwan AA (1992) Dating of late Proterozoic ophiolites in Egypt and Sudan using the single grain zircon evaporation techniques. Precambr Res 59:15–32 Martinez F, Taylor B (2003) Controls on back-arc crustal accretion: insights from the Lau, Manus and Mariana basins. In: Larter RD, Leat ET (eds) Intra-oceanic Subduction Systems: Tectonic and Magmatic Processes. Geological Society, London, Special Publication, 219, pp 19–54 Moore JG (1979) Vesicularity and CO2 in mid-ocean ridge basalts. Nature 282:250–253 Moores EM (1982) Origin and emplacement of ophiolites. Rev Geophys Space Phys 20:735–760 Pearce JA, Lippard SJ, Roberts S (1984) Characteristics and tectonic significance of suprasubduction zone ophiolites. In: Kokelaar BP, Howell MF (eds) Marginal Basin Geology. Geological Society of London, pp 77–94 Pearce JA (2008) Geochemical fingerprinting of oceanic basalts with applications to ophiolite classification and the search for Archean oceanic crust. Lithos 100:14–48 Ries AC, Shackleton RM, Graham RH, Fitches WR (1983) Pan-African structures, ophiolites and melange in the Eastern Desert of Egypt: A traverse at 26o N. J Geol Soc London 140:75–95 Saunders AD, Tarney J, Marsh NG, Wood DA (1980) Ophiolites as ocean crust or marginal basin crust: a geochemical approach. In: Panayiotou A (ed) Ophiolites-Proceedings of the International Ophiolite Symposium, Cyprus, pp 193–204 Serri G (1981) The petrochemistry of ophiolitic gabbroic complexes: a key for the classification of ophiolites into low-Ti and high-Ti types. Earth Planet Sci Lett. 52:203–212 Saunders AD, Tarney J (1984) Geochemical characteristics of basaltic volcanism within back-arc basins. In: Kokelaar BP, Howells MF (eds), Geol Sot London, Spec Publ, 16: 59–76 Shackleton RM, Ries AC, Graham RH, Fitches WR (1980) Late Precambrian ophiolitic mélange in the Eastern Desert of Egypt. Nature 285:472–474 Shervais JW (1982) Ti-V plots and the petrogenesis of modern and ophiolitic lavas. Earth Planet Sci Lett 59:101–118 Stern RJ (1994) Arc assembly and continental collision in the Neoproterozoic East Africa Orogen: Implications for the consolidation of Gondwanaland. Ann Rev Earth Planet 22:319–351 Stern RJ, Johanson PR, Kröner A, Yibas B (2004) Neoproterozoic ophiolites of the Arabian-Nubian Shield. In: Kusky TM (ed) Precambrian ophiolites and related rocks. In Developments in precambrian geology: Amsterdam, Elsevier, 13, pp 95–128 Stern RJ (2008) Neoproterozoic crustal growth: The solid Earth system during a critical episode of Earth history. Gondwana Res 14:33–50 Sun S, Nesbitt RW (1982) Geochemical regularities and genetic significance of ophiolitic basal &. Geology 6: Taylor B, Martinez F (2003) Back-arc basin basalt systematic. Earth and Planetary Sci Lett 210:481–497 Vuagnat, M (1975) Pillow lava flows: isolated sacks or connected tubes. Bull Volcanol 39:581–589 Zimmer M, Kröner A, Jochum KP, Reischmann T, Todt W (1995) The Gabal Gerf complex: a Precambrian N-MORB ophiolite in the Nubian shield, NE Africa. Chem Geol 123:29–35

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Evidence for Mesoproterozoic Components in the Arabian-Nubian Shield Hamdy H. Abd El-Naby

Abstract

This chapter reviews and summarizes the early ideas on the existence of Mesoproterozoic components in the Arabian-Nubian Shield (ANS). It sheds more light on their significance in crustal evolution of the ANS. Nowadays, there is a general agreement among geologists that the Late Mesoproterozoic to Cambrian period bears the most remarkable geological record in Earth’s history. It began with the assembly of Rodinia supercontinent (1300–900 Ma) and ended by the amalgamation of Gondwana supercontinent (around 550 Ma). It is widely accepted that the ANS were formed by juvenile magmatic arc accretion and subsequent shield-wide post-tectonic magmatism. However, the recorded Mesoproterozoic crust, with subordinate Palaeoproterozoic and rare Archaean components, proves that ANS is less juvenile than previously thought, and may provide important clues for the reconstruction of the crustal evolution of the ANS. Recent studies by several workers have shown that inherited zircon grains from igneous rocks within the ANS may provide evidence for Mesoproterozoic crustal growth. The Mesoproterozoic components of the ANS were interpreted as pervasively reworked pre-existing crust during the Pan-African orogeny, which was responsible for the final juxtaposition of the collage of terranes now seen, e.g. reworked pre-Neoproterozoic crustal fragments at Sa’al Mesoproterozoic rocks, Wadi Rutig volcano-sedimentary succession and Wadi Solaf metapsammitic and granodioritic biotite gneisses in Sinai (Egypt); Khida subterrane in eastern Saudi Arabia; the arc–gneiss collages of the Precambrian basement of Yemen and the eastern Ethiopian–northwestern Somalian crustal block. Sediments are good natural samplers of the existing continental crust at the time of erosion and sedimentation. Several workers have presented detrital H. H. A. El-Naby (&) Faculty of Earth Sciences, King Abdul Aziz University, Jeddah, Saudi Arabia

zircon isotopic dataset from the Palaeozoic sedimentary sequences of the ANS (e.g. Meinhold et al. 2020). They interpreted the Ediacaran to middle Tonian zircon grains as being derived from igneous rocks of the ANS, whereas Palaeoproterozoic and Archaean grains were interpreted as a primary derivation from Palaeoproterozoic and Archaean basement found in some parts of the ANS or crustal components of the eastern Saharan Metacraton. This is consistent with palaeocurrent indicators, which suggest sediment transport from the hinterland of Gondwana. In the light of the above, future efforts to understand ANS crustal evolution will require more zircon geochronology and related studies to answer two main questions: (1) how isotopically juvenile Neoproterozoic magmas were contaminated with preNeoproterozoic zircon? and (2) was zircon assimilated from unexposed older crust that underlies the contaminated shield or derived from Neoproterozoic sediments sourced from a neighbouring pre-Neoproterozoic craton? Keywords








Mesoproterozoic Gondwana supercontinent Neoproterozoic magmas Zircon geochronology

14.1

Introduction

Although Mesoproterozoic witnessed high amounts of igneous activity, possibly due to large-scale mantle “superplumes” forming underneath the first continental or supercontinental granitic masses (Gladkochub et al. 2016; Tretyakova et al. 2016), only a few igneous complexes whose Mesoproterozoic age has been confirmed by reliable geochronological dating in the ANS. This may mean that a Palaeoproterozoic supercontinent did not fully fragment before Rodinia began to form (Condie 2001, Table 14.1). The Late Mesoproterozoic began with the formation of the

© The Author(s), under exclusive license to Springer Nature Switzerland AG 2021 Z. Hamimi et al. (eds.), The Geology of the Arabian-Nubian Shield, Regional Geology Reviews, https://doi.org/10.1007/978-3-030-72995-0_14

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344

H. H. A. El-Naby

Table 14.1 Proterozoic Aeon

Source International Commission on Stratigraphy’s Chronostratigraphic Chart 2018

Rodinia supercontinent through worldwide orogenic events between *1300 Ma and 900 Ma involving most continental blocks at that time (Cawood 2005; Li et al. 2008). This period is also known for the first evidence of the long-lived Mesoproterozoic mountain-building Grenville Orogeny, in which assembly of the supercontinent Rodinia occurred. The Mesoproterozoic is also the most prolific time for Stromatolites (Grotzinger and Knoll 1996; Knoll and Semikhatov 1998), becoming even more common ca. 2.2 Ga and began to decline during the Neoproterozoic. The construction of this chapter is based primarily on a compiled data of previous geological and geochronological studies on the Mesoproterozoic components in the ANS. This will be helpful in understanding the significance of these components in crustal evolution of the ANS. The present work examines and summarizes the early ideas on the existence of pre-Pan-African crust and Mesoproterozoic or older xenocrystic and detrital zircons from many recorded and common occurrences in the ANS, with shedding some light on the nature of the older cratonic areas flanking the ANS.

14.2

Early Ideas on the Existence of Pre-Pan-African Crust in the ANS

Based on Rb-Sr and K-Ar ages in Africa, the term Pan-African was launched by Kennedy (1964) for a tectonothermal event around 500 ± 100 Ma, where a series of mobile belts in Africa formed among much older African cratons. He believed that the Pan-African episode

had led to the structural differentiation of the entire continent into cratons and orogenic areas. This event has been extended to as much as *1000–450 Ma (Kröner 1984; Rogers and Santosh 2004) and recognized in other Gondwana continents where regional names such as Brasiliano for South America, Adelaidean for Australia and Beardmore for Antarctica have been proposed (Kröner 1984; Stern 1994; Kröner and Stern 2004; Rino et al. 2008). Another view regarding the identification of the term Pan-African has been put forward by Abdel-Rahman (1995). He proposed that the term Pan-African should be redefined as the geodynamic and geochemical differentiation of East Africa and Arabia into mobile zones and tectonic provinces, which took place during the ca. 950–550 Ma tectonothermal orogenic event. Kröner and Stern (2004) stated that Pan-African cannot be a single event but must be a protracted cycle reflecting the opening and closing of large ocean realms as well as the accretion and collision of buoyant crustal blocks. However, the recent study by Fritz et al. (2013) has shown that the term “Pan-African” has lost its significance as defining a particular event. Two broad types of orogenic or mobile belts can be distinguished within the Pan-African domains (Kröner and Stern 2004). The first type is dominated by Neoproterozoic crustal and magmatic rocks, such as ophiolite suites, granites associated with subduction or collision, arcs or passive continental margin assemblages, and some exotic terranes. Such belts include the ANS in northeastern Africa (Fig. 14.1). The other type of mobile belts is dominantly manifested in the polydeformed high-grade metamorphic

14

Evidence for Mesoproterozoic Components …

assemblages from the middle-lower crust of the region and is mostly Mesoproterozoic to Archaean with strong reworking in the Neoproterozoic. Such belt includes the Mozambique belt (including the Madagascar extending to western Antarctica), the Zambezi Belt of northern Zimbabwe and Zambia and, possibly, the little known migmatitic terranes of Chad, the Central African Republic, the Tibesti Massif in Libya and the western parts of Sudan and Egypt (Kröner and Stern 2004). In the last three decades, there has been considerable debate among geologists about the contribution of pre-Pan-African (>1 Ga) crust to the gneissic rocks or granitoids of the ANS. In Egypt, El Gaby et al. (1988b) identified a group of pre-Pan-African rocks that consist mainly of Archaean to Early Proterozoic granites, gneisses and schists and their mylonitized and remobilized equivalents cropping out at Gabal Uweinat in the south-western corner of Egypt. The gneissic complexes, especially the Meatiq, El Sibai, El-Shalul and Migif-Hafafit complexes of the Eastern Desert and the Feiran complex in Sinai are included in this group (El Gaby et al. 1990). These gneissic

345

complexes were thought of being the eastern continental margin or as a minimum comprising parts of the pre-PanAfrican (Archaean/Early Proterozoic) Saharan Metacraton, previously known as Nile Craton or Sahara Congo Craton or Eastern Saharan Craton, onto which the volcanosedimentary-ophiolite and related granitoids of the ANS were accreted (Abdelsalam et al. 2002). Pohl (1979) inferred that these gneissic complexes represent the northern continuation of the Mozambique belt. Kröner (1985) and Kröner et al. (1987) offered a model whereby the disputed pre-PanAfrican continental crust in the Central Eastern Desert was proposed to exist as small continental fragments within the Mozambique Ocean. The age of the gneisses and their tectonic environment in the ANS is crucial for settling the controversy of whether the high-grade metamorphic gneisses and amphibolites in Eastern Desert and Sinai represent part of the older continental foreland during the Pan-African orogeny (pre-PanAfrican) or they are metamorphosed Pan-African rocks. Several workers are of the opinion that the gneisses at Meatiq and Hafafit in the Eastern Desert of Egypt represent

Fig. 14.1 Map of Gondwana at the end of Neoproterozoic time (*540 Ma) showing the general arrangement of Pan-African belts. Modified from Kusky et al. (2003), Kröner and Stern (2004), Küster et al. (2008)

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deeper levels of juvenile Neoproterozoic crust or Archaean/Palaeoproterozoic crust that was remobilized during Neoproterozoic time. Hume (1934) identified the high-grade metamorphic rocks at Meatiq area west of Quseir, Egypt, as a Protarchaean “fundamental basement” of old continental crust, separated by an unconformity from younger series. This fundamental basement extended beneath the entire Eastern Desert and Sinai. Habib et al. (1985a, b) identified the “fundamental basement” of Hume as being infracrustal basement overthrusted by a supracrustal cover. The infracrustal rocks were developed as a result of an old orogeny referred to as the Meatiqian orogeny where granitic gneisses, migmatitic gneisses and migmatized amphibolites were formed. Similarly, Abdel-Monem and Hurley (1979) suggested a pre-Pan-African age (1770 Ma) for Migif-Hafafit gneisses. Akaad and Noweir (1980) and Ries et al. (1983) considered the quartzofeldspathic gneisses at Meatiq and Hafafit as Neoproterozoic (or older) shelf sediments metamorphosed during the Pan-African orogeny. El Gaby (1983), El Gaby et al. (1988a) and Hassan and Hashad (1990) believe that the Mid-Proterozoic continental crust extends into the Eastern Desert and crops out in gneiss domes underneath overthrusted Pan-African island arc volcanics and volcaniclastics and associated ophiolites. Kröner et al. (1987) suggest that the ANS is a collage of small crustal blocks or microcontinents immersed in more abundant juvenile material. In agreement with the abovementioned authors, Neumayr et al. (1996) proposed a pre-Pan-African polymetamorphic evolution for Meatiq basement dome, which has been reworked during the Pan-African orogeny. In contrary, other workers (e.g. Stern and Hedge 1985; Kröner et al. 1994) suggested Pan-African emplacement ages of migmatitic orthogneiss protoliths that cluster around 687 Ma. This is in agreement with the findings of Abd El-Naby et al. (2008) who reported Sm–Nd age of 593 ± 4 Ma for Hafafit hornblende gneiss and an age of 585 ± 8 Ma for Hafafit amphibolites. They considered these ages to mark the end of peak metamorphic temperatures. Moreover, Liégeois and Stern (2010) concluded that the protolith for the Meatiq and Hafafit gneisses were juvenile Neoproterozoic igneous rocks. The conclusion was put forward primarily based on the Rb–Sr and U–Pb zircon ages of 600–750 Ma that give no support for any significant contribution of pre-Neoproterozoic crust in these two sections of Eastern Desert crustal infrastructure. This in agreement with the isotopic dating of gneiss domes at Meatiq (631 ± 2 Ma, Andresen et al. 2009) and Hafafit (659 ± 5, Lundmark et al. 2012). Hamimi et al. (1994) concluded that the El-Shalul area, central Eastern Desert of Egypt, is covered by two main lithotectonic units, the first unit is the infrastructure granitic gneisses and the other unit is the suprastructure ophiolitic

H. H. A. El-Naby

mélange. The infrastructure granitic gneisses were interpreted as being pre-Pan-African, based on Rb-Sr whole rock ages of ca. 1200 Ma (El-Manharawy 1977). This is in agreement with the conclusion of El Gaby (1983) who considered El-Shalul as old continental crust (extension of the Sahara Metacraton) of pre-Pan-African age that had been mylonitized during the Pan-African orogeny. On the other hand, recent radiometric dating by Ali et al. (2012) supports a juvenile origin for the El-Shalul granitic gneisses. Their conclusion is evidenced primarily by U-Pb zircon dating of El-Shalul gneisses (631 ± 6 Ma). El Sibai area of the Central Eastern Desert (CED) of Egypt consists of an ophiolitic association of arc metavolcanics, ophiolitic rocks, mélange, metasediments with minor mafic intrusions; and a gneissic association of amphibolite, gneissic diorite, tonalite, granodiorite and granite. Previous studies of El Sibai area have identified the gneissic association as a lower crustal infrastructure in sheared contact with upper crustal ophiolitic association suprastructure and have presented it as an example of a metamorphic or magmatic core complex. Similar to Meatiq and Hafafit gneissic complexes, El Sibai gneissic association was regarded by early workers as a pre-Pan-African lower crustal infrastructure in sheared contact with upper crustal arc metavolcanics and ophiolitic association suprastructure and have presented it as an example of a metamorphic or magmatic core complex (El Gaby 1983; El Gaby et al. 1988a; Kamal El Din et al. 1992; Khudeir et al. 1992). On contrary, Fowler et al. (2007) concluded that El Sibai complex is not a core complex, but exemplifies the overlap of NW–SE folding and NW–SE extensional which is a significant theme of CED regional structure. In addition, the Pan-African age of El Sibai Complex (680 Ma) has been demonstrated by Bregar et al. (2002) and Augland et al. (2012). Similar to previously mentioned gneissic complexes in Egypt, the age dating of Feiran metamorphic belt is a matter of debate. Many authors have considered the Feiran belt to be of Pan-African age (Bielski 1982; Stern and Manton 1987; Kröner et al. 1994), while others have considered it to be of pre-Pan-African age (Schürmann 1966; Siedner et al. 1974; Shimron 1980a, b, 1984a, b; El-Shafei and Kusky 2003). Abu Anbar et al. (2004) reported Sm–Nd age of 1419 ± 178 Ma for Feiran gneisses that is considered as the age of protolith formation (crystallization of parent rock), while the Sm–Nd isotopes of amphibolites yielded an age of 1118 ± 222 Ma for the protolith. These rocks had been later metamorphosed during the Pan-African event (around 630 Ma) which is contemporaneous with the intrusion of syn-orogenic granodiorites in the northern and eastern parts of Feiran gneiss belt (Abu Anbar et al. (2004). The Uweinat-Kamil basement complex, located at the borderline between Egypt and Sudan, is a unique inlier-outcrop in NE Africa that contains the oldest

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Evidence for Mesoproterozoic Components …

347

pre-Pan-African rocks in the Saharan Metacraton. Archaean ages as old as 3.3 Ga with SHRIMP U-Pb magmatic zircon were reported by Bea et al. (2011a) for the tonalite–trondhjemite–granite (TTG) gneisses of Gabal Kamil (Table 14.2). Based on petrography and Rb/Sr whole rock age data, Klerkx (1980) recognized two major basement series: (1) the high-grade granulitic Granoblastite Formation (Karkur Murr Series) as lower unit at the eastern and southeastern slopes of Jebel Uweinat, overlain by (2) the clearly remobilized Anatexite Formation (Ain Dua Series). Remnants of the pre-Pan-African crust were identified in Gabal El Asr inlier with Palaeoproterozoic Sm–Nd whole rock model age of 2399–1839 Ma for migmatites (Schandelmeier et al. 1988) and conventional U–Pb zircon age of 1.9–2.1 Ga for an anorthosite (Sultan et al. 1994). Gabal El Asr inlier probably represents the eastern transition between the Palaeoproterozoic to Archaean terranes of the Sahara Metacraton (i.e. Uweinat-Kamil inlier) and the juvenile Neoproterozoic terranes of the ANS. Similar to Egypt, pre-Pan-African gneisses have been recorded as older basement inliers in Somalia. These inliers were reactivated during Pan-African tectonothermal activities (Schandelmeier et al. 1990). Systematic geological mapping in northern Somalia suggested the presence of an older gneissic basement underlying the predominantly quartzofeldspathic rocks of the Mozambique sequence (Hunt 1958; Gellatly 1960). Previous work in the Bur area of Southern Somalia suggested that pelitic gneisses, quartzites, banded iron formation (BIF), dolomite/marble and orthoand para-amphibolite of the Dinsor formation overlie older rocks comprising paragneisses and amphibolites of the Ontole formation (D’amico et al. 1981). They attributed the

Table 14.2 Isotopic age determination for the basement complex of the Uweinat-Kamil inlier

high-grade metamorphism of these rocks to the Mozambique dynamothermal peak rather than an earlier event of reworking. However, the remnant lithologies, particularly the BIF, contrast with the predominantly clastic Upper Proterozoic sequence of Mozambique sediments and therefore are tentatively regarded as older reworked basement, but age determinations are needed to verify the stratigraphy proposed for the Bur Region (Warden and Horkel 1984). Pan-African syn- to post-tectonic granites intruded the high-grade paragneisses, migmatites, granitic orthogneisses and amphibolites, although the age of the metamorphism is uncertain. In northwestern Somalia and extending into eastern Ethiopia, the reactivated Early to Mid-Proterozoic basement consists of paragneisses and migmatites with intercalations of amphibolites and local occurrences of carbonate rocks, quartzite, and acidic to basic metavolcanics. There are several syn- to post-tectonic intrusions of granites, diorites and gabbros. The Northern Somalian Belt is regarded as a Pan-African ensialic mobile belt, consisting of a completely rejuvenated older crust (Western sector) and juvenile Pan-African terrains (Eastern Sector) (Sassi et al. 1993). The later basement can be correlated with that of Yemen on the opposite side of the Gulf of Aden. In the Ethiopian Precambrian, three complexes are recognized (Kazmin 1971, 1975). The Lower Complex formed of high-grade gneisses representing older cratonic basement (older than 2500 m.y.). The Middle Complex (clastic metasediments) is presumably the Lower to Middle Proterozoic platform cover. The Upper Complex consists of low-grade rocks in the following succession: ophiolitic rocks, andesitic metavolcanics and associated metasediments, clastic and to less extent carbonate sediments.

Location

Rock type

Method

Age (Ma)

References

Jebel Kamil

Anorthositic gneiss

U/Pb

2063–2629 1922–2141

Sultan et al. (1994)

Jebel Uweinat (Wadi Wahesh)

Mylonite

RB/Sr, regression line

2637 (± 393)

Cahen et al. (1984), Klerkx and Deutsch (1977)

Jebel Uweinat

Granulitic gneiss

Karkur Murr gneisses

Jebel Uweinat (South)

2556 (± 142) RB/Sr, model ages

2919–2904

Sm/Nd model ages

3000–3200

U-pb zircon crystallization age

3000

Harris et al. (1984) Bea et al. (2011a)

Sm–Nd model ages

3000–3200

Harris et al. (1984)

Migmatitic biotite gneiss

RB/Sr, isochron

1784 (± 126)

Cahen et al. (1984), Klerkx and Deutsch (1977)

Granodiorite gneiss

K/Ar biotite

1878 (± 64)

Hunting Geology and Geophysics Ltd. (1974)

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Kazmin et al. (1978) reconstructed the following sequence of events for the Ethiopian crust: (1) formation of ancient high-grade gneisses, mainly Archaean that have undergone at least two stages of tectonism and metamorphism prior to the Pan-African event. (2) Accumulation of platform-type psammitic and pelitic sediments of the Wadora Group, most probably belong to the Middle or the Lower Proterozoic. (3) Rifting of older basement in the Red Sea area and opening of zones with oceanic crust (ophiolites). Many authors have reported polydeformed high-grade metamorphic assemblages in the western parts of Sudan (Vail 1983; Abdel-rahman et al. 1993; Stern et al. 1994; Küster and Liégeois 2001). These rocks consist predominantly of much older Mesoproterozoic to Archaean continental crust that was strongly reworked during the Neoproterozoic. A pre-Pan-African sialic basement complex in the Arabian Shield has been documented by many earlier works (e.g. Baubron et al. 1976; Delfour 1980; Stacey and Stoeser 1983). Khida terrane, in the eastern Arabian Shield, is regarded as one of the significant component of Late Archaean or Early Palaeoproterozoic age. However, Late Archaean source region for some rocks of the Khida terrane are not presently known in Saudi Arabia. However, the postulated Archaean source region may correlate with the Late Archaean gneisses in the Al-Mahfid terrane of eastern Yemen, which have been dated at ca. 2.55 Ga (U-Pb zircon, Whitehouse et al. 1998; Sm–Nd, Windley et al. 1996). Based on common lead data for feldspar, whole rock and galena samples, ANS can be divided into two main groups (Stacey et al. 1980; Stacey and Stoeser 1983; Stacey and Hedge 1984). The first group has oceanic (mantle) characteristics and are found in rocks from the Red Sea Hills of Egypt and the western and southern parts of the Arabian Shield, whereas the other group has incorporated a continental-crustal component of at least Early Proterozoic age and are found in rocks from the northeastern and eastern parts of the Arabian Shield, as well as from the southeastern Shield near Najran. They are also found in rocks to the south in Yemen, to the east in Oman, and to the west at Aswan, Egypt. Based on isotopic data, Li et al. (2018) demonstrated that reworking of old continental crust played important roles in generation of oceanic arcs in the northwestern ANS that is likely much less juvenile than previously thought. Although most of magmatic rocks of the Arabian Shield are proven to be a juvenile Pan-African crust, they contain zircon fractions that yielded older ages (Archaean to Proterozoic). Calvez et al. (1985) interpreted the U–Pb age of 2067 ± 74−72 Ma of a zircon fraction, separated from gabbroic-trondhjemitic rocks in the Al Amar suture zone, as evidence for the presence of much older crust in that part of the Shield. Similarly, Kennedy et al. (2004) found minor inherited zircon in volcanic rocks from the central Afif and Hijaz terranes that yielded U–Pb ages as old as 1565–

H. H. A. El-Naby

2750 Ma. It is worthy to mention that these zircon grains may have experienced isotopic disturbance. Kennedy et al. (2004) interpreted the pre-Neoproterozoic ages as evidence of inheritance from older crustal material. Although sedimentary origins for the inherited zircon are possible, some inherited grains exhibit characteristics more consistent with juvenile crystals, such as euhedral morphologies and a lack of recognizable cores, which suggest that some grains were assimilated in situ from their original host igneous rocks (i.e. from cryptic pre-Neoproterozoic crust). The age of the cryptic basement, if it exists, is inferred from the ages of inherited zircon grains to be dominantly Mesoproterozoic, with subordinate Palaeoproterozoic and rare Archaean components (Hargrove et al. 2006).

14.3

Mesoproterozoic or Older Xenocrystic and Detrital Zircons in the ANS

Detrital zircon U–Pb geochronology is a powerful tool to determine the maximum depositional age for clastic sediments, to characterize potential provenance areas and to assess palaeogeographic models, especially in areas where other methods, such as palaeomagnetism or fossil distribution, cannot be used. The growing number of ion microprobe zircon ages in the ANS has revealed that a significant number of ANS igneous rocks contain zircons that are meaningfully older than the crystallization ages of the host rocks (Küster et al. 2008; Be’eri-Shlevin et al. 2009a; Stern et al. 2010). In general, the ANS comprises Early Neoproterozoic to Cambrian (*850–530 Ma) tectonostratigraphic terranes formed by the closure and accretion of juvenile volcanic arcs and back-arc basins associated with Gondwana assembly. However, much older zircon ages than this have recorded in many areas in the ANS, e.g. the Palaeoproterozoic of Khida subterrane in Saudi Arabia (*1.8–1.67 Ga) and terranes of Yemen, and Mesoproterozoic volcano-sedimentary rocks of Sinai, Egypt (1.03 and 1.11 Ga). Though, there is another emerging trend where some researchers believe in the presence of pre-Neoproterozoic crust at depth, in spite of Nd isotopic evidence that ANS crust is overwhelmingly juvenile (Stern et al. 2010). Recent studies by several workers (e.g. Stern et al. 2010) show that about 5% of individually dated zircons from ANS Neoproterozoic igneous rocks have ages older than 880 Ma, with concentrations in the Tonian– Mesoproterozoic (0.9–1.15 Ga); Late Palaeoproterozoic (1.7–2.1 Ga); Palaeoproterozoic–Neoarchaean (2.4–2.8 Ga); and Early Archaean (>3.2 Ga). Inheritance in the Arabian Shield was first indicated by a report of zircons dating 1986 ± 200 Ma (Calvez et al. 1985) in Neoproterozoic plagiogranite immediately west of Al Amar fault. Explanations for inheritance include:

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Evidence for Mesoproterozoic Components …

(1) contamination during processing; (2) assimilation from cryptic Early Neoproterozoic to Archaean basement; (3) assimilation of terrigenous sediment shed from nearby passive margins transported by rivers or glaciers; and (4) inheritance from a mantle source. Meinhold et al. (2020) presented detrital zircon U–Pb age dataset from Palaeozoic sandstones of Saudi Arabia. Five main age populations are present in varying amounts in the zircon age spectra, with age peaks at *625 Ma, *775 Ma, *980 Ma, *1840 Ma and *2480 Ma. They interpreted the Ediacaran to middle Tonian zircon grains (625 Ma, *775 Ma, *980 Ma) as being derived from igneous rocks of the ANS, whereas Palaeoproterozoic and Archaean grains (*1840 Ma and *2480 Ma) may be explained in two alternatives: (a) A primary derivation from Palaeoproterozoic and Archaean basement, as rocks of such age occur in the vicinity, e.g. *1.66 Ga crust of Khida terrane in the eastern Arabian Shield (Whitehouse et al. 2001a, b), *2.55–2.95 Ga crust in the Arabian Shield of Yemen (Whitehouse et al. 1998), and the Palaeoproterozoic and Archaean crustal components of the eastern Saharan Metacraton (e.g. Küster et al. 2008; Bea et al. 2011b; Zhang et al. 2019). b) They may represent xenocrystic cores originally derived from igneous rocks of the ANS, but this would only account for a small number of grains (Stern et al. 2010). Garzanti et al. (2013) identified similar age populations in two Palaeozoic sandstone samples from the Tabuk area. The Devonian samples of Palaeozoic sandstones of Saudi Arabia yielded the oldest zircon ages yet recorded from Arabia and the surrounding area: 3771 ± 37 Ma and 3850 ± 25 Ma in the Jubah Formation (Meinhold et al. 2020). The rounded morphologies of many inherited grains are consistent with a sedimentary origin, whereby sediments were derived from a nearby continent, transported fluvially or by glaciers, and deposited into basins undergoing arc volcanism. According to the Gondwana super-fan model (Fig. 14.2), Saudi Arabia would lie more distal along the sediment path to the northeast. Similar Palaeozoic sandstone is also recorded around the Mekelle Basin in the Tigray province of northern Ethiopia and, to a minor extent, in the Blue Nile region in the west of the country, and comprise Upper Ordovician and Carboniferous–Permian glaciogenic sandstones (e.g. Lewin et al. 2018, and references therein). A Stenian–Tonian (ca. 1 Ga) zircon population characterizes the Upper Ordovician–

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Lower Silurian Enticho Sandstone. Lewin et al. (2018) described the Stenian–Tonian zircons as being derived probably from the centre of the East African Orogen and were supplied via the Gondwana super-fan system (Fig. 14.2). This material was transported by the Late Ordovician glaciers and formed the Enticho Sandstone, whereas Tonian (ca. 800 Ma) zircons represent the earliest formation stage of the southern ANS. Another pre-Pan-African detrital zircons ages have been recorded in the granulite facies terrigenous garnet–cordierite–gneisses at Sabaloka in Sudan (Kröner et al. 1987). These Mesoproterozoic depositional ages (TDM age of 1.69 Ga) are similar to the 1.58 and 1.63 Ga TDM ages reported by Harris et al. (1984) for meta-sedimentary gneisses from the Rahaba mine area of central Bayuda. Küster and Liégeois (2001) attributed the Mesoproterozoic TDM Nd model ages to a mixture of detritus from Palaeoproterozoic and Neoproterozoic sources that were shed into the basin at the same time. Küster et al. (2008) attributed the old metamorphic zircons from basement rocks of the Bayuda and Sabaloka terrane in Sudan to an orogenic event starting in the Early Neoproterozoic named as Bayudian event (920–900 Ma) and ending during Late Pan-African times (ca. 600–580 Ma). Zircon xenocrysts within the arc-related calc-alkaline volcanic and intrusive rocks of the Sa’al metamorphic complex (SMC) in Sinai (Egypt) yielded ages of ca. 1.03– 1.02 Ga (Be’eri-Shlevin et al. 2012). These ages may represent the oldest magmatic events in the northernmost ANS. Be’eri-Shlevin et al. (2009b) recorded similar ages (1.03 and 1.11 Ga) from the pelitic sediments of Ra’ayan formation (SMC). These ages may provide a possible connection between latest Mesoproterozoic ocean closure during the assembly of Rodinia and the later build-up of Gondwana (Be’eri-Shlevin et al. 2012). Shimron et al. (1993a, b) suggested that the calc-alkaline volcanic rocks and the pelitic sediments of Sa’al metamorphic complex represent the island arc and its fore-arc and/or trench sediments. On the other hand, tholeiitic to calc-alkaline Neoproterozoic magmatism characterizes the earliest island arcs throughout the ANS (Roobol et al. 1983; Johnson and Woldehaimanot 2003; Ali et al. 2010). Lancelot and Bosch (1991) gave another direct evidence for older basement in Zabargad Island, Egypt. They interpreted the initial Nd isotopic data of granulite gneisses from Zabargad Island as indicate extraction of the protoliths from depleted mantle after 1200 Ma, and possibly as early as 1700 Ma. Those model ages are comparable to Mesoproterozoic ages for inherited zircon in the Arabian Shield (Kennedy et al 2004; Hargrove et al. 2006).

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Fig. 14.2 Simplified geological map of North Gondwana showing present-day distribution of Precambrian basement rocks and Cambrian– Ordovician sediments. Modified after Avigad et al. (2012)

14.4

Nature of the Older Cratonic Areas Flanking the ANS

Palaeoproterozoic and older continental crust have been identified in the cratonic areas flanking the ANS (Fig. 14.1) that extends about 3000 km north to south and more than 500 km on either side of the Red Sea (Kröner and Stern 2004). It is flanked to the west by the Saharan Metacraton which includes a broad tract of older crust that was remobilized during Neoproterozoic time along with a significant amount of juvenile Neoproterozoic crust (Abdelsalam et al. 2002) and documented ca. 0.9–0.92 Ga orogeny in the Bayuda Desert (Küster et al. 2008). The extent of juvenile Neoproterozoic crust to the east in the subsurface of Arabia is not well defined, but it appears that Pan-African crust underlies most of this region. To the east, scattered older continental outcrops in Saudi Arabia (Palaeoproterozoic to Archaean continental crust of Khida terrane and Mesoproterozoic age crust of the Afif terrane); and Palaeoproterozoic

and Archaean crust in Yemen that was overprinted by Pan-African tectonomagmatic events (Stacey and Hedge 1984; Whitehouse et al. 1998, 2001a, b, Kröner and Stern 2004). To the north, the west margin of the ANS is not defined because it is covered by Mesozoic to Cenozoic sedimentary rocks but it extends along the line of the Nile Valley and crops out in the Keraf arc-continent suture in northern Sudan (Abdelsalam et al. 1998). To the south, the Mozambique Belt (Fig. 14.1) which represents the southern part of the axial region between converging Gondwana cratonic blocks during the culmination of the supercontinental cycle that began with the breakup of the supercontinent Rodinia and ended with the assembly of the supercontinent Gondwana (Stern 1994). The ANS is distinguished from the Mozambique Belt by its dominantly juvenile nature, relatively low grade of metamorphism, and abundance of tonalitic to granodioritic rocks and ophiolites. The Saharan Metacraton stretches between the Tuareg Shield in the west, the ANS in the east and the Congo and Tanzanian Cratons in the south (Fig. 14.1). Its borders have

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Evidence for Mesoproterozoic Components …

been defined based on lithological, tectonic and geochronological evidences. Rocks of the Saharan Metacraton are patchily exposed in northwestern, central and western Sudan and in south-western Egypt (Fig. 14.1). They are predominantly made up of high-grade metamorphic gneisses, migmatites and supracrustal rocks, commonly foliated along ENE–WSW to NE–SW structures. Phanerozoic sedimentary rocks and desert sands overlie more than 50% of the Saharan Metacraton (Abdelsalam et al. 2011). To the north, rocks of the Metacraton disappear beneath the Phanerozoic cover of the northern margin of Africa. However, in its northern part, Küster and Liégeois (2001) have found Neoproterozoic high-grade rocks indicating that the boundary between the ANS and the Saharan Metacraton may lie further west or that these high-grade Neoproterozoic rocks have been thrusted across the Metacraton. The southern boundary is not well defined, but is marked by the Oubanguides orogenic belt, which separates it from the Congo and Tanzanian Cratons (Abdelsalam et al. 2002). Nd isotopic data of the high-grade basement of the northeastern Saharan Metacraton suggest Archaean, Palaeoproterozoic and Neoproterozoic ancestries with a direction of approximate crustal growth from NW to SE. Archaean crust (Nd TDM ages > 3100 Ma, Harris et al. 1984) is exposed only in the Uweinat massif of SE Libya/SW Egypt. In SW Egypt (Gebel Kamil, Gebel El Asr), NW Sudan (Nubian Desert, Gebel Rahib) and W Sudan (Gebel Marra region) isotopic data of tonalitic and granitic orthogneisses and of migmatites have confirmed the existence of mainly Palaeoproterozoic crust (Nd TDM ages between 1900 and 2500 Ma, Harris et al. 1984; Harms et al. 1990); slightly younger Nd TDM ages of 1600–1700 Ma were found in tonalitic gneisses at Wadi Howar, Sudan and Chad (Harms et al. 1990). This Palaeoproterozoic basement was intensely remobilized and deformed during the Neoproterozoic Pan-African orogeny (Harms et al. 1990; Schandelmeier et al. 1990). The rest of the Saharan Metacraton comprises mostly granitoids and granulites, paragneisses and various schists with imprints of several fold stages (e.g. Fleck et al. 1973; Huth et al. 1984), giving Neoproterozoic radiometric ages between 500 and 700 Ma that correspond to the Pan-African tectonothermal event (e.g. Kennedy 1965), including the time of the East and West Gondwana collision (e.g. Shang et al. 2010). The East African Orogen (EAO) is a Neoproterozoic– Early Cambrian mobile belt that today extends south along eastern Africa and western Arabia from Sinai and Jordan in the north to Mozambique and Madagascar in the south (Fig. 14.1). Jacobs et al. (1998), Jacobs and Thomas (2004) and Grantham et al. (2011) proposed a southern continuation of the EAO, from Mozambique to Antarctica. In the north, the EAO is subdivided into the ANS which composed largely of juvenile Neoproterozoic crust (e.g. Stern 1994, 2002; Johnson and Woldehaimanot 2003; Johnson

351

et al. 2011), and the Mozambique Belt in the south comprising mostly pre-Neoproterozoic crust with a Neoproterozoic–Early Cambrian tectonothermal overprint (e.g. Fritz et al. 2005; Collins 2006; De Waele et al. 2006; Viola et al. 2008; Bingen et al. 2009). The ANS is interpreted to represent the northern continuation of the higher grade and more strongly deformed (amphibolite and granulite facies) rocks of the Mozambique Belt in East Africa and Antarctica. The ANS and Mozambique Belt, both of which are sandwiched between East and West Gondwana, together constitute the East African Orogen (Fig. 14.1). The Nubian or African part of the ANS includes Egypt, Sudan, Ethiopia and Somalia, whilst Saudi Arabia and Yemen constitute the Arabian part of the Shield. Whilst the Mozambique Belt is bounded by well-defined pre-Neoproterozoic cratonic blocks, the western, northern and northeastern margins of the ANS are less well defined or known, being masked by cover sequences. Beyond the Kenya-Tanzania province, the Mozambique Belt splits into two segments. One segment continues northeast through southeastern Ethiopia and Somalia into Yemen across the Gulf of Aden, while the other segment runs northwestwards into Sudan. Both segments, characterized by high-grade gneisses and metasediments, are separated by large wedges of low-grade volcano-sedimentary rocks containing ophiolites. Azania Block constitutes one or more continental blocks between the Indian Shield and Congo–Tanzanian Cratons (Fig. 14.1). It was defined by geochronological data (Collins and Windley 2002) as a micro-continent of Archaean and Palaeoproterozoic crust (2900–2450 Ma), extending over Madagascar, Somalia, and Arabia (Afif terrane). The eastern and western margins of Azania Block are marked by Neoproterozoic volcano-sedimentary sequences and rocks formed in an oceanic environment (Collins and Pisarevsky 2005). The ocean separating Azania from East Africa was interpreted to have started by an intra-oceanic arc in the Tonian (Jöns and Schenk 2008; Collins et al. 2012). This ocean is thought to have finally closed by continent–continent collision between Azania and the Congo–Tanzania Craton at *630 Ma (Collins and Windley 2002; Collins 2006) causing high-grade metamorphism and contractional deformation in south-western Madagascar (Jöns and Schenk 2011) and eastern Africa (Möller et al. 2000; Hauzenberger et al. 2004, 2007). The Congo Craton refers to the amalgamated central African landmass at the time of Gondwana assembly (De Waele et al. 2008), and is comprised of Mesoarchaean blocks and Palaeo- to Mesoproterozoic mobile belts amalgamated during the assembly of Gondwana (Goodwin 1996; De Waele et al. 2008), with regions of extensive Phanerozoic cover (De Waele 2005). The Tanzanian Craton is located on the eastern side of the Congo Craton and is

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interpreted to constitute single continental block from the Early Proterozoic onwards (Fig. 14.1). It is made up of two blocks separated by the Dodoma Schist Belt. The northern Nyanzian Block is comprised of greenstone belts and granites ranging in age from 2.80 to 2.66 Ga, overlain by a molasse that is intruded by 2.60 Ga granitoids. The southern Dodoman Block consists of 2.93–2.85 Ga granodiorites, granitic gneisses, migmatites and other high-grade metamorphic rocks (Cahen et al. 1984; Schulter 1997; Manya and Maboko 2003; Begg et al. 2009 and references therein). North of the Congo and Tanzanian Cratons lies the Neoproterozoic Oubanguides Belt which resulted from the collision between the Congo Craton, the Säo Francisco Craton and the West African Craton during the formation of Gondwana (Castaing et al. 1994; Toteu et al. 2004). The mobile belt consists of 0.90–1.7 Ga sedimentary rocks over a 2.0 Ga basement (Begg et al. 2009 and references therein). Abdelsalam et al. (2002) have put forward another view regarding the formation of Oubanguides Belt. According to them, the formation of the Oubanguides orogenic belt was resulted from the collision between the Saharan Metacraton and Congo and Tanzanian Cratons. East of the Tanzania Craton is the Neoproterozoic Mozambique Belt which is the longest terrane in Africa, extending along the east African coast from Mozambique to southern Egypt (Schulter 1997). Clifford (1970) coined the term Kalahari Craton for the first time for the Precambrian shield of Southern Africa. It consists of three major crustal domains, the Kaapvaal Craton to the south, the Zimbabwe Craton to the north and the Limpopo Belt between them. These domains made up of crustal components of different ages, comprising a composite Archaean nucleus partly surrounded by progressively younger accreted belts of Palaeoproterozoic and Mesoproterozoic age (Hartnady et al. 1985). Based on U–Pb and Lu– Hf isotope analyses of granitoid samples, Zeh et al. (2009) concluded that Kalahari Craton consists of at least five distinct terranes, which underwent different crustal evolutions, and was successively accreted at ca. 3.23 Ga, 2.9 Ga and 2.65–2.7 Ga.

14.5

Significance of the Mesoproterozoic Rocks in Crustal Evolution of the ANS

There is general consensus that the ANS formed by juvenile magmatic arc accretion that evolved during the Neoproterozoic above a multiplicity of subduction zones in the Mozambique Ocean. The arcs have tholeiitic to calc-alkaline, and locally MORB chemistry and include basalt, boninite, andesite and rhyolite, voluminous volcaniclastic and sedimentary rocks, and large amounts of diorite, tonalite, trondhjemite and granodiorite. Abundant Neoproterozoic ophiolites (ca. 870–690 Ma) in the ANS are

consistent with a convergent margin setting (Berhe 1990; Stern et al. 2004). However, Mesoproterozoic rocks prove that ANS less juvenile than previously thought, many plate reconstruction models infer the existence of a supercontinent prior to the formation of Gondwana. This supercontinent commonly referred to as Rodinia. The breakup of Rodinia was initiated most likely by a mantle superplume (Stein and Goldstein 1996) and amalgamated ca. 1.3–0.9 Ga, and its following breakup provided the components for Gondwana (Meert 2003; Li et al. 2008). Based on growth rates in the ANS, two schools of thoughts have been recognized, the first one assumed normal growth rates for continental crust in the ANS, i.e. simple arc-accretion model from 900 to 600 Ma (e.g. Dixon and Golombek 1988). The second school of thoughts modelled growth rates of an order of magnitude greater than the rates estimated for arcs since the end of the Palaeozoic. (e.g. Reymer and Schubert 1986; Stein and Hofmann 1994; Stein and Goldstein 1996). They suggested that ANS assembly involved the accretion of oceanic plateaux as well as arc terranes, because plateaux are the largely unsubductable surface expressions of voluminous juvenile lithosphere emplaced by mantle plumes over brief time periods. Be’eri-Shlevin et al. (2012) interpreted the Mesoproterozoic rocks (ca. 1.03–1.02 Ga) of Sa’al metamorphic complex (SMC) in Sinai (Egypt) in two ways, the first is that these rocks represent the earliest stages of the ANS, implying a link between latest Mesoproterozoic ocean closure during the assembly of Rodinia and the later build-up of Gondwana. The second interpretation is the SMC is unrelated to the ANS, implying that rocks from more than one oceanic basin are represented in the shield. The later scenario is supported by Küster et al. (2008) who gave evidence for the presence of an orogenic phase represented by ca. 0.9–0.92 Ga (Bayudian event) that preceded the Pan-African orogenic cycle (860–590 Ma) at the eastern boundary of the Saharan Metacraton. Stern et al. (2010) conclude that the presence of pre-880 Ma zircon xenocrysts in ANS igneous rocks with mantle-like isotopic compositions indicates either incorporation of sediments or inheritance from the mantle source region, or both. In addition to the Sa’al Mesoproterozoic rocks in Sinai (Egypt), reworked pre-Neoproterozoic crustal fragments are recognized at Wadi Rutig volcano-sedimentary succession, Sinai (ca. 900–1100 Ma, Samuel et al. 2011), Wadi Solaf metapsammitic and granodioritic biotite gneisses, Sinai (1.0 Ga zircons, Abu El-Enen and Whitehouse 2013), Khida subterrane in eastern Saudi Arabia (Agar et al. 1992; Stacey and Agar 1985; Whitehouse et al. 2001a, b), the arc–gneiss collages of the Precambrian basement of Yemen (Whitehouse et al. 2001a, b; Windley et al. 1996; Yeshanew et al. 2015) and the eastern Ethiopian–northwestern Somalian crustal block (Kröner and Sassi 1996; Teklay et al. 1998;

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Evidence for Mesoproterozoic Components …

Yeshanew et al. 2017). The recognition of older (i.e. pre-Pan-African > 1000 Ma) crustal components in ANS either as discrete rocks or through Pb isotope systematics has revived interest in models that involve repeated rifting of an Archaean to Mid-Proterozoic craton (Garson and Shalaby 1976; Kröner 1979; Kemp et al. 1980; Stern 1981; Calvez and Kemp 1982; Stacey and Stoeser 1983; Stacey and Hedge 1984; Claesson et al. 1984). Several workers are of the opinion that the Mesoproterozoic zircons are a key feature of the Gondwana super-fan system and were likely transported from regions in the centre of the East African Orogen, i.e. from the Mozambique Belt, or from regions of the SE Saharan Metacraton towards the continental margin of Gondwana during the Early Palaeozoic (Meinhold et al. 2013, 2020). After the assembly of Gondwana during the Early Palaeozoic, a thick pile of quartz-rich sandstones was deposited across northern Gondwana (Fig. 14.2), originally eroded from the Pan-African Orogen (e.g. Garfunkel 2002; Burke et al. 2003; Avigad et al. 2005). The exact provenance of these quartz-rich sandstones remains unclear. Based on detrital zircon U–Pb age spectra in Palaeozoic sandstones throughout Gondwana, Squire et al. (2006) suggested super-fan model where most of these zircons were derived from the East African Orogen towards the continental margins, rather than from the adjacent ANS. Later, the super-fan model was extended to the northern Gondwana margin, i.e. North Africa, the Sinai Peninsula and the northwesternmost part of the Arabian Peninsula (Meinhold et al. 2013). Based on detrital zircon age spectra from Lower Palaeozoic sedimentary rocks, Stephan et al. (2019a, b) suggested three contrasting provenance end-members of the former Gondwanan shelf, namely the Avalonian, West African and East African—Arabian zircon provinces. The East African–Arabian zircon province corresponds to the Gondwana super-fan system of Meinhold et al. (2013).

14.6

Conclusions

Although much of the ANS is isotopically juvenile, some previous studies reported pre-Neoproterozoic rocks from different occurrences of the ANS, implying a process driven primarily by reworking of pre-existing crust. In contrary, available Pb and Nd isotopic data do not support the existence of widespread substratal basement beneath the roots of the ANS. Whereas recent studies support local mixing of juvenile Neoproterozoic magmas with more isotopically evolved material, such as localized deposits of terrigenous sediment, highly attenuated crust, or continental microplates of pre-Neoproterozoic age. This is consistent with the widely variable initial Sr isotopic ratios, which argue for a widespread source of contamination.

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However, controversy persists on several important points, in particular, the rates at which juvenile crust is formed and the importance of the extraction of juvenile continental crust from the mantle versus the recycling of pre-existing crust. Two models have been proposed for the rate of continental growth through time in the ANS and are mostly based on geochronological and isotopic data from rocks and minerals. Model one assumes normal growth rates for continental crust in the ANS, i.e. simple arc-accretion model from 900 to 600 Ma. Whereas the second model proposes growth rates of an order of magnitude greater than the rates estimated for arcs since the end of the Palaeozoic. Therefore, understanding the detailed chronology of events and interpreting, the origin of Mesoproterozoic crustal components such as those in the ANS contributes to our understanding of this important period in Earth’s history.

References Abd El-Naby HH, Frisch W, Siebel W (2008) Tectonic–metamorphic evolution of the Wadi Hafafit culmination (central Eastern Desert, Egypt). Implication for Neoproterozoic core complex exhumation in NE Africa. Geol Acta 6:293–312 Abdel-Monem AA, Hurley PM (1979) U-Pb dating of zircon from psammitic gneisses, Wadi Abu Rosheid-Wadi Sikait area, Egypt. Institute of Applied Geology Bulletin, King Abdulaziz University Jeddah, Saudi. Arabia 2:165–170 Abdel-Rahman AM (1995) Tectonic-magmatic stages of shield evolution: the Pan-African belt in northeastern Egypt. Tectonophysics 242:223–240 Abdelrahman EM, Matheis G, Schandelmeier H, Karfis MA, Gadir I, El Khedir M (1993) Evolution of the Keraf back-arc basin: constraints on the Nubian Shield margin. In: Thorweihe U, Schandelmeier H (eds) Geoscientific research in NE Africa. Balkema, Amsterdam, pp 93–98 Abdelsalam MG, Stern RJ, Copeland P, Elfaki EM, Elhur B, Ibrahim FM (1998) The Neoproterozoic Keraf Suture in NE Sudan: sinistral transpression along the eastern margin of West Gondwana. J Geol 106:133–147 Abdelsalam MG, Liegeois JP, Stern RJ (2002) The Saharan Metacraton. J Afr Earth Sci 34:119–136 Abdelsalam M, Gao S, Liégeois J-P (2011) Upper mantle structure of the Saharan Metacraton. J Afr Earth Sci 60(5):328–336 Abu Anbar M, Solís Pichardo G, Hernández Bernal MS, Morales, Contreras J, Hernández Treviño T (2004) Sm-Nd and Rb-Sr isotopes of Feiran gneisses and amphibolites: evidences of Pre-Pan-African continental crust in Sinai, Egypt. In: 6th International Conference Geochemical, Alex Univ, Egypt, 727–745 Abu El-Enen MM, Whitehouse MJ (2013) The Feiran-Solaf metamorphic complex, Sinai, Egypt: geochronological and geochemical constraints on its evolution. Precambr Res 239:106–125 Agar RA, Stacey JS, Whitehouse MJ (1992) Evolution of the southern Afif terrane—a geochronologic study. Saudi Arabian Deputy Ministry for Mineral Resource Open File Report DGMR-OF-10–15, 41p Akaad MK, Noweir AM (1980) Geology and lithostratigraphy of the Arabian Desert Orogenic Belt of Egypt between Lat. 25 35′and 26 30′. Bull Inst Appl Geol King Abdulaziz Univ 3:127–135 Ali KA, Stern RJ, Manton WI, Johnson PR, Mukherjee SK (2010) Neoproterozoic diamictite in the Eastern Desert of Egypt and

354 Northern Saudi Arabia: evidence of *750 Ma glaciation in the Arabian Nubian Shield? Int J Earth Sci 99:705–726 Ali K, Andresen A, Manton WI, Stern RJ, Somar A, Maurice AE (2012) U-Pb zircon dating and Sr-Nd-Hf isotopic evidence to support a juvenile origin of the *634 Ma El Shalul granitic gneiss dome, Arabian-Nubian Shield. Geol Mag 149:783–797 Andresen A, El-Rus MA, Myhre PI, Boghdady GY, Corfu F (2009) U-Pb TIMS age constraints on the evolution of the Neoproterozoic Meatiq Gneiss Dome. Int J Earth Sci, Eastern Desert, Egypt. https:// doi.org/10.1007/s00531-007-0276-x Augland LE, Andresen A, Boghdady GY (2012) U-Pb ID-TIMS dating of igneous and metaigneous rocks from the El-Sibai area: time constraints on the tectonic evolution of the Central Eastern Desert, Egypt. Int J Earth Sci 101:25–37 Avigad D, Sandler A, Kolodner K, Stern R, McWilliams M, Miller N, Beyth M (2005) Mass-production of Cambro-Ordovician quartzrich sandstone as a consequence of chemical weathering of Pan-African terranes: environmental implications. Earth Planet Sci Lett 240:818–826 Avigad D, Gerdes A, Morag N, Bechstädt T (2012) Coupled U-Pb-Hf of detrital zircons of Cambrian sandstones from Morocco and Sardinia: Implications for provenance and Precambrian crustal evolution of North Africa. Gondwana Res 21(2–3):690–703 Baubron JC, Delfour J, Vialette J (1976) Geochronological measurements (Rb-Sr; K-Ar) on rocks of the Arabian Shield. French Bureau de Recherches Géologiques et Minières. Technical Record 76-JED-22, 152p Be’eri-Shlevin Y, Katzir Y, Whitehouse MJ, Kleinhanns IC (2009a) Contribution of pre Pan-African crust to formation of the Arabian-Nubian Shield: new secondary ionization mass spectrometry U-Pb and O studies of zircon. Geology 37:899–902 Be’eri-Shlevin Y, Katzir Y, Whitehouse MJ (2009b) Post-collisional tectonomagmatic evolution in the northern Arabian-Nubian Shield: time constraints from ion-probe U-Pb dating of zircon. J Geol Soc Lond 166:71–85 Be’eri-Shlevin Y, Eyal M, Eyal Y, Whitehouse MJ, Litvinovsky B (2012) The Sa’al volcano-sedimentary complex (Sinai, Egypt), a latest Mesoproterozoic volcanic arc in the northern Arabian Nubian Shield. Geology 40:403–406 Bea F, Montero P, Abu Anbar M, Talavera C (2011a) SHRIMP dating and Nd isotope geology of the Archean terranes of the Uweinat-Kamil inlier, Egypt-Sudan-Libya. Precambr Res 189:328–346 Bea F, Montero P, Abu Anbar M, Molina JF, Scarrow JH (2011b) The Bir Safsaf Precambrian inlier of South West Egypt revisited. A model for *1.5 Ga TDM late Pan-African granite generation by crustal reworking. Lithos 125:897–914 Begg GC, Griffin WL, Natapov LM, O’Reilly SY, Grand SP, O’Neill CJ, Hronsky JMA, Poudjom Djomani Y, Swain CJ, Deen T, Bowden P (2009) The lithospheric architecture of Africa: Seismic tomography, mantle petrology, and tectonic evolution. Geosphere 5:23–50 Berhe SM (1990) Ophiolites in Northeast and East Africa: implications for Proterozoic crustal growth. J Geol Soc Lond 147:41–57 Bielski M (1982) Stages in the Arabian-Nubian Massif in Sinai. Unpublished PhD thesis, Hebrew Univ., Jerusalem, 155p Bingen B, Jacobs J, Viola G, Henderson IHC, Skara Ø, Boyd R, Thomas RJ, Solli A, Key RM, Daudi EXF (2009) Geochronology of the Precambrian crust in the Mozambiquebelt in NE Mozambique and implications for Gondwana assembly. Precambr Res 170:231–255 Bregar M, Bauernhofer A, Pelz K, Kloetzli U, Fritz H, Neumayr P (2002) A late Neoproterozoic magmatic core complex in the Eastern Desert of Egypt: emplacement of granitoids in a wrench-tectonic setting. Precambr Res 118:59–82

H. H. A. El-Naby Burke K, MacGregor DS, Cameron NR (2003) Africa’s petroleum systems: four tectonic “Aces” in the past 600 million years. In: Arthur TJ, MacGregor DS, Cameron NR (eds) Geological Society. Special Publications, London, pp 21–60 Cahen L, Snelling NJ, Delhal J, Vail JR (1984) The geochronology and evolution of Africa. Clarendon Press, Oxford, p 512 Calvez JY, Kemp J (1982) Geochronological investigations in the Mahd adh Dhahab quadrangle, central Arabian Shield, Saudi Arabian Deputy Ministry for Mineral Resources Technical Record BRGM-TR-02-5, 41p Calvez JV, Delfour J, Feybesse JL (1985) 2000-million-year old inherited zircons in plutonic rocks from the Al Amar region: New evidence for an Early Proterozoic basement in the eastern Arabian Shield? Saudi Arabian Deputy Ministry for Mineral Resources Open File Report BRGM-OF-05-11, 28 p Castaing C, Feybesse JL, Thieblemont D, Triboulet C, Chevremont P (1994) Paleogeographical reconstructions of the Pan-African/ Brasiliano orogen:closure of an oceanic domain or intracontinental convergence between major blocks? Precambr Res 69:327–344 Cawood PA (2005) Terra Australis Orogen: rodinia breakup and development of the Pacific and Iapetus margins of Gondwana during the Neoproterozoic and Paleozoic. Earth-Sci Rev 69:249– 279 Claesson S, Pallister JS, Tatsumoto M (1984) Samarium-neodymium data on two late Proterozoic ophiolites of Saudi Arabia and implications for crustal and mantle evolution. Contrib Mineral Petrol 85:244–252 Clifford TN (1970) The structural framework of Africa. In: Clifford TN, Gass IG (eds) African magmatism and tectonics. Edinburgh, Oliver and Boyd, pp 1–26 Collins AS (2006) Madagascar and the amalgamation of Central Gondwana. Gondwana Res 9:3–16 Collins AS, Windley BF (2002) The tectonic evolution of central and northern Madagascar and its place in the final assembly of Gondwana. J Geol 110:325–339 Collins AS, Pisarevsky SA (2005) Amalgamating eastern Gondwana; the evolution of the Circum-Indian Orogens. Earth Sci Rev 71: 229–270 Collins GS, Melosh HJ, Osinski GS (2012) The impact-cratering process. Elements 8:25–30 Condie KC (2001) Continental growth during formation of Rodinia at l.35—0.9Ga. Gondwana Res 4:5–16 D’Amico C, Ibrahim HA, Sassi NP (1981) Outline of the Somali Basement. Geol Rundsch 70–3:882–896 De Waele B (2005) The Proterozoic geological history of the Irumide belt, Zambia. PhD Thesis, Curtin University of Technology, Perth, 468 p De Waele B, Kampunzu AB, Mapani BSE, Tembo F (2006) The Mesoproterozoic Irumide belt of Zambia. J Afr Earth Sci 46:36–70 De Waele B, Johnson SP, Pisarevsky SA (2008) Palaeoproterozoic to Neoproterozoic growth and evolution of the eastern Congo Craton: its role in the Rodinia puzzle. Precambr Res 160:127–141 Delfour J (1980) Geologic, tectonic and metallogenic evolution of the northern part of the Precambrian Arabian shield (Kingdom of Saudi Arabia). Bulletin Bureau des Récherches Géologiques et Minieres 11:1–20 Dixon TH, Golombek MP (1988) Late Precambrian crustal accretion rates in Northeast Africa and Arabia. Geology 16(11):991–994 El Gaby S (1983) Architecture of the Egyptian basement complex. Inter Basement Tectonics Assoc Publ 5:1–18 El Gaby S, Khudeir AA, Youssef M, Kamal El Din G (1988a) Low pressure metamorphism in Hammamat sediments at Wadi Um Had area, central Eastern Desert, Egypt. Bull Fac Sci Assuit Univ 17:51–71 El Gaby S, List F, Tehrani R (1988b) Geology, evolution and metallogenesis of the Pan-African belt in Egypt. In: El-Gaby S,

14

Evidence for Mesoproterozoic Components …

Greiling RO (eds) The Pan-African Belt of Northeast Africa and Adjacent Areas. Vieweg and Sohn, Weisbaden, pp 17–68 El Gaby S, List F, Tehrani R (1990) The basement complex of the Eastern Desert and Sinai. In: Said R (ed) The geology of Egypt. Balkema, Rotterdam, pp 175–200 El-Manharawy MA (1977) Geochronological investigations of some basement rocks in the central Eastern Desert, Egypt, between lat. 25° and 26°N. Unpublished PhD thesis, Cairo Univ., 211 p El-Shafei MK, Kusky TM (2003) Structural and tectonic evolution of the Neoproterozoic Feiran-Solaf metamorphic belt Sinai Peninsula: implications for the closure of the Mozambique ocean. Precambr Res 123:269–293 Fleck RJ, Coleman RG, Cornwall HR, Greenwood WR, Hadley DG, Prinz WC, Ratte JS, Schmidt DL (1973) Potassium-argon Geochronology of the Arabian Shield. United States Department of International Geological Survey, Saudi Arabia Project Report 165, 1–40 Fowler A, Khamees H, Dowidar H (2007) El-Sibai gneissic complex, Central Eastern Desert, Egypt: folded nappes and syn-kinematic gneissic granitoid sheets—not a core complex. J Afr Earth Sci 49:119–135 Fritz H, Tenczer V, Wallbrecher CA, Hauzenberger E, Hoinkes G, Muhongo S, Mogessie A (2005) Central Tanzanian tectonic map: a step forward to decipher Proterozoic structural events in the East AfricanOrogen. Tectonics 24, TC6013. http://dx.doi.org/10.1029/ 2005TC001796 Fritz H, Abdelsalam M, Ali KA, Bingen B, Collins AS, Fowler AR, Ghebreab W, Hauzenberger CA, Johnson PR, Kusky TM, Macey P, Muhongo S, Stern RJ, Viola G (2013) Orogen styles in the East African Orogen: a review of the Neoproterozoic to Cambrian tectonic evolution. J Afr Earth Sci 86:65–106 Garfunkel Z (2002) Early Paleozoic sediments of NE Africa and Arabia: products of continental-scale erosion, sediment transport and deposition. Isr J Earth Sci 51:135–156 Garson S, Shalaby IM (1976) Precambrian-lower Palaeozoic plate tectonics and metallogenesis in the Red Sea region. Geol Assoc Can Spec Pap 14:574–596 Garzanti E, Vermeesch P, Andò S, Vezzoli G, Valagussa M, Allen K, Kadi KA, Al-Juboury AIA (2013) Provenance and recycling of Arabian desert sand. Earth Sci Rev 120:1–19 Gellatly D (1960) Geology of the Las Dureh area, Burao district: Somalia Geological Survey Report no. 6, 40 p Gladkochub DP, Donskaya TV, Mazukabzov AM, Pisarevsky SA, Ernst RE, Stanevich AM (2016) The Mesoproterozoic mantle plume beneath the northern part of the Siberian craton. Russ Geol Geophys 57:672–686 Goodwin AM (1996) Principles of Precambrian geology. Academic Press, San Diego Grantham GH, Manhica ADST, Armstrong RA, Kruger FJ, Loubser M (2011) New SHRIMP, Rb/Sr and Sm/Nd isotope and whole rock chemical data from central Mozambique and western Dronning Maud Land, Antarctica: implications for the nature of the eastern margin of the Kalahari Craton and the amalgamation of Gondwana. J Afr Earth Sci 59:74–100 Grotzinger JP, Knoll AH (1996) Stromatolites in Precambrian carbonates: evolutionary mileposts or environmental dipsticks? Annu Rev Earth Planet Sci 27:313–358 Habib ME, Ahmed AA, El Nady OM (1985a) Two orogenies in the Meatiq area of the Central Eastern Desert, Egypt. Precambr Res 30:83–111 Habib ME, Ahmed AA, El Nady OM (1985b) Tectonic evolution of the Meatiq infrastructure, Central Eastern Desert, Egypt. Tectonics 4:613–627

355 Hamimi Z, El Amawy MA, Wetait M (1994) Geology and structural evolution of El Shalul dome and environs, Central Eastern Desert, Egypt. Egypt J Geol 38:575–595 Hargrove US, Stern RJ, Kimura J-I, Manton WI, Johnson PR (2006) How juvenile is the Arabian-Nubian Shield? Evidence from Nd isotopes and pre-Neoproterozoic inherited zircon in the Bi’r Umq suture zone, Saudi Arabia. Earth Planet Sci Lett 252:308–326 Harms U, Schandelmeier H, Darbyshire DPF (1990) Pan-African reworked early/middle Proterozoic crust in NE Africa west of the Nile: Sr and Nd isotope evidence. J Geol Soc Lond 147:859–872 Harris NBW, Hawkesworth CJ, Ries AC (1984) Crustal evolution in northeast and East Africa from model Nd ages. Nature 309:773–776 Hartnady CJH, Joubert P, Stowe CW (1985) Proterozoic crustal evolution in southwestern Africa. Episodes 8:236–244 Hassan MA, Hashad AH (1990) Precambrian of Egypt. In: Said R (ed) The geology of Egypt. Balkema, Rotterdam, pp 201–245 Hauzenberger CA, Bauernhofer AH, Hoinkes G, Wallbrecher E, Mathu EM (2004) Pan-African high pressure granulites from SE-Kenya: petrological and geothermobarometric evidence for a polycyclic evolution in the Mozambique belt. J Afr Earth Sci 40:245–268 Hauzenberger CA, Sommer S, Fritz H, Bauernhofer A, Kröner A, Hoinkes G, Wallbrecher E, Thöni M (2007) SHRIMP U-Pb zircon and Sm–Nd garnet ages from granulite facies basement of SE-Kenya: evidence for Neoproterozoicpolycyclic assembly of the Mozambique belt. J Geol Soc Lond 164:189–201 Hume WF (1934) Geology of Egypt 2, Part 1: The metamorphic rocks. Geol Surv Egypt, 293p Hunt JA (1958) Geology of the Adadleh area, Hargeisa and Berbera districts: Somaliland British Protectorate Geological Survey Report no. 2, 16 p Hunting Geology and Geophysics Ltd. (1974) Geology of the Jabal Al Uwaynat area, Libyan Arab Republic. Report to Industrial Research Centre, Department of Geological Research and Mining, Tripoli. Boreham Wood, U.K Huth A, Franz G, Schandelmeier H (1984) Magmatic and metamorphic rocks of NW Sudan: a reconnaissance survey. Berliner geowissenschaft liche Abhandlungen Reihe A50:7–21 Jacobs J, Thomas RJ (2004) Himalayan-type indenter-escape tectonics model for the southern part of the late Neoproterozoic-early Paleozoic East-African-Antarctic Orogen. Geology 32:721–724 Jacobs J, Fanning CM, Henjes-Kunst F, Olesch M, Paech H-J (1998) Continuation of the Mozambique into East Antarctica: Grenville-age metamorphism and polyphase Pan-African high grade events in central Dronning Maud Land. J Geol 106:385–406 Johnson PR, Woldehaimanot B (2003) Development of the Arabian– Nubian shield: perspectives on accretion and deformation in the northern East African Orogen and the assembly of Gondwana. In: Yoshida M, Windley BF, Dasgupta S (eds) Proterozoic East Gondwana: supercontinent assembly and breakup, vol 206. Geol. Soc. Lond. Spec. Publ.:290–325 Johnson PR, Andresen A, Collins AS, Fowler AR, Fritz H, Ghebreab W, Kusky T, Stern RJ (2011) Late Cryogenian-Ediacaran history of the Arabian-Nubian Shield: a review of depositional, plutonic, structural, and tectonic events in the closing stages of the northern East African Orogen. J Afr Earth Sci 61:167–232 Jöns N, Schenk V (2008) Relics of the Mozambique Ocean in the central East African Orogen: evidence from the Vohibory Block of southern Madagascar. J Metamorph Geol 26:17–28 Jöns N, Schenk V (2011) The ultrahigh temperature granulites of southern Madagascar in a polymetamorphic context: implications for the amalgamation of the Gondwana supercontinent. Eur J Mineral 23:127–156

356 Kamal El Din GM, Khudeir AA, Greiling RO (1992) Tectonic evolution of a Pan-African gneiss culmination, Gabal El-Sibai area, Central Eastern Desert, Egypt. Zbl Geol Paläont Teil 1(1991):2637– 2640 Kazmin V (1971) The Precambrian of Ethiopia. Nature Phys Sci Lond 280:176–177 Kazmin V (1975) The Precambrian of Ethiopia and some aspects of the geology of the Mozambique Belt. Bull Geophys Obs Addis Ababa 15:27–43 Kazmin V, Shifferaw A, Balch T (1978) The Ethiopian Basement: Stratigraphy and Possible Manner of Evolution. Geol Rundsch 67:531–546 Kemp J, Pellaton C, Calvez J V (1980) Geochronological investigatiom and geological history in the Precambrian of northwestern Saudi Arabia. Bureau des Recherches Gtologiques et Minieres Open File-01-1 Kennedy WO (1964) The structural differentiation of Africa in the Pan-African tectonic episode. In: 8th Ann. Rep., Res. Inst. Afr. Geol., Leeds Univ., 48–49 Kennedy WQ (1965) The influence of basement structure on the evolution of the coastal (Mesozoic and Tertiary) basins of Africa. Salt Basins around Africa. Institute of Petroleum, London, pp 7–16 Kennedy A, Johnson PR, Kattan FH (2004) SHRIMP geochronology in the northern Arabian Shield, Part 1: Data acquisition, Saudi Geological Survey Open File Report SGS-OF-2004-11, 28 p Khudeir AA, El-Gaby S, Greiling RO, Kamal El Din GM (1992) Geochemistry and tectonic significance of polymetamorphosed amphibolites in the Gabal Sibai window, Central Eastern Desert, Egypt. In: 1st International conference on geology of the Arab World, Cairo Univ, 461–476 Klerkx J (1980) Age and metamorphic evolution of the basement complex around Jabal al’Awaynat. In: Salim MJ, Busrewil MT (eds) The Geology of Libya 2. Academic Press, London, pp 901–906 Klerkx J, Deutsch S (1977) Resultats preliminaires obtenus par la methode Rb/Sr sur l’age des formations Precambriennes de la region d’Uweinat (Libye). Musee R. l’Afr. Centrale. Déepart. Géeol. Minéeral Rapp Ann, 83–94 Knoll AH, Semikhatov MA (1998) The genesis and time distribution of two distinctive Proterozoic stromatolitic microstructures. Palaios 13:407–421 Kröner A (1979) Pan-African mobile belts as evidence for a transitional tectonic regime from intraplate orogeny to plate margin orogeny. In: Al Shanti AMS (ed.) Evolution of the Arabian-Nubian Shield vol. 1, Institute of Applied Geology Bulletin 3:21–37 Kröner A (1984) Late Precambrian plate tectonics and orogeny: a need to redefine the term Pan-African. In: Klerkx J, Michot J (eds) African Geology. Tervuren, Belgium, pp 23–28 Kröner A (1985) Ophiolites and the evolution of tectonic boundaries in the late Proterozoic Arabian-Nubian shield of northeastern Africa and Arabia. Precambr Res 27:277–300 Kröner A, Sassi FP (1996) Evolution of the northern Somali basement: new constraints from zircon ages. J Afr Earth Sci 22:l–15 Kröner A, Stern RJ (2004) Pan-African orogeny. In: Selley RC, Cocks LRM, Plimer IR (eds) Encyclopedia of Geology, 1. Elsevier, Amsterdam, pp 1–12 Kröner A, Greiling R, Reischmann T, Hussein IM, Stern RJ, Durr S, Krüger J, Zimmer M (1987) Pan-African crustal evolution in the Nubian segment of Northeast Africa. Am Geophys Union Geodyn Ser 17:235–257 Kröner A, Kruger J, Rashwan AA (1994) Age and tectonic setting of granitoid gneisses in the Eastern Desert of Egypt and south-west Sinai. Geol Rundsch 83:502–513 Küster D, Liégeois J-P (2001) Sr, Nd isotopes and geochemistry of the Bayuda Desert high-grade metamorphic basement (Sudan): An

H. H. A. El-Naby early Pan-African oceanic convergent margin, not the edge of the East Saharan ghost craton? Precambr Res 109(1):1–23 Küster D, Liégeois J-P, Matukov D, Sergeev S, Lucassen F (2008) Zircon geochronology and Sr, Nd, Pb isotope geochemistry of granitoids from Bayuda Desert and Sabaloka (Sudan): Evidence for a Bayudian event (920–900 Ma) preceding the Pan-African orogenic cycle (860–590 Ma) at the eastern boundary of the Saharan Metacraton. Precambr Res 164(1–2):16–39 Kusky TM, Abdelsalam M, Stern RJ, Tucker RD (2003) Evolution of the East African and related Orogens, and the assembly of Gondwana. Precambr Res 123:82–85 Lancelot JR, Bosch D (1991) A Pan African age for the HP–HT granulite gneisses of Zabargad Island; implications for the early stages of the Red Sea rifting. Earth Planet Sci Lett 107:539–549 Lewin A, Meinhold G, Hinderer M, Dawit EL, Bussert R (2018) Provenance of sandstones in Ethiopia during Late Ordovician and Carboniferous-Permian Gondwana glaciations: Petrography and geochemistry of the Enticho Sandstone and the Edaga Arbi Glacials. Sediment Geol 375:188–202 Li X-H, Abd El-Rahman Y, Abu Anbar M, Li J, Ling X-X, Wu L-G, Masoud AE (2018) Old continental crust underlying juvenile oceanic arc: Evidence from northern Arabian-Nubian Shield, Egypt. Geophys Res Lett 45:3001–3008 Li ZX, Bogdanova SV, Collins AS, Davidson A, DeWaele B, Ernst RE, Fitzsimons ICW, Fuck RA, Gladkochub DP, Jacobs J, Karlstrom KE, Lu S, Natapov LM, Pease V, Pisarevsky SA, Thrane K, Vernikovsky V (2008) Assembly, configuration, and break-up history of Rodinia: a synthesis. Precambr Res 160:179–210 Liégeois JP, Stern RJ (2010) Sr–Nd isotopes and geochemistry of granite–gneiss complexes from the Meatiq and Hafafit domes, Eastern Desert, Egypt: no evidence for pre-Neoproterozoic crust. J Afr Earth Sci 57:31–40 Lundmark AM, Andresen A, Hassan MA, Augland LE, Abu El-Rus MA, Boghdady GY (2012) Repeated magmatic pulses in the East African Orogen of Central Eastern Desert, Egypt: an old idea supported by new evidence. Gondwana Res 22:227–237 Manya S, Maboko MAH (2003) Dating basaltic volcanism in the Neoarchaean Sukumaland Greenstone Belt of the Tanzania Craton using the Sm–Nd method: implications for the geological evolution of the Tanzania Craton. Precambr Res 121:35–45 Meert JG (2003) A synopsis of events related to the assembly of eastern Gondwana. Tectonophysics 362:1–40 Meinhold G, Morton AC, Avigad D (2013) New insights into peri-Gondwana paleogeography and the Gondwana super-fan system from detrital zircon U-Pb ages. Gondwana Res 23:661–665 Meinhold G, Bassis A, Hinderer M, Lewin A, Berndt J (2020) Detrital zircon provenance of north Gondwana Palaeozoic sandstones from Saudi Arabia. Geological Magazine Pub, 24 June 2020. https://doi. org/10.1017/s0016756820000576 Möller A, Mezger K, Schenk V (2000) U-Pb dating of metamorphic minerals: Pan-African metamorphism and prolonged slow cooling of high pressure granulites in Tanzania East Africa. Precambr Res 104:123–146 Neumayr P, Mogessie P, Hoinkes G, Puhl J (1996) Geological setting of the Meatiq metamorphic core complex in the Eastern Desert of Egypt based on amphibolite geochemistry. J Afr Earth Sci 23 (3):331–345 Pohl WL (1979) Metallogenic/Minerogenic Analysis—Contribution to the Differentiation between Mozambiquian Basement and Pan-African Superstructure in the Red Sea Region. Annals Geol Survey Egypt 9:32–44 Reymer A, Schubert G (1986) Rapid growth of some major segments of continental crust. Geology 14(4):299–302

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Ries AC, Shackleton RM, Graham RH, Fitches WR (1983) Pan-African structures, ophiolites and mélanges in the east Desert of Egypt: a traverse at 26 north. J Geol Soc Lond 140:75–95 Rino S, Kon Y, Sato W, Maruyama S, Santosh M, Zhao D (2008) The Grenvillian and Pan-African orogens: World’s largest orogenies through geologic time, and their implications on the origin of superplume. Gondwana Res 14:51–72 Rogers JJW, Santosh M (2004) Continents and Supercontinents. Oxford University Press, USA, p 289p Roobol MJ, Wright JV, Smith AL (1983) Calderas or gravity-slide structures in the Lesser Antilles island arc? J Volcanol Geotherm Res 19:121–134 Samuel MD, Be’eri-Shlevin Y, Azer MK, Whitehouse MJ, Moussa HE (2011) Provenance of conglomerate clasts from the volcanosedimentary sequence at Wadi Rutig in Southern Sinai, Egypt as revealed by SIMS U-Pb dating of zircon. Gondwana Res 20:450– 464 Sassi FP, Visona D, Ferrara G, Gatto GO, Ibrahim HA, Said AA, Tonarini S (1993) The crystalline basement of northern Somalia: lithostratigraphy and the sequence of events. In: Abbate E, Sagri M, Sassi FP (eds) Geology and mineral resources of Somalia and surrounding regions: relazioni e monographie agrarie subtropicali e tropicali nuova serie. Florence, Istituto Agronomico l’Otremare, pp 3–40 Schandelmeier H, Darbyshire DPF, Harms U, Richter A (1988) The east Saharan Craton: evidence for pre-Pan-African crust in NE Africa west of the Nile. In: El-Gaby S, Greiling RO (eds) The Pan-African belts of Northeast Africa and adjacent areas. Vieweg & Sohn, Wiesbaden, pp 69–94 Schandelmeier H, Richer A, Harms U, Abdel-Rahman EM (1990) Lithology and structure of the late Proterozoic Jebel Rahib fold-and-thrust belt (SW Sudan). Berliner Geowissen Abher (A) 120(1):15–30 Shang CK, Satir M, Morteani G, Taubald H (2010) Zircon and titanite age evidence for coeval granitization and migmatization of the early Middle and early Late Proterozoic Saharan Metacraton: example from the central north Sudan basement. J Afr Earth Sci 57:492–524 Schulter T (1997) Geology of East Africa. Beträge zur Regionale Geologie der Erde, Band 27, Gebrüder Borntraeger, Verslagsbuchhandlung, 484 pp Schürmann HME (1966) The Precambrian Along the Gulf of Suez and the Northern Part of the Red Sea. EJ Brill, Leiden, p 404p Shimron AE (1980a) Proterozoic island arc volcanism and sedimentation in Sinai. Precambr Res 12:437–458 Shimron AE (1980b) Late phase deformation and mylonite belts in Sinai: Pan African thrust fault tectonics. Geol Surv Isr Curr Res 1980:75–80 Shimron AE (1984a) Metamorphism and tectonism of a Pan-African terrain in southeastern Sinai—a discussion. Precambr Res 24:173– 188 Shimron AE (1984b) Evolution of the Kid Group, southeast Sinai Peninsula: thrusts, mélanges, and implications for accretionary tectonics during the late Proterozoic of the Arabian-Nubian Shield. Geology 12:242–247 Shimron AE, Furnes H, Roberts D, Bogoch R (1993a) Petrogenesis of the Late Proterozoic Sa’al Group—southern Sinai Peninsula. Geol Surv Israel Curr Res 8:24–29 Shimron AE, Bogoch R, Furnes H, Roberts D (1993b) The Sa’al Group: an ensialic island arc sequence in Sinai. In: Thorweihe U, Schandelmeier H (eds) Scientific research in Northeast Africa. Balkema, Rotterdam, pp 49–53 Siedner G, Shimron A, Pringle JR (1974) Age relationships in basement rocks of the Sinai Peninsula: Internat. Meet., Geochron. Cosmo. Isotop. Geol., Paris (Abstract)

357 Squire RJ, Campbell IH, Allen CM, Wilson CJL (2006) Did the Transgondwanan Supermountain trigger the explosive radiation of animals on Earth? Earth Planet Sci Lett 250:116–133 Stacey JS, Agar RA (1985) U-Pb isotopic evidence for the accretion of a continental micro-plate in the Zalm region of the Saudi Arabian Shield. J Geol Soc Lond 142:1189–1203 Stacey JS, Hedge CE (1984) Geochronologic and isotopic evidence for early Proterozoic crust in the eastern Arabian Shield. Geology 12:310–313 Stacey JS, Stoeser DB (1983) Distribution of oceanic and continental leads in the Arabian-Nubian Shield. Contrib Mineral Petrol 84:91– 105 Stacey JS, Doe BR, Roberts RJ (1980) Lead isotope study of mineralization in the Saudi Arabian Shield. Contrib Mineral Petrol 74:175–188 Stein M, Goldstein SL (1996) From plume head to continental lithosphere in the Arabian-Nubian Shield. Nature 382:773–778 Stein M, Hofmann AW (1994) Mantle plumes and episodic crustal growth. Nature 372:63–68 Stephan T, Kröner U, Romer RL (2019a) The pre-orogenic detrital zircon record of the Peri-Gondwanan crust. Geol Mag 156:281–307 Stephan T, Kröner U, Romer RL, Rösel D (2019b) From a bipartite Gondwanan shelf to an arcuate Variscan belt: the early Paleozoic evolution of northern Peri-Gondwana. Earth-Sci Rev 192:491–512 Stern RJ (1981) Petrogenesis and tectonic setting of late Precambrian ensimatic volcanic rocks, central Eastern Desert of Egypt. Precambr Res 16:195–230 Stern RJ (1994) Arc assembly and continental collision in the Neoproterozoic East African Orogen: implications for the consolidation of Gondwanaland. Annu Rev Earth Planet Sci 22:319–351 Stern RJ (2002) Crustal evolution in the East African Orogen: a neodymium isotopic perspective. J Afr Earth Sci 34:109–117 Stern RJ, Hedge CE (1985) Geochronologic and isotopic constraints on late Precambrian crustal evolution in the Eastern Desert of Egypt. Am J Sci 285:97–127 Stern RJ, Manton WI (1987) Age of Feiran basement rocks, Sinai: implications for late Precambrian evolution in the northern Arabian-Nubian Shield. J Geol Soc Lond 144:569–575 Stern RJ, Kröner A, Bender R, Reischmann T, Dawoud AS (1994) Precambrian basement around Wadi Halfa, Sudan: a new perspective on the evolution of the East Saharan Craton. Geol Rundschau 83:564–577 Stern RJ, Johnson PR, Kröner A, Yibas B (2004) Neoproterozoic ophiolites of the Arabian-Nubian Shield. In: Kusky TM (ed) Precambrian ophiolites and related rocks. Developments in Precambr Res 13:95–128 Stern RJ, Ali KA, Liegeois JP, Johnson PR, Kozdroj W, Kattan FH (2010) Distribution and significance of pre-Neoproterozoic zircons in juvenile Neoproterozoic igneous rocks of the Arabian-Nubian Shield. Am J Sci 310:791–811 Sultan M, Tucker RD, El Alfy Z, Attia R, Ragab A (1994) U-Pb (zircon) ages for the gneissic terrane west of the Nile, Southern Egypt. Geol Rundsch 83:514–522 Teklay M, Kröner A, Mezger K, Oberhänsli R (1998) Geochemistry, Pb-Pb single zircon ages and Nd-Sr isotope composition of Precambrian rocks from southern and eastern Ethiopia: implications for crustal evolution in East Africa. J Afr Earth Sci 26:207–227 Toteu SF, Penaye J, Djomani YP (2004) Geodynamic evolution of the Pan-African belt in Central Africa with special reference to Cameroon. Can J Earth Sci 41:73–85 Tretyakova AA, Kovachb VP, Shatagin KN (2016) Sources of Mesoproterozoic Igneous Rocks and Formation Time of the Continental Crust of the Kokchetav Massif (Northern Kazakhstan). Dokl Earth Sci 471–2:1312–1315

358 Vail JR (1983) Pan-African crustal accretion in NE Africa. J Afr Earth Sci 1:285–294 Viola G, Henderson IHC, Bingen B, Thomas RJ, Smethurst MA, deAzavedo S (2008) Growth and collapse of a deeply eroded orogen: insights from structural, geophysical, and geochronological constraints on the Pan-African evolution of NE Mozambique. Tectonics 27, TC5009. http://dx.doi.org/10.1029/ 2008TC002284 Warden AJ, Horkel AD (1984) The Geological Evolution of the NE-Branch of the Mozambique Belt (Kenya, Somalia, Ethiopia). Mitt österr Geol Ges 77:161–184 Whitehouse MJ, Windley BF, Ba-Bttat MAO, Fanning CM, Rex DC (1998) Crustal evolution and terrane correlation in the eastern Arabian Shield, Yemen: geochronological constraints. J Geol Soc Lond 155:281–295 Whitehouse MJ, Stoeser DB, Stacey JS (2001a) The Khida terrane geochronological and isotopic evidence for Paleoproterozoic and Archean crust in the eastern Arabian Shield of Saudi Arabia. Gondwana Res 4:200–202 Whitehouse MJ, Windley BF, Stoeser DB, Al-Khirbash S, Ba-Bttat MAO, Haider A (2001b) Precambrian basement character of Yemen

H. H. A. El-Naby and correlations with Saudi Araba and Somalia. Precambr Res 105:357–369 Windley BF, Whitehouse MJ, Ba-Bttat MAO (1996) Early Precambrian gneiss terranes and Pan-African island arcs in Yemen; crustal accretion of the eastern Arabian Shield. Geology 24:131–134 Yeshanew FG, Pease V, Whitehouse MJ, Al-Khirbash S (2015) Zircon U-Pb geochronology and Nd isotope systematics of the Abas terrane, Yemen: Implications for Neoproterozoic crust reworking events. Precambr Res 267:106–120 Yeshanew FG, Pease V, Abdelsalam MG, Whitehouse MJ (2017) Zircon U-Pb ages, d18O and whole-rock Nd isotopic compositions of the Dire Dawa Precambrian basement, Eastern Ethiopia: Implications for the assembly of Gondwana. J Geol Soc London 174:142–156 Zeh A, Gerdes A, Barton JM Jr (2009) Archean accretion and crustal evolution of the Kalahari Craton—the zircon age and Hf isotope record of granitic rocks from Barberton/Swaziland to the Francistown Arc. J Petrol 50(5):933–966 Zhang W, Pease V, Whitehouse MJ, El-Sankary MM, Shalaby MH (2019) Pre-Neoproterozoic basement evolution of southwestern Egypt. Int Geol Rev 61(15):1909–1926

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Najd Shear System in the Arabian-Nubian Shield Zakaria Hamimi and Abdel-Rahman Fowler

Abstract

The Arabian-Nubian Shield (ANS) was assembled from juvenile crust during a three-stage Neoproterozoic tectonic evolution involving: (1) intra-oceanic subduction and arc accretion stage, (2) orogenic extension stage, and (3) post-extensional compressional stage. Stage 1 is manifested by ophiolite-decorated arc-arc high-strain zones (suture zones) and calc-alkaline magmatism. The orogenic extensional tectonic stage generated dyke swarms, bimodal volcanism, molasse basins, A-type granite magmatism, low angle normal faulting (LANFs) and metamorphic gneiss complexes. The geological features attributed to this stage have been interpreted in terms of continental rifting, gravitational collapse, crustal and mantle delamination, transpression, escape/extrusion tectonism, and gravitational uplift. The post-extensional compressional stage is typified by dominantly NW–SE trending folds and thrusts, E–W transpression, and the N– S shortening zones. The Najd Fault System (NFS) (ca. 630–540 Ma) to be described in this chapter is attributed by some workers to the orogenic extension tectonic stage and by others to the post-extensional compressional stage. Earlier interpretation connects the NFS to the Najd Orogeny (570–520 Ma). The NFS is one of the largest transcurrent shear systems worldwide and deciphering its kinematic history adds considerably to our understanding of the cratonization of Gondwana, and specifically to mechanisms of exhumation of metamorphic complexes in the ANS. The NFS extends in a NW–SE direction across the Arabian Shield (e.g., Ajjaj, Qazaz, Ruwah, Ar Rikah, and Halaban strands of the NFS) for more than 1300 km Z. Hamimi Geology Department, Faulty of Science, Benha University, 13518 Benha, Egypt A.-R. Fowler (&) Geology Department, Faculty of Science, United Arab Emirates University, Al Ain, United Arab Emirates e-mail: [email protected]

(*400 km wide) and continues beneath Phanerozoic cover in Yemen. The NFS is believed to extend into the Nubian Shield (Egyptian Eastern Desert and Sinai). The dominant sense of shearing along the NW–SE trending Najd megashears is sinistral, however, evidence exists for an earlier phase of dextral slip. NE- (to ENE-) oriented shear zones (e.g. the Ad-Damm, Fatima, Idfu-Mersa Alam, Qena-Safaga shear zones) could be Najd-related conjugates or earlier fault systems. The shear and volume strain aspects of Najd shears are described, as are the stress controls on the brittle evolution of Najd faults. The role of Najd brittle structures in hydrothermal mineral deposits and ground water flow patterns are also covered in this chapter. Keywords








Najd shears Transpression Escape tectonics Gravitational collapse Gneiss complexes

15.1




Introduction

The Najd Fault System (NFS) dissecting the Arabian-Nubian Shield (ANS) represents one of the most prominent examples of post-amalgamation wrench fault systems, which are also recorded in other shield areas. It is the best exposed and may be the largest pre‐Mesozoic zone of transcurrent faulting on Earth (Stern 1985). The geology, structural setting, tectonic evolution, origin, dimensions, geochronology, kinematic history, sense of shearing, modeling, mechanism of development, role in exhumation of the core complexes and gneissic domes, as well as its control on post-amalgamation volcano-sedimentary basins, mineral resources (e.g. gold, copper, zinc, silver, and lead) and hydrothermal alterations, have all been the subjects of numerous detailed studies since the 1960s (Brown and Jackson 1960; Delfour 1970, 1977, 1979; Brown and

© The Author(s), under exclusive license to Springer Nature Switzerland AG 2021 Z. Hamimi et al. (eds.), The Geology of the Arabian-Nubian Shield, Regional Geology Reviews, https://doi.org/10.1007/978-3-030-72995-0_15

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Coleman 1972; Hadley 1974, 1976; Fleck et al. 1976; Howland 1979; Moore 1979a, b; Schmidt et al. 1979; Davies 1984; Schmidt and Brown 1984; Stacey and Agar 1985; Stern 1985; Burke and Sengör 1986; Agar 1986, 1987; Sultan et al. 1988; Andre 1989; Stern et al. 1989, Berhe 1990; Duncan et al. 1990; Quick 1991; Bonovia and Chorowicz 1992; Abdelsalam and Stern 1996; Sengör and Natalin 1996; Ghebreab 1998; Fritz and Messner 1999; Blasband et al. 2000; El-Rabaa et al. 2001; Matsah and Kusky 2001; Johnson et al. 2001; Bregar et al. 2002; Fritz et al. 2002; Genna et al. 2002; Brooijmans et al. 2003; Kusky and Matsah 2003; Fowler and El-Kalioubi 2004; Abdeen and Greiling 2005; Shalaby et al. 2006; Fowler et al. 2007; Abdeen et al. 2008; Abd El-Naby et al. 2008;; Mogren et al. 2008; Abu-Alam and Stüwe 2009; Fowler and Osman 2009; Abd El-Wahed 2010; Abu-Alam et al. 2010; Johnson et al. 2011; Zoheir 2011; Abdelazeem et al. 2013; Fritz et al. 2013; Abdeen et al. 2014; Ali-Bik et al. 2014; Hamimi et al. 2014a,b; Meyer et al. 2014; Stüwe et al. 2014; Fowler et al. 2015; Abd El-Wahed et al. 2016; Hassan et al. 2016a, b; Makroum 2017; Sultan et al. 2017; Emam et al. 2018; Hagag et al. 2018; Stern 2018, Zoheir et al. 2018; Hamimi et al. 2018, 2019; El-Fakharani et al. 2019; Fowler and Hamimi 2020; Hamimi and Abd El-Wahed 2020; Hamimi et al. 2020). This contribution summarizes the results of the previously published literature and discusses field relations and structural interpretations of the Najd system. It provides an account of the nature and timing of Najd deformation and reviews proposed tectonic models, describes geometry, and kinematics and discusses along-strike deformational style variation and other enigmatic aspects of the NFS. We aim to offer an up-to-date synthesis on this global scale tectonic element that developed during the closing stages of Neoproterozoic crustal evolution and played an important role in shaping the ANS.

15.2

The Arabian-Nubian Shield Within the East African Orogen

The rugged mountainous belts exposed on either side of the Red Sea form counterparts: the Arabian Shield in western Saudi Arabia and Yemen, and the Nubian Shield in Egypt, and including Sudan, Eritrea, Ethiopia, Somalia, and Kenya. The Arabian Shield and the Nubian Shield are together known as the Arabian-Nubian Shield (ANS). The ANS is regarded as the largest tract of juvenile continental crust on Earth (Patchett and Chase 2002). It is dominated by Neoproterozoic rocks (*980 Ma–535 Ma; Nance et al. 2014) comprising greenschist volcanic, plutonic, and sedimentary belts, and represents the northern continuation and the upper crustal equivalent of the Mozambique belt (MB), which encompasses high-grade metamorphic belts. Both the ANS

Fig. 15.1 The East African-Antarctic Orogen (EAAO) with the Arabian-Nubian Shield (ANS) at its northern end. The Najd Fault System (NFS) is located in the northeastern part of the ANS (from Jacobs et al. 2008)

and the MB are components of the East African Orogen (EAO) that tectonically evolved during the Pan-African Orogeny (900–550 Ma) (Stern 1994) (Fig. 15.1) The EAO is regarded as a northern part of the larger East African-Antarctic Orogen (EAAO) (Jacobs et al. 2008) that extends from the EAO into Eastern Antarctica and more completely describes the collision zone between East and West Gondwana (Fig. 15.1). The EAO is a Neoproterozoic– early Cambrian mobile belt and is considered one of the Earth’s great collision zones, where East and West Gondwana collided to form the supercontinent ‘Greater Gondwana’ or ‘Pannotia’ at the end of Neoproterozoic time (Stern 2002). It is comparable with the *7500 km long Cenozoic Alpine–Himalayan orogenic system (Fritz et al. 2013). The formation and construction of the EAO impacted atmospheric oxygenation (Och and Shields-Zhou 2012) and

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Fig. 15.2 Aerial view of brittle, ESE-trending sinistral Riedel shear (dashed line) displacing paragneiss and orthogneiss of the Hijaz gneiss belt along the sinistral Ruwah fault zone and a thrust sheet of Bani Ghayy group marble. The Riedel fault is centered about 20°50′06″ N, 43°57′13″ E

consequently influenced the evolution of the higher organisms (Squire et al. 2006). The EAO originated within the Mozambique Ocean (Stern 1994). It connected the 900– 720 Ma break-up of Rodinia (Li et al. 2008) and the *650– 530 Ma multiphase assembly of Gondwana, concurrently with the East African Orogeny (Collins and Pisarevsky 2005). It can be recognized from space (Fig. 15.2) and is about 6000 km long, extending from North Africa to Antarctica through 350 million years of evolution (Stern 1994). A reconstruction made by Jacobs et al. (1998, 2008) extended the EAO further south to East Antarctica (Fig. 15.1). Several lines of evidence indicate that the ANS crust was juvenile (mantle-derived), formed throughout the accretion of intra-oceanic arcs (Kröner et al. 1987) to form a mosaic of island-arc and minor continental tectonostratigraphic terranes. The terranes amalgamated to form the ANS predating the E–W collision that formed the supercontinent Gondwana (Abdelsalam and Stern 1996). The ANS terranes

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were accreted onto the Saharan metacraton and are separated along ophiolite decorated arc-arc and/or arc-continent sutures as well as major high strain zones. The arc-arc sutures separate the arc terranes that formed between *800 and 700 Ma and the arc-continent suture juxtaposes the ANS from West Gondwana that formed between 700 and 650 Ma (Stoeser and Camp 1985; Kröner et al. 1992) (Fig. 15.3). The ANS evolution from juvenile oceanic lithosphere to island arcs and continental lithosphere is evidenced lithologically by the occurrence of ophiolite belts, quartzofeldspathic and carbonate sedimentary sequences, and calc-alkaline igneous rocks, besides other lithologies. On the other hand, the ANS had a prolonged polyphase deformation history, recorded in arc-arc collision in the Cryogenian (780–700 Ma), followed by a complex history of terrane collision (680–640 Ma), Najd wrench tectonics (630–540 Ma) and tectonic escape (590–540 Ma) in the Ediacaran (Shackleton et al. 1980; El-Ramly et al. 1984; Greiling et al. 1994; Fritz et al. 1996; Johnson et al. 2011). The evolution of the ANS and the entire EAO was and still is a matter of much discussion. Four tectonic models have been proposed; infracrustal, Turkic-type, hot-spot, and arc-assembly-orogenic models. Among these models, the arc-assembly model is currently widely accepted in the geoscientific community. This model was proposed by Vail (1985) and Stoeser and Camp (1985), and was subsequently modified by Stern (1994) and Stern and Abdelsalam (1998). It comprises four main geological stages that manifested a Wilsonian orogenic cycle (Stern 1994). These are: (1) rifting and break-up of Rodinia (900–850 Ma), (2) sea-floor spreading, arc and back-arc basin formation, and terrane accretion (870–690 Ma), (3) continental collision and formation of the EAO (630–600 Ma), and (4) continued shortening, escape tectonics and orogenic collapse (590– 540 Ma). Such stages define a single supercontinent cycle, beginning with the fragmentation of end-Mesoproterozoic Rodinia Supercontinent in the early Neoproterozoic (Hoffman 1999).

15.2.1 The Najd Orogeny in the Tectonic Frame of the Arabian Shield The Arabian Shield tectonic events were originally divided by Brown (1972) and Brown and Coleman (1972) into the earlier Hijaz Orogeny, associated with the island arc and subduction stage, and the later Najd Orogeny, marked by regional wrench faulting. Moore and Al Shanti (1973) suggested 660–570 Ma as the duration of the Najd Orogeny. Greenwood et al. (1976), in their work in the central and southern areas of the shield, renamed the first event as the ‘Hijaz Orogenic Cycle’, followed by Najd Faulting. The same early events in the eastern and central areas of the

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Fig. 15.3 Map of the main terranes of the ANS and the sutures at their boundaries (from Kröner and Stern 2004)

shield were termed ‘Najd Orogenic Cycle’ by Delfour (1982). The shortage of precise geochronological data before the 1990s hindered the setting up of unified orogenic nomenclature. Stoeser et al. (1985) bracketed the Hijaz Orogenic Cycle to >950–680 Ma, and stated that it was terminated by the Nabitah Orogeny (680–630 Ma), with the consequent generation of strong N–S to NE–SW structural trends throughout the shield. Stoeser et al. divided the following Najd Orogeny (630–550 Ma) into two stages. The first stage was dominated by the development and inversion of the foreland Murdama basins (615–570 Ma), involving sedimentation, deformation, metamorphism, and granite intrusion. This was succeeded by the second stage in which

the sinistral Najd Fault System formed, with strike-slip related basins being the sites of sedimentation for the Jibalah Group (590–530 Ma). Detailed geochronology by Kozdrój et al. (2018) confirms the range of Najd faulting activity as 636–577 Ma. Two chronological tables for the Najd in the frame of Arabian Shield events are shown in Fig. 15.4. In Fig. 15.4a Johnson and Woldehaimanot’s (2003) chart shows Najd faulting as a separate but time-overlapping event with respect to the Nabitah Orogeny (shearing associated with collision of the Afif terrane with the Asir and other terranes). In Fig. 15.4b Genna et al.’s (2002) chart combines the arc terrane collisions with the Najd in the continuous Nabitah Orogeny.

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Fig. 15.4 Chronological tables showing the time frame for Najd faulting and other major events in the Arabian-Nubian Shield. a Redrafted from Johnson and Woldehaimanot (2003). The time ranges for the main shearing events are shown as red broken wavy lines. Sedimentation and deformation of basins are shown. The arrows show the orientation of the regional forces, with the N direction upwards. b From Genna et al. (2002). The Nabitah Orogeny is shown to include the Najd faulting event. 690 Ma marks the beginning of the Pan-African events, and the closure of the oceanic domain. 590–530 Ma marks the beginning and end of the regional extensional event

15.3

The Najd Fault System

The NFS is a complex left-handed (sinistral) strike-slip fault system that crosses the Pan-African belt of the Arabian Shield in a NW-SE direction. The pioneering work revealing this system was by Brown and Jackson (1960). Brown and Coleman (1972) introduced the term Najd Fault System to the ANS lexicon. This system, which was developed during

a limited interval of the Neoproterozoic crustal evolution (620–540 Ma, Stern 1985; 580–530 Ma, Fleck et al. 1980, 630–600 Ma, Stacey and Agar 1985), extends for 1300 km (maximum width *400 km) in western Saudi Arabia. It is inferred to continue into the Eastern Desert of Egypt and Sinai (Abuzeid 1984; Sultan et al. 1988). Perhaps the most southerly possible equivalents of Najd are the sinistral NWto NNW-trending Aswa and Nyangere Faults that lie near the southern tip of the ANS in northern Tanzania, and

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Fig. 15.5 Map of the ANS showing the main Najd faults of the Arabian Shield and Egyptian Nubian Shield (from Johnson et al. 2011). Gneiss domes and molasse basins associated with the Najd are also shown

continue in the NW direction to separate the Saharan and Congo cratons (Katumwehe et al. 2016). These faults may have had a suture origin and show a significant pure shear component. They have been dated to around 593 Ma (Fritz et al. 2013). Najd Fault System matches fault trends in South Yemen (Andre and Blodget 1984; Andre 1989), which has led many workers (Brown 1972; Moore and Al-Shanti 1979) to suppose its extension beneath and beyond Al-Rub’ Al-Khali, yielding a total length of at least 2000 km. Secondary structures associated with the NFS include oblique-, normal- and reverse-faults, pull-aparts, thrusts, folds, and dyke swarms. Although the NFS is restricted to a broad zone (*400 km width), four to five principal strands

(Halaban-Zarghat, Ar Rika, Ajjaj, and Qazaz Faults) reflect the history of NFS deformation and kinematic movement (Moore and Al-Shanti 1979; Schmidt et al. 1979; Davies 1984; Stern 1985) (Fig. 15.5). The remarkable and long-lasting kinematic stage along the NFS was transcurrent with sinistral sense of shearing. Left lateral displacement may reach up to 300 km, as documented by Brown and Coleman (1972), Schmidt et al. (1979), Davies (1984) and many others. However, the tectonic significance and kinematic history of the NFS remain matters of much debate. Some workers have reported right-lateral shearing and others believe that the NFS kinematic history switched from dextral at the beginning of movement to sinistral. Yet another group

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Najd Shear System in the Arabian-Nubian Shield

believes that the main sinistral shearing stage is overprinted by an attenuated dextral one. In this context, Stacey and Agar (1985) stated that the NFS was active between 630 and 600 Ma as a dextral system and then as a sinistral system up to *530 Ma. The same conclusion was given by Agar (1987) who indicated that in the southwest of the Najd system, near Zalm, initial faulting was dextral and began earlier than formerly thought (Kusky and Matsah 2003). Emplacement of a plutonic complex was controlled by Najd fractures of dextral geometry and displacements. However, there is nearly a total agreement that the transcurrent shearing was accompanied by compression one, making the NFS a typical example of a transpression zone. Major dextral faults with N-S to NNE-SSW orientation, e.g. Ad Damm Fault (Davies 1984; Hamimi et al. 2014) have also been thought to be conjugates of the sinistral Najd, in this context (Fig. 15.5). In the following text, the characteristics of the main fault strands of the NFS (Fig. 15.5) are presented before consideration of the faults on the system scale.

15.3.1 The Main Najd Megashears This section highlights the characteristic features of Ruwah, Ar Rika, Qazaz, and Halaban-Zarghat fault zones which are the most prominent of the NFS in the Arabian Shield (Fig. 15.5).

15.3.1.1 Ruwah Fault Zone The Ruwah Fault Zone (RFZ) is a conspicuous NW-trending high strain zone extending from the Asir terrane in the west to the Afif terrane in the east. The sense of shearing along the RFZ is sinistral. Various lithologies can be observed in the vicinity of the RFZ, including amphibolite mylonitic gneisses (ortho- and para-), schists, gabbros, and metaultramafics, along with slightly deformed greenschist facies metavolcaniclastics. These units are steeply dipping and elongated parallel to the main fault zone. The longest single gneiss belt in the Arabian Shield is the Hajizah–Tin gneiss, located along the RFZ in the eastern shield (Johnson et al. 2011). Johnson et al. (2011) believed that the RFZ originated as a strand of the major Hulayfah-Ad Dafinah-Ruwah suture zone (HARS) at the western/southwestern margin of the Afif terrane, and was reactived by Najd shearing. It is worth mentioning that the HARS itself is a subvertical sinistral transpressional zone (600 km long and 5–30 km width) extending in a gentle concave-to-the NE curve across the northeastern Arabian Shield. It has been recognized as the northern continuation of the Nabitah suture, however, this claim has been challenged by Johnson and Kattan (2001). These authors believed that the southern extension of the HARS is concealed beneath the Phanerozoic cover to the east of the Tathlith-Malahah terrane

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in the Arabian Shield. A granodiorite gneiss protolith age of 683 ± 9 Ma is considered by Stacey and Agar (1985) to represent the age of suturing (age of deformation) of the Hulayfah-Ad Dafinah-Ruwah suture zone. Near its northeastern continuation, the RFZ controls the basin boundaries of the Bani Ghayy Group (*650–620 Ma; Johnson et al. 2011). The marble belonging to this group occurs as an overthrust sheet. The deformation/metamorphism terminated by 592 ± 4 Ma, the time of intrusion of the Najd-related hornblende-biotite monzogranite (Stoeser and Stacey 1988). The RFZ is transected on the northwest by a brittle fault affecting the Siham Group of the Afif terrane and is also dissected in the south by the Hasak semi-brittle shear zone.

15.3.1.2 Ar Rika Fault Zone Based on their origin, time, tectonic implications, and related structures, Johnson (1996) classified the NFS into two major groups, namely Ar Rika and Ruwah fault zones. Ar Rika Fault Zone (ARFZ) encompasses the Kirsh, An Nakhil, Wajiyah, Hammadat, Ajjaj, and Qazaz gneiss-schist belts, along with overprinting brittle faults. Both groups are separated by relatively less deformed areas that contain NW–SE translation planes, along which bending of the N–S Nabitah structures is obvious (Mogren et al. 2008). Ar Rika Fault Zone (ARFZ) is a conspicuous structural element traversing the Afif terrane in the Arabian Shield. It represents part of an en échelon set of shear zones extending from the Qazaz-Ajjaj Shear Zone in the northwest to the Ar Rika shear zone in the southeast. This set is known as the Qazaz-Ar Rika shear zone, which hosts the ARFZ, the Hamadat, and Kirsh gneiss belts. Shearing on this set is thought to have occurred between *635 Ma and *573 Ma (Calvez et al. 1984; Kennedy et al. 2010). Johnson et al. (2011) believed that the initial movements on the ARFZ followed the deposition of the post-Nabitah collision Murdama group (*650–625 Ma), and subsequent movements took place after the intrusion of 611 Ma granitoids. In the southeastern part of the ARFZ, the gneisses and pegmatites outcropping in the Al Hawriyah antiform have SHRIMP U-Pb ages demonstrating zircon growth at 623 Ma, 602 Ma, and 589 Ma, and probable metamorphism at 600 Ma (Kennedy et al. 2005, 2011). These ages led Johnson et al. (2011) to consider that slip on the ARFZ occurred between 625 and 590 Ma. The obtained 40Ar/39Ar isochron ages (557 ± 15 Ma) given by Al-Saleh (2012) from the Kirsh biotite paragneiss may represent the exhumation age of Kirsh gneiss. 15.3.1.3 Qazaz Fault Zone The Qazaz fault zone (QFZ) represents the northwestern continuation of the abovementioned ARFZ. Careful examination reveals that both faults are characterized by a sinistral left-stepping en échelon fault set. The QFZ (3–4 km wide) is

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a prominent NW- (to WNW-) trending shear zone belonging to the NFS, traversing the Midyan terrane and showing steep mylonitic foliations, gently plunging stretching lineations, and sinistral transpressional shearing. El-Fakharani et al. (2019) carried out integrated remote sensing and field studies to define the tectonic styles of Hanabiq, Ajjaj, and Qazaz zones. They came to the conclusion that both Qazaz and Ajjaj shear zones formed in a top-to-NW or top-to-WNW sinistral transpressional regime. The striking feature along the strike of the QFZ is its splitting into two noticeable branches. Shearing along both branches is bracketed between *630 and 580 Ma (Calvez et al. 1984; Kennedy et al. 2009). The two branches enclose a triangular dome-like structure, known as the Qazaz dome. The dome is dominated in its inner core by medium- to high-grade gneisses, flanked by low-grade mylonitic zone, low-grade metapelites, conglomerates, and volcanic rocks of the Thalbah and Bayda Groups. Meyer et al. (2014) studied in some detail the Qazaz metamorphic dome and described a new mechanism whereby core complexes can be exhumed along crustal-scale strike-slip fault systems that accommodated crustal shortening. They indicated that the dome was exhumed along a gently dipping jog in a crustal-scale vertical strike-slip fault zone that caused more than 25 km of exhumation of lower crustal rocks by 30 km of lateral motion. In the same context, Hassan et al. (2016a) concluded that the Qazaz metamorphic dome recording peak metamorphic conditions of 560–640 °C and 7–8 kbar and exhumation occurs at about 25 km due to the activity of the NFS.

15.3.1.4 Ajjaj Fault Zone The Ajjaj fault zone (AJFZ) is an impressive transcurrent shear zone measuring some 160 km long (*30 km wide) in the northwestern part of the Arabian Shield. The AJFZ is not parallel with the neighboring QFZ and other faults of the NFS. It is generally an anastomosing structural element, but to the south of Baladiyah complex it possesses an average WNW-ESE strike. A N–S trending dextral strand of the AJFZ, the Hanabiq shear zone, is also recorded. As in the QFZ and most of the Najd-related megashears in Western Arabia and Central Eastern Desert of Egypt, there is a geometric relationship between the AJFZ and a gneiss dome, known as the Hamadat anticlinorium, consisting of NW-plunging tight folds and SE-plunging more open folds (Genna et al. 2002). The dome (*100 km long diagonal) extends parallel to the AJFZ. It is occupied by garnetiferous gneisses and low-grade greenschist up to amphibolite facies grade rocks of the volcano sedimentary Zaam and Bayda groups, along with igneous rocks (Abu-Alam et al. 2014). El-Fakharani et al. (2019) carried out detailed microstructural analysis of the Ajjaj–Qazaz–Hanabiq shear zones in order to determine the shear sense and the

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kinematic history. The authors identified four phases of deformation. The oldest (Hijaz) D1 structures are represented by WNW to E–W striking foliations and interfolial folds. D2 microstructures are represented by NNW- to N–S trending foliation and folds parallel to the Nabitah suture. The earlier structures are overprinted by the transpressional Najd Fault System. The geometry and overprinting relations of microstructures (including sigmoidal structures, mantled porphyroclastic structures, asymmetric boudins, tension gashes, fish-shaped structures, oblique foliation, bookshelf sliding structures, slickenlines, and C–S tectonites) led the authors to conclude top-to-the-NW and top-to-the-WNW slip shearing in the AJFZ, and top-to-the-NNE in the Hanabiq shear zone.

15.3.1.5 Halaban-Zarghat Fault Zone The Halaban-Zarghat fault zone (HZFZ) is the most prominent Najd-related structural feature in the Arabian Shield. The HZFZ can easily be traced using Landsat and ASTER imagery. Various post-amalgamation volcano sedimentary depositional pull-apart basins are genetically connected to the HZFZ, being formed at releasing bends. Among these basins is the Jifan basin that includes the type section of the Jibalah Group, which consists of conglomerate, sandstone, shale, volcaniclastics, and limestone. High-grade mylonitic gneiss belts are also recorded along the HZFZ. The HZFZ has a prolonged history of deformation, and the early kinematic history was dextral, while the latest most significant movement was sinistral (Matsah and Kusky 2001). The early dextral sense along the HZFZ is proposed to be 624.9 ± 4.2 Ma (U/Pb zircon dating of rhyolite erupting into the Jifan basin). This age is interpreted by Matsah and Kusky (2001) as predating the E-W Gondwana collision where the accreted terranes caught between these two supercontinents attempted to escape toward an oceanic free face to the north. The 576.6 ± 5.3 Ma U/Pb zircon minimum age given for a felsic dyke intruding the Jibalah group is considered by these authors to represent the end of collision and escape (extrusion or indentation) tectonics. Johnson et al. (2011) postulated that the Halaban part of the HZFZ was initiated at about 680 Ma as a suture zone at the eastern border of Afif composite terrane, and its trajectory was turned to the NW during a later stage transcurrent deformation. On the other hand, Cole and Hedge (1986) believed that the movement along the HZFZ is principally sinistral postdating the deposition of the Murdama group, although minor dextral slip existed predating the emplacement of 621 Ma Ediacaran diorite. Although the kinematic history had terminated at about 595 Ma (the deposition time of the Jurdhawiyah group; Kennedy et al. 2004, 2005), the HZFZ was active during and after deposition of the Jibalah group at *590–560.

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15.4

Evidence for the Timing of Deformation on the Najd Fault System

On the system scale, the activity of Najd faulting is broadly accepted as extending from *630 Ma to *540 Ma (Stern 1985; Stoeser and Camp 1985; Stacey and Agar 1986; Agar 1987; Sultan et al. 1988; Johnson et al. 2011) (Fig. 15.4a). An alternative range given by Johnson and Woldehaimanot (2003) is 640–560 Ma (Fig. 15.4a). These boundary dates are defined by the Najd displacement of the youngest sutures (end of Nabitah suturing at 640 Ma, Idsas suture *640 Ma), and the post-activity andesitic flows (K-Ar age 540 ± 18 Ma, Brown 1972) in the Jibalah basins located along the Najd faults. However, up to the end of the 1970s the popular view was that Najd faulting began after 600 Ma, i.e. extended from 590 or 580 Ma to 530 Ma (Fleck et al. 1976; Schmidt et al. 1979). This narrower period was based on the age range of felsic dykes intruded along Najd faults (Fleck et al. 1976) and the observation that Najd fault activity followed the intrusion of the majority of the Najd granites (Stoeser et al. 1985), and followed the early stages of deformation of the 615–595 Ma Murdama basins (Stacey et al. 1984). The end of Najd was also limited by the existence of undisturbed platform sandstones and limestones with early Middle Cambrian fauna, overlying the northwestern end of the Najd system in Jordan (Schmidt et al. 1979). Fritz et al. (2013) date the end of Najd activity via E-W dyke intrusions (590–545 Ma) that heralded the onset

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of a N-S extensional tectonic regime in the final stages of the Najd event. Early evidence in support of extending the Najd activity to pre-600 Ma times was given by Agar (1984, 1987) who described the Najd-assisted intrusion of the Haml granodiorite in the Zalm Quadrangle, subsequently dated by Stacey and Agar (1985) at 637 ± 11 Ma. Bani Ghayy Group molasse sediments were deposited between *650 Ma and 620 Ma in strike-slip pull-apart basins along the Najd Ruwah Fault Zone at right-stepping releasing bends. Stacey et al. (1984) found that Murdama basin sediments (*650– 620 Ma) were deformed in the early stages of Najd. A growing database of precise ages for synkinematic gneissic intrusions along strands of the Najd in the northern parts of the shield added further to the evidence for pre-600 Ma Najd activity. Davies (1982, 1984) noted that plutons as old as 625 Ma were elongated and concordant within the Najd shear zones but circular outside these zones, indication existence of Najd-related shears at *625 Ma. Evidence for the time range of Najd activity on the major shear zones of this system are set out in Johnson et al. (2011, their Table 3) and are summarized in Table 15.1. The generally accepted date of the onset of Najd activity at 630 Ma has been extended to > 680 Ma by Nehlig et al. (2001, 2002) who regarded the Nabitah stage as continuous with Najd and drew attention to the sigmoid appearance of the N-S Nabitah faults at their intersections with the NW-SE Najd trends. For the Ruwah Fault, Nehlig et al. (2001)

Table 15.1 Constraints on the timing of Najd Fault System activity Najd onset must postdate/predate …

Najd end must postdate/predate …

Probable range

HalabanZarghat

Postdate 625 ± 4 Ma rhyolite that floors dextral pull-apart basin (Kusky and Matsah 2003) Postdate 621 ± 7 Ma plutons (Kusky and Matsah 2003; Cole and Hedge 1986) Predate 620 ± 7 Ma pluton (Cole and Hedge 1986) Predate 584 ± 10 Ma oldest seds in pull-apart basin (Nettle 2009)

Predate undeformed 577 ± 5 Ma felsite dyke (Kusky and Matsah 2003) Postdate 573 ± 8 Ma synkinematic pluton (Kennedy et al. 2005) Postdate 574 ± 28 Ma pluton (Calvez et al. 1984)

625–560 Ma

Ar Rika

Predate 600 ± 20 Ma gneiss (Kennedy et al. 2005)

Postdate 579 ± 19 Ma faulted pluton (Calvez et al. 1983)

625–590 Ma

Ruwah

Predate 620 Ma end of Bani Ghayy deposition (Stacey and Agar 1985)

Predate 591 ± 6 Ma undeformed pluton (cited by Kellogg et al. 1986). Predate 592 ± 4 Ma undeformed pluton (Stoeser and Stacey 1988)

620–588 Ma

QazazAjjaj

Postdate 670 ± 10 Ma gneiss (Kennedy et al. 2005) Postdate 630 ± 19 Ma pluton (Kemp et al. 1980) Postdate 626 ± 4 Ma pluton (Kennedy et al. 2011) [Ajjaj] predate 601 ± 3 Ma synkinematic pluton (Hassan et al. 2016a)

Predate 573 ± 6 Ma undeformed dykes (Kennedy et al. 2011). Postdate 575 ± 10 Ma deformed pluton (Kennedy et al. 2011). Postdate 560 ± 4 Ma zircon in Jibalah basin faulted seds (Kennedy et al. 2011) [Ajjaj] predate 581 ± 4 Ma undeformed pluton (Hassan et al. 2016a)

635–573 Ma [Ajjaj] 604–581 Ma (Hassan et al. 2016a) [Ajjaj] 630–560 Ma (Robinson et al. 2017) [Qazaz] 635–573 Ma (Robinson et al. 2017)

Ad Damm

Postdate 620 Ma deformed pluton (Fleck and Hadley 1982)

620–540 Ma

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recognized the Subay granodiorite (683 ± 9 Ma) as synkinematic during Najd shearing. Stacey and Agar (1985) placed this pluton in the early stages of the Nabitah Orogeny. Similarly, Nehlig et al. (2001) characterized the 680– 630 Ma elongated gneiss domes along the Ar Rika Faults as intrusions into the early stage Najd. For the Qazaz Fault, Nehlig et al. (2002) provided a similar interpretation of the synkinematic tonalites along this fault, dated at 676 ± 4 Ma and 672 ± 30 Ma by Hedge (1984).

15.5

Kinematic History of the NFS

During the more than fifty years of study of the NFS, the detailed field mapping of the fault zones and areas between the faults, combined with regional aeromagnetic geophysical surveys and strain profiles constructed across the individual faults, has revealed much about the kinematic history of the NFS. While the NFS is understood to be a major sinistral strike-slip system, it has been suspected from early surveys that there was a minor dextral slip-sense phase in the history of this system. Early evidence for local dextral slip-sense on the Najd was given by Davies (1980) from northwestern parts of the shield. Agar (1984) and Stacey and Agar (1985) presented evidence for an early stage of dextral slip on Najd in the form of dextral sense of pull-aparts in the Zalm area (Fig. 15.6) (including faults of the Ruwah strand of the Najd), in which the pre-620 Ma Bani Ghayy molasse sediments were deposited. They also noted the dextral-shear assisted mechanism of emplacement of the *637 Ma Haml granodiorite along the same fault, which was followed by the earliest signs of a switch to sinistral displacement at *620 Ma. Agar (1987) further suggested that early dextral switching to sinistral shear sense may characterize the entire NFS, though did not explain the cause of the slip sense

Fig. 15.6 Chronostratigraphic chart for the Zalm area (from Stacey and Agar 1985). An early dextral stage for the Najd shearing (Ruwah Fault) is shown as 640–620 Ma, corresponding to the time of deposition of the Bani Ghayy Group, and intrusion of the Hufayrah Complex granite and Haml suite granodiorite

change. The dextral stage of slip on the Ruwah Fault has been rejected by Johnson et al. (2001), and the general model of an early dextral stage of NW-SE trending Najd has been challenged by several workers. Moore and Al Shanti (1979) suggested that the Najd dextral shears could be antithetic Riedel (R’) shears that were rotated towards parallelism with the Najd, or perhaps scissor-fault activity, and Brown et al. (1989) expressed similar doubts about the existence of early dextral Najd. Subsequent studies have confirmed the early dextral slip sense on most Najd faults, e.g. Kusky and Matsah (2003) for the Halaban-Zarghat Fault; and Baggazi et al. (2019) for the Hanabiq shear zone. The total displacement across the 400 km wide NFS was investigated by Brown (1972), resulting in his estimate of 240 km of sinistral strike-slip displacement, based on offsets of the ophiolite-decorated N-S to NE-SW oriented suture zones. Davies’ (1984) figure for total displacement was 280 ± 120 km. Moore and Al Shanti (1979) noted that some oblique-slip was common, and this must result in variable vertical (up and down) displacements on the shears. Davies’ (1982) strain profile across the Ajjaj shear zone yielded 60 ± 20 km sinistral strike-slip displacement and 5 ± 1.5 km vertical displacement. Meyer et al. (2014) concluded that there had been 60 km of sinistral displacement along the Qazaz shear zone, with the first 30 km resulting in a 25 km uplift of a core complex bounded by the elements of the shear zone. Hassan et al. (2016a, b) estimated an exhumation from 58 km to 44 km depth (14 km vertical displacement) of the Hamadat complex along the Ajjaj shear zone during shear zone activity. The longest faults are known to have displacements of a few tens of kms, whereas a smaller net slip of a few kms is typical for most individual faults (Moore and Al Shanti 1979). Delfour (1977) illustrated sinistral displacements in the Nuqra area (between the Halaban-Zarghat and Ar Rika fault zones) of

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up to 40 km on some individual faults, though the majority of small faults showed only about 1 km of slip (Fig. 15.7). Moore and Al Shanti (1979) noted that there is some evidence for a displacement-fault length proportionality relationship, i.e. longer faults showed greater displacement than shorter faults. He also reported that the maximum displacement occurred at approximately the midpoint of the fault, with slip decreasing towards the fault tips. The fracture network alongside individual throughgoing Najd faults is commonly about 10 km wide (Fig. 15.7). These brittle fault zones are bordered by ductile sheared wallrocks. The ductile margins extend away from the faults to form broad ductile foliated zones an order of magnitude wider (100 km) than the brittle fault zones. The association of ductile foliated zone and brittle fault zones, however, is variable, with the more northwesterly Najd shears (e.g. the Ajjaj) having brittle fault zones lying within intensely foliated greenschist or higher grade, broad, enveloping ductile sheared zones (Stern 1985). The Najd shears in the northeast (e.g. the Ad Dawadmi) show little ductile-associated deformation (Stern 1985). The following section describes the characteristics of the ductile deformed rocks associated with the brittle Najd faults, and the partitioning of displacement between them, as revealed by metamorphic and field studies, strain analysis, remote sensing, and geophysics.

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15.6

Metamorphic Conditions of Najd Shearing

One of the intriguing features of the Najd Fault System is the common lateral transition from ductile-dominated to brittle-dominated shearing that roughly correlates with decreased metamorphic temperatures within the shear zones and in its surroundings (Johnson et al. 2011). Most studies of the ductile dimensions of the NFS have been located in the northwestern shears, in particular, the Qazaz, Ajjaj, and Hanabiq shear zones. The comparative variations in metamorphic grade between inner and outer parts of these Najd shears and their enclosing rocks have been recently reported by Meyer et al. (2014), Abu-Alam et al. (2014), Hassan et al. (2016a, b). Metamorphic grade maps of Ajjaj-Hamadat– Hanabiq by Hassan et al. (2016a, b) show that the least deformed parts of the shear zones are broadly low metamorphic grade, with elongate zones of medium grade gneissic bodies along the shears (Fig. 15.8). These medium grade rocks are cut by thin high-grade metamorphic belts that correspond to high strain zones in the inner parts of the gneiss (Davies 1982). Similar observations are reported for the Qazaz shear zone by Meyer et al. (2014), where amphibolite facies gneissic bodies in the inner part of the shear zone are enclosed by high strain mylonitic enclosing shears showing a syn-shearing greenschist overprint, which

Fig. 15.7 Map of Nuqra Quadrangle showing the typical magnitude of brittle displacements (few 10s km) on Najd faults (modified from Moore 1979a). The displacement is demonstrated on the ophiolitic bodies. These are simplified by marking their midline in yellow. Fracture networks associated with the main faults are also about 10 km wide

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Fig. 15.8 Metamorphic map of the Ajjaj shear zone (from Hassan et al. 2016a, b)

separate the inner shear from the greenschist wallrocks (Fig. 15.9). Along the Ajjaj and Hanabiq, the geobarometry suggests shearing took place at up to 60 km depth in this broadly foliated zone (Hassan et al. 2016a). Significant depths of shearing (16 km, based on metamorphic P of about 4.5 kbar) were reported for the Baladiya gneissic complex in the eastern part of the Qazaz shear zone by Abu-Alam et al. (2014), matching similar depths reported by Meyer et al. (2014). The rocks beyond the shear zones showed metamorphism at depths of only 1.5–3.5 km indicating substantial uplift nearer the shear zones.

15.7

Strain Analysis Across the Najd Shear Zones

Strain studies of the NFS and its individual shears have involved finite strain estimates, shear strain intensity maps, and shear strain profiles (Davies 1982, 1984; Hamimi et al. 2014a; Stüwe et al. 2014; Baggazi et al. 2019; Kahal et al. 2019). Early views of the deformation mechanism of the NFS by Moore and Al Shanti (1979) was that simple shear was sufficient to explain the evolution of the system. Davies’ (1982) strain study of the Ajjaj shear zone used the changing angles between fold axial planes and shear zone margin, from outside the shear zone (*60o) to inside the shear zone (*15o), and rotation of dykes that had intruded along tension gashes to compute both the shear strain magnitude c

and volume change dV. He contoured the values dV on a map to show that there had been volume loss inside the shear zones (partly by early pressure solution effects) and volume gain in the wallrocks outside the shear zones (partly by space-filling granite intrusions and vein systems). This necessitated a model including a pure shear component added to the simple shear of the early models. He also contoured c values to find that the inner parts of the shears had highest shear strains. Integration of the c-distance curve in shear strain profiles across the shears gave 60 ± 20 km ductile displacement. Davies (1984) followed the same approach for four Najd faults (Ajjaj, Arja/Jifn, Ad Dawadmi, Ad Damm) to give a total sinistral ductile displacement of 280 ± 120 km (Fig. 15.10b). He found that all of the shears in this regional profile deviated from simple shear. The amount of shortening (pure shear) across the zones increased towards the east. For the individual shears, the ductile displacement varied from 15.5 ± 7 km (Ad Damm fault) to 180 ± 55 km for the Arja plus Jifn faults, with the last showing additional 40 km of brittle slip. For the regional profile, the spacing of c maxima and the amount of shortening decreased maybe due to increased ductility towards the east. He concluded that the shear zones in the east may represent higher sections through Najd, with the implication that the zone of ductile strain increased in width with depth. He found the Najd faults to parallel the trend of shear strain contours (Fig. 15.10a) and also found volume loss within shear zones partly compensated by volume loss outside shear zones (Fig. 15.10b).

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Fig. 15.9 Metamorphic map of the Qazaz shear zone and gneissic dome (from Meyer et al. 2014)

AlKashghari (2017) conducted strain profiles across the Qazaz, Ajjaj, and Hanabiq shear zones. He found that the shear zones had a high pure shear component (88%) making them pure-shear dominated transpression zones. The shear zones corresponded to subhorizontal shortening, with maximum extension parallel to the shears. The gneiss domes however were characterized by subhorizontal foliations indicating vertical shortening. Strain studies have been conducted for the NE trending dextral Fatima Shear Zone by Kassem and Hamimi (2018) and Baggazi et al. (2019). These will be described in Sect. 15.11 on the NE-trending Najd faults.

15.8

Remote Sensing and Geophysical Studies of the Najd Fault System

Remote sensing studies of the NFS have used enhanced Landsat images to delineate foliation trends and the surface extent of foliated zones enclosing the NFS brittle faults (Duncan et al. 1990; AlWash and Zakir 1992; Divi and Zakir 2001). The continuity of the shear zones and faults has also been mapped using Remote Sensing (RS). Duncan et al. (1990) successfully used remote sensing to estimate 25– 30 km displacement in the western part of the Ajjaj shear

zone. AlWash and Zakir (1992), Hamimi et al. (2014a), Samkari (2015), Bishta et al. (2015) and El-Sawy and El-Shafei (2019) have used processed Landsat, ASTER, and TIRS imagery to outline fault-related lineaments of the Ad Damm fault in the Jeddah terrane (Fig. 15.11). More recently remote sensing has been used to geologically map the Qazaz–Ajjaj–Hanabiq shear branches in the Midyan terrane by AlKashghari (2017). Extracted lineaments from Zouaghi et al.’s (2019) aeromagnetic data revealed fine details of the N-S Farwah dextral fault in the southern part of the Arabian Shield. Aeromagnetic datasets for the Saudi shield were mainly collected during 1962–1983 under the supervision of the USGS and BRGM. Moore and Al Shanti (1979) commented on the magnetic signature of the Najd Faults being significantly broader and laterally continuous that the surface expression of the faults, indicating wider zones of Najd shearing at depth. Moore and Al Shanti (1979) commented that the Najd faults at the surface are underlain by much more continuous probably ductile or partly ductile shear zones that extend for hundreds of kilometres in length. He proposed that the brittle faults merged downwards into these shear zones. Mogren et al. (2008) found that the Ar Rika and Ruwah faults were associated with magnetic anomalies (range 50–300 nT) (Fig. 15.12a) and 20 km wide negative

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Fig. 15.10 Strain maps and profiles across major Najd faults derived from Davies (1984). a Contoured values of shear strain (c) across the Ajjaj shear zone. The sinistral faults are approximately parallel to the trend of the shear strain contours. b Shear strain (c) profiles with red line generalized pattern of strain variations for four groups of Najd shears (from left to right: the Ajjaj; Arja and Jifn; Ad Dawadmi; and Ad Damm faults). The figures below, with blue line generalized patterns, are the variations in the y-axis strain along the profiles of the same four shears. High c within the shear zones is associated with y-axis strains suggestive of volume loss. Low c outside the shear zones is associated with y-axis strains suggestive of volume gain

gravity anomalies ranging from 20 to 30 mGal that was probably due to lower density of crushed rock. Nehlig et al. (2001, 2002) and Stern and Johnson (2010) illustrate 10– 20 km wide magnetic anomalies with intensities of about 140 nT outlining the Najd faults, particularly where gneissic domes lie within the fault zones (Fig. 15.12b). The broad

ductile zones enclosing the Najd faults are also evident by the width of the zone of shear curvature affecting the regional N-S magnetic anomalies as they approach the Najd faults (Fig. 15.12a) (Mogren 2004), and the cross-cutting and displacement of N–S magnetic lineations of the earlier terrane collision event (Johnson and Stewart 1995).

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Fig. 15.11 Remote sensing image (left) and corresponding geological map (right) of the Ad Damm Fault area (from Hamimi et al. 2014a)

(b) (a)

100 km

Fig. 15.12 Aeromagnetic maps of the Arabian Shield showing N-S magnetic anomalies cross-cut by NW-SE trending anomalies associated with Najd faulting. a Map of the area of the Ruwah and Ar Rika Faults (from Mogren et al. 2008), showing curvature of the magnetic trends as they approach the Najd faults. (b) Map of the entire Arabian Shield (from Nehlig et al. 2001). Even at this scale the Najd faults are visible and are expressed as magnetic anomalies 10–20 km wide

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Both NE- and NW-trending Najd fault elements were revealed in the aeromagnetic studies of Al-Saud (2008a) in the Jeddah terrane, though these were not referred to as conjugates. The Najd faults are also characterized by gravity anomalies (positive) due to abundant mafic dykes along the faults which are believed to pass downwards into larger mafic intrusions (Gettings et al. 1986). This explains simultaneous gravity and magnetic anomalies along the faults. Subsurface modeling of the Najd fault system was performed by Mogren (2004) using gravity profiles.

15.9

Brittle Evolution of the Najd Fault Zone

The mechanism of propagation of the Najd faults and the origin of the second-order structures associated with them (e.g. splays, secondary faults, etc.) have been studied by Moore and Al Shanti (1973), who applied stress trajectory analysis to a group of NW-trending sinistral Najd faults in the Ad Dawadmi area in the eastern part of the Arabian Shield. In their study area, the Najd faults terminated within a large intrusion of isotropic adamellite, allowing a simple stress analysis of the fault tips and recognition of second-order structures. The r1 and r3 trajectories were guided by the numerous dykes and extension fissures (normal to r3) in the wall rocks. From the r1 and r2 trajectories, the surfaces of maximum shear stress s could also be constructed (Fig. 15.13a). Moore and Al Shanti’s (1973) observations included that (a) dykes were intruded along curved principal stress (r1) trajectories (i.e. the dyke curvature was not strain related) (Fig. 15.13b), (b) dykes intruded different fracture sets controlled by the manner in which sporadically active faults relieved stresses on the fractures, (c) second-order structures were splays and horsetail structures, and secondary faults often followed the trend of maximum shear stress trajectories (Fig. 15.13a), and (d) the Najd faults propagated along splays from the SE to the NW. They also found that the style of second order structures was most likely due to uniaxial stress with r1 at a large angle (35–45o) to the main faults. The gentle pitch of striations indicated that the principal stresses were not perfectly vertical or horizontal. Further stress studies by Moore and Al Shanti (1979) explored the interaction of hydrothermal fluids with the Najd Fault stress regime. They noted that the most complex areas for secondary structures were in zones between en échelon faults where compression and tension could occur on opposite sides of the faults. Braided normal faults adjacent to the main faults could produce pop-up and sag (basin) structures. Complex fold patterns could also result with folds inclined to the main faults or parallel to them. Broadly,

though, the regional dykes described E-W orientations as expected for NW-SE sinistral shear sense. Moore and Al Shanti (1979) pointed out the importance of fracture systems not physically connected to the main faults (i.e. brittle structures originating from the same stresses that formed the faults, but not branching from the fault planes) (Fig. 15.13c). These independent fractures included Riedel fractures, particularly R synthetic (sinistral) faults at 15° counterclockwise to the main Najd faults. The antithetic R’ Riedel faults were N-S to NE-SW trending (75° or more clockwise from the main fault) and were much less well developed. Moore explained the role of R in the development of sheeted main faults, pinnate zones. He noted that the main faults could cut through their secondary structures or be displaced by them. While not physically attached to the main faults, the Riedels appeared to lie above active Najd faults. Stress tensor studies of the Ad Dawadmi area by El-Sawy and Masrouhi (2019) revealed that the left-stepping sinistral en échelon main faults defined a pull-apart structure (Fig. 15.14).

15.10

Role of Transpression in the Evolution of the Najd Fault System

Davies’ (1982, 1984) early strain studies of the rocks enclosing the Najd faults showed conclusively that the strain patterns could not be explained by progressive simple shear mechanism alone. Davies (1984) added a pure shear (actually dilational) component to the model mechanism. Sanderson and Marchini (1984) also observed that the Najd fault system showed numerous features consistent with transpression at strike-slip fault bends and terminations. On the other hand, Moore and Al Shanti (1979) reported that (R) Riedel shears in the Najd system lay typically 15° to the main shears. According to Naylor et al.’s (1986) sandbox experiments, that modeled the progressive development of faults in sedimentary cover above a basement strike-slip fault, the Riedel (R) shears are about 17° to the basement fault if there is no transpression. In Naylor et al.’s experiments, with significant transpression, R was found to lie at higher angles (35°) to the basement fault. Nevertheless, they found that wide fault zones, such as those of the Najd, are more likely to form in transpression. Naylor et al. (1986) noted that transpression was also marked by development of P shears that link the Riedel faults to produce fault lenses, such as those characterizing the Najd system, however, there are very few mentions of P shears in the literature of Najd (e.g. Zoheir and Moritz 2014). The above deviations of Najd Riedel shear orientations from those expected for transpressive shear zones can be explained as follows. Jones and Tanner (1995) argued that

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Najd Shear System in the Arabian-Nubian Shield

Fig. 15.13 a Pattern of conjugate maximum shear stress trajectories around the terminations of faults in the Ad Dawadmi area (from Moore and Al Shanti 1973). The trajectories show the expected orientations of secondary faults. Red lines mark the actual mapped faults in this area, which closely follow the expected fault trends. b Principal stress (r1 and r3) trajectories on either side of the Hamrur Fault in the Ad Dawadmi area (from Moore and Al Shanti 1973). The pattern of dykes in the area suggests that they intruded along the fault plane and along the curved r1 trajectories. c Generalized model depicting the secondary faults (normal, thrust, and strike-slip) and folds associated with the main Najd faults. C denotes compression structures. D denotes dilatational or extensional structures. R denotes synthetic Riedel shears (from Moore 1979a). Yellow diamond symbol shows the typical dilatational sites preferred for veins mineralization in the Najd fractures (see text Sect. 15.6 for discussion)

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Fig. 15.14 The two faults of the Ad Dawadmi area interpreted as elements of a sinistral pull apart structure (from El-Sawy and Masrouhi 2019)

Fig. 15.15 Model from Jones and Tanner (1995) showing different degrees of partitioning with depth for brittle slip and ductile shear

although the upper and lower crust may both experience transpression, the lower crust rheology may favor homogeneous transpression, while the upper brittle crust promotes partition of the transpression into brittle shears (taking up much of the simple shear strain), while the blocks between

the brittle shears show mainly pure shear strain (Fig. 15.15). If this is the case then the angle of convergence would appear to be different between the upper and lower crust, and the upper crust faults may appear much less transpressive. The tendency for faults to show regional curvature can also explain changes in the degree of transpression along them. Johnson and Kattan (2001) describe the curved boundary of the Afif terrane with the Hijaz, Jiddah, and Asir terranes, as a sinistral transpressive shear zone composed of the Ruwah, Ad Dafinah, and Hulayfah fault zones. The Hulayfah-Ad Dafinah faults approach N-S orientation and the early folds associated with them suggest NW-SE compression and vertical extension leading to steeply plunging stretching lineations (Fig. 15.16). The Ruwah fault is NW-SE trending and has stretching lineations that are subhorizontal, and the deformation was simple shear. So the NW-SE compression produced simple shear on the Ruwah but sinistral transpression on the Hulayfah–Ad Dafinah. Al-Saleh and Kassem (2012) considered the Kirsch dome as a metamorphic core complex exhumed by local extension along a releasing bend in the Ar Rika Fault, following a transpressional orogeny.

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Fig. 15.16 Diagram from Johnson and Kattan (2001) showing how the plunge of stretching lineations changes from subhorizontal (on the NE-trending Ruwah Fault) to subvertical (along the N-S trending Ad Dafinah Fault)

Transpression effects are particularly well developed in the Qazaz, Ajjaj, and Hanabiq shear zones. Genna et al. (2002) developed a model to explain the foliations and lineations in the Qazaz gneiss dome as a result of transpression in the enclosing shear zone. They illustrated the vertical changes in the strain ellipsoid in the gneiss dome, so that there were vertical foliations at depth and subhorizontal foliations in the upper parts of the dome (Fig. 15.17). Stretching lineations were parallel to the dome axis and intense in the core of the dome. They described this configuration as representing simple shear dominance over pure shear, with upward expulsion of material forming the anticlinal geometry of the gneiss dome. Genna et al.’s structural model could also reflect upward changes in ratios of lateral to vertical extrusion (Jones et al. 1997), related to the upward convex geometry of the boundary shears. While subhorizontal stretching lineations are usually taken to

indicate simple shear dominated transpression (Fossen 2010), Jones et al. (1997) found that subhorizontal stretching lineations can also form in pure shear dominated transpression if both lateral and vertical extrusion (unconfined transpression) have occurred. Studies of the Ajjaj shear zone have found that it experienced transpression between 605 and 580 Ma (Stüwe et al. 2014; Meyer et al. 2014; Hassan et al. 2016a). Meyer et al. (2014) studied exhumation mechanisms of gneiss domes in the Saudi Shield and compared these mechanisms with those of the CED core complexes (Figs. 15.18a, b). They noted that the CED complexes were exhumed by extension (Fig. 15.18a), while the Saudi ones were also exhumed in transpression (Fig. 15.18b). Some evidence for this was the existence of HT metamorphic rocks in the enclosing shear zones, as well as in the gneissic cores. Similar high-pressure kyanite-bearing paragneisses envelope the Kirsh gneiss

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transpression was followed by N-directed extension and tectonic escape (*620–580 Ma). Abu-Alam et al. (2014) found oblique sinistral transpression to correlate with the collision of the accreted terranes with the Keraf suture within a wider time range of 650–560 Ma. This is termed their D3 event and was followed by D4 tectonic escape. Robinson et al. (2017) date transpression on the Najd to around 620 Ma.

15.11

Fig. 15.17 Model of strain variations in transpressive Najd-hosted gneiss domes (from Genna et al. 2002). In this model, simple-shear dominates pure-shear deformation. Vertical foliations at depth give way to subhorizontal foliations in higher parts of the gneiss dome

along the Ar Rika Fault, also suggesting transpressionassisted uplift was involved in gneiss dome exhumation (Al-Saleh 2012). Meyer et al. (2014) and Meyer (2015) describe the kinematics of branching ductile shears of the Najd Qazaz shear zone (Fig. 15.18c), and reported significant vertical shear displacements as well as lateral ones. They emphasized the importance of rheological changes associated with strain softening for the nucleation of shear zones. Meyer (2015) noted that the ratio of the pure shear (shortening) and simple shear components in the shear zones changed during the transpression history of the shear zones. AlKashghari (2017) found the Qazaz, Ajjaj, and Hanabiq shear zones to be pure-shear dominated transpression zones, with shear sense and orientation controlled by exhumation of high-grade gneissic complexes. El-Fakharani et al. (2019), studied the same shear zones and concluded that sinistral transpression accompanied the D3 (Najd) deformation along the Qazaz and Ajjaj, while dextral sense was recorded along the more northerly trending Hanabiq, due to exhumation of the gneissic bodies. The timing of Najd transpression was dated by Johnson et al. (2011) to *630 Ma and correlated with the E-W or ENE-WSW convergence of east and west Gondwana. The

N-S to NE-SW and Other Trending Dextral Strike-Slip Faults of the Southern Saudi Shield

The NE-oriented faults in southern Arabia have been previously suggested to be contemporaneous conjugates shears of the NFS (Davies 1984). These faults exhibit dextral sense of shearing and are typified by Ad-Damm and Fatima fault zones in Western Arabia. Other fault zones in the Eastern Desert of Egypt, namely the Qena-Safaga and Idfu-Mersa Alam faults have also been included in this set. The NW-SE trending Ruwah fault zone is commonly presented as the SW margin of the Najd Fault system in the Arabian Shield. To the southwest of this fault are a number of N-S to NE-SW trending faults, typically with dextral shear sense, which are of questionable relationship to the Najd system (Fig. 15.19a). The Nabitah-Hamdah and Junaynah are N-S trending fault structures, and appear to be distinctly older than the Najd system as they are cut by the Najd NW-SE trending faults. The Nabitah is generally considered as a major suture, marking the collision of the Afif terrane with the western arc terranes at about 680– 640 Ma (Johnson et al. 2011), though Flowerdew et al. (2013) have argued that it is a minor, rather than major, suture. Dextral transpression on the Nabitah suture has been described by Quick (1991), Johnson and Kattan (2001), Johnson and Woldehaimanot (2003) and Johnson et al. (2011). The deformation history of the N-S trending Yiba fault zone has been described by Hamimi et al. (2013, 2014b). This fault experienced dextral transpression soon after the Nabitah event (*630 Ma) during which the Ablah basin sediments were deposited in a pull-apart structure. A more southerly part of the Ablah graben was studied by Bamousa (2013). He found that the N-S thrusts and folds affecting the Ablah Group were related to N-S dextral shearing, and were separated in time from the NW-trending sinistral Najd faults by E-W trending dextral faults. The ophiolitic bodies decorating the NE-SW trending Bir Umq fault zone indicate it is also a suture zone, dated to 780– 750 Ma (Abdelsalam and Stern 1996; Johnson et al. 2011). Dextral transpression accompanied amalgamation of the Hijaz and Jiddah terranes along this suture (Wipfler 1996; Johnson and Woldehaimanot 2003; Johnson et al. 2004).

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Fig. 15.18 Models for the exhumation of the Qazaz gneissic complex (from Meyer et al. 2014). a typical core complex exhumation mechanism involving extension. b Meyer et al.’s (2014) model of core complex exhumation in a strike-slip environment. c stages of exhumation of the Qazaz gneissic complex (from earliest at top to latest at bottom)

The Ad Damm shear zone (Fig. 15.19a, b) is one of the most prominent NE-striking fault structures in the Arabian Shield. It separates the Jiddah terrane on its NW from the Asir terrane to its SE, and has been considered as a suture zone or a major wrench fault system (Matsah et al. 2004). The presence of ophiolite in the (NW) hangingwall is consistent with origins as a suture (Matsah et al. 2004). Early deformation events involved NW-SE compression that generated a N-plunging megafold in the psammitic gneisses of the hangingwall, during amphibolite facies metamorphism, and thrusted these rocks over the footwall granitic gneisses (El-Sawy and El-Shafei 2019). The following event was a dextral shearing or dextral transpression (Hamimi et al. 2014a; Samkari 2015; El-Sawy and El-Shafei 2019) which occurred at about 620–540 Ma (Hamimi et al. 2014a; El-Sawy and El-Shafei 2019). The dextral transpression resulted in greenschist facies shears and mylonites cutting through the amphibolite facies rocks. The NE-trending major shear zone along Wadi Fatima in the western side of the shield (Fig. 15.19a, c) has also

experienced several deformation stages (Hamimi et al. 2012). The first two events were associated with terrane amalgamation, involving NW-SE compression during the assembly of the Jeddah and Asir terranes (795 Ma) (Abd-Allah et al. 2014). The presence of dismembered ophiolite (now represented as amphibolites and schists) along the margin of Wadi Fatima Fault hints at the possibility that this structure originated as a suture at this time (Hamimi et al. 2014b). The NW-SE compression generated complex NW-vergent thrusting and tight folding, and was followed by dextral transpression, associated with a more E-W convergence direction. This transpression event is thought to have occurred around 670 Ma but could have continued beyond that date (Abd-Allah et al. 2014). The post-amalgamation Fatima Group molasse was deposited unconformably on the sheared basement rocks in a NE-trending basin, either after or in the late stages of the transpression. The Fatima Group deposition dates from soon after the Nabitah Orogeny (Hamimi et al. 2012), and separates the dextral shearing event in time from the NW-SE sinistral Najd stage. The Fatima

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Fig. 15.19 a Map of the SW side of the Arabian Shield showing the main NE-trending dextral fault zones (from Al Saud 2008). b Map of the Ad Damm fault zone (from Samkari 2015). c Map of the Wadi Fatima fault zone (from Baggazi et al. 2019)

Group was itself later deformed by NW-vergent folds and thrusts during renewed NW-SE compression. Significant shortening (>50%) was involved in the NW-ward compression events (El-Shafei 2017), expressed as oblate strain ellipsoids (Kassem and Hamimi 2018), though Baggazi et al. (2019) reported the bulk strains during transpression as varying from plane strain to constrictional. The idea that both Najd-related NW-oriented sinistral faults and the NE-oriented dextral faults are contemporaneous conjugate pairs was challenged by Stern (1985) on the grounds that: (1) these latter faults are short in length; (2) they are relatively rare; (3) they have nowhere been observed to displace any of the principal Najd strands; and (4) they are confined to the southern parts of the Arabian Shield. Stern (1985) found the NE-oriented faults to be older than the NFS. In general, with the exception of the Ad Damm shear zone, the N-S to NE-SW trending dextral shear zones in the Arabian Shield are distinctly closer in time to the terrane amalgamation (Nabitah and earlier) stages than to the Najd event, and in several cases are separated from the Najd shearing by molasse deposition. Most of these dextral structures appear to have been initiated as terrane sutures, which were later reactivated by dextral transpression as the tectonic convergence direction swung from NW-SE to E-W. The Ad Damm shear zone, with dextral shearing overlapping

in time with the Najd sinistral shearing is, however, a possible candidate for conjugate shear relationship to the Najd sinistral shears (Fleck et al. 1980; Davies 1984; El-Sawy and El-Shafei 2019), though the evidence that these shears existed before Najd argues against them being conjugated in the true sense. Other than N-S and NE-SW dextral faults, Fleck et al. (1976) suggested that E-W dextral faults in the Khaybar Quadrangle (southern part of the Arabian Shield centred at 18°45′N/42°45′E) could also be considered as Najd conjugate shears with dihedral angle of about 40° between the conjugates.

15.12

Discussion of Najd as a Regional System

15.12.1 Along-Strike Trends in NFS Characteristics There exist distinct variations in style of the NFS between the southeastern fault strands and northwestern shear zones. Observations on the regional characteristics of the NFS have been given by Brown (1972), Hadley (1976), Moore and Al Shanti (1979), Moore (1979a, b), Davies (1982, 1984), Stern (1985) and Johnson et al. (2011).

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The southeastern NFS elements (the Halaban-Zarghat, Ar Rika, and Ruwah faults) are represented by very long (100 km) continuous, connected fault segments. Fault wallrocks have low metamorphic grade, and fault zone margins show only narrow zones of foliated rock or no foliations at all. Very slender Jibalah basins lie within the brittle fault zones, and gneissic complexes are also strongly elongated. There are abundant dyke intrusions along the faults. Fault slip included relatively minor vertical displacements. Transpression effects are less obvious, probably due to effective strain partitioning. Volume strains are minor. The initiation of NFS activity in the SE was around 625– 620 Ma. The characteristics of the NFS in the northwest (the Ajjaj, Qazaz, and Hanabiq shear zones) are rather distinct. They are represented by dominantly broad curviplanar ductile shear zones, which show a strong tendency to branch even at large angles of divergence. The shears show amphibolite facies metamorphic grades and enclose lensoidal or more complex-shaped gneissic bodies. Vertical displacements may be extreme. Jibalah basins are rare and there are abundant granitoid intrusions into the shears. Volume strains are significant. Transpression effects are major. Brittle faults are subdued and are curviplanar or represented by groups of short semi-brittle fault strands. The ductile shearing activity appears to have commenced at around 635 Ma. Any model for the origins of the Najd Fault System must explain these along-strike variations. Moore (1979a) and Davies (1984) proposed that the above style differences, from SE to NW, could be explained as shallower SE crustal levels passing into deeper crustal levels in the NW. Al-Damegh et al. (2005) observed that the present-day crustal thickness of the northwestern part of the Arabian shield is 33–37 km, while the southeastern part is 41–53 km thick, which could be explained by greater uplift and erosion of the northwestern parts, exposing deeper crustal levels. This trend in eastwards thickening lithosphere was confirmed by Stern and Johnson (2010). An alternative factor is heat flow. Stern (1985) noted that there had been a higher heat flow in NW, compared to the SE, which could explain the more ductile aspects in the northwestern NFS.

15.12.2 Lithospheric Structure of the NFS Another important question relevant to the entire NFS is the depth to which the strike-slip main faults and shear zones extend. Lemiszki and Brown (1988) compiled deep seismic reflection profiles of continental transforms (forming either conservative boundaries between continental plates or forming as an accommodation to intracontinental rifting) and

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intra-continental strike-slip faults (commonly forming within continental plates as a result of continent-continent collisions) in order to answer this question. They concluded that continental transforms could extend vertically downwards to the base of the crust, or deeper into the lithospheric mantle, while the intra-continental strike-slip faults tended to extend only to the level of subhorizontal midcrust detachments (Fig. 15.20a). Stern’s (2018) work on Egyptian Eastern Desert Najd faults describes Najd flattening out at a midcrustal level corresponding to the extension-related Eastern Desert Shear zone, and not continuing to the base of the crust (Fig. 15.20b). A different approach to the question of the depths penetrated by continental wrench faults was taken by Vauchez and Tommasi (2003), who studied progressively more deeply eroded strike-slip systems in ancient orogens. They found that some intra-continental strike-slip faults developed a listric geometry as they approach midcrust horizontal flow planes at *20 km depth (and T *650 °C), however, even in these examples the presence of mantle-derived magmas along the vertical shears suggested a connection to the upper mantle. In deeper eroded terrains (e.g. 30 km depth for the Mozambique Belt) they found that strike-slip mylonitic shears with subhorizontal stretching lineations persisted at T exceeding 750 °C cutting granulite facies rocks. They discovered that CO2 from fluid inclusions in the shear zones were mantle-derived while CO2 in the wallrocks gave crustal signatures. Mantle flow directions beneath the transpressive strike-slip shears in Taiwan lie parallel to the shears suggesting that the shears are profound enough to influence the directions of tectonic escape. They noted that in these deepest crustal sections transpression is homogeneous (not partitioned) and was associated with thinning by lengthening (lateral escape) of the transpression zone. Cao and Neubauer (2016), though, reported that transpression also resulted in vertical exhumation if buoyancy effects are significant. The abundance of mafic magmas intruding along the NFS is consistent with this system having some connection with lower crust and/or lithospheric mantle, though Idris (2006) concluded that some of the mantle-derived mafic intrusions along the faults are related to the Red Sea spreading in the Tertiary. The P-T studies by Hassan et al. (2016a, b) of gneissic complexes exhumed along the Ajjaj and Hanabiq shear zones suggest that shear took place under conditions reaching 14 kbars (45 km depth), before vertical exhumation along the shears. The original crustal thickness of the Arabian Shield soon after the Nabitah Orogeny and before delamination (at 635 Ma) has been estimated at >50 km by Avigad and Gvirtzman (2009). The geobarometric data of Hassan et al. (2016a, b) point to the Najd shears in the Arabian Shield penetrating at least to the base of the crust.

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Fig. 15.20 a Model of intracontinental transform faults that penetrate the crust and mantle lithosphere (from Cao and Neubauer 2016). b Model of Najd Faulting in the Central Eastern Desert of Egypt (from Stern 2018), involving Najd faults curving towards the Eastern Desert Detachment (EDD) at a depth of about 10–12 km. The EDD separates crustal superstructure from infrastructure

15.12.3 Najd Shear Corridor in Egyptian Eastern Desert Sultan et al. (1988) generated digital images using ratios of Landsat thematic mapper bands (bands 5/4  3/4, 5/1, 5/7) to follow the extension of the NFS from Midyan terrane of Western Arabia into the Eastern Desert of Egypt. They concluded that the Ajjaj Shear Zone marks the intersection of the Najd System with the eastern margin of the Red Sea in the Midyan region. Common features between this zone and the Central Eastern Desert of Egypt include (1) NW— elongation because of folding, with fine‐scale lithologic heterogeneity at the outcrop scale attributed to the deformation that accompanied faulting; (2) NW—trending sinistral strike-slip faults and ductile shear zones; (3) subhorizontal NW-trending stretching lineations, and (4) tectonic contacts. Subsequently, Fritz et al. (1996) defined the Najd Shear Corridor which is a narrow zone between Wadi Kharit-Wadi Hodein shear zone to the south and Duwi shear zone to the north. The most obvious Najd-related shear zones in this corridor are Wadi Kharit-Wadi Hodein shear zone (KHSZ), Nugrus shear zone,

and Atalla shear zone. These shear zones have been the subject matter of detailed studies before, and Hamimi et al. (2019) has summarized the most important structural features of these zones.

15.12.4 Tectonic Models for the NFS One of the most contentious aspects of the Najd Fault System is the tectonic origins of this system. Reviews of the progress in discovering the tectonic mechanisms for the NFS have been provided by Abdelsalam and Stern (1996) and Johnson et al. (2011). The first steps in understanding the NFS came with recognition of the Najd “Orogeny” (Brown 1972). The Najd Orogeny was regarded as a shallower event, compared to the earlier Hijaz (Nabitah) Orogeny. Moore (1979a) observed that the similarity of secondary Najd faults to experiments that generated Riedel shears (Tanner 1962; Wilcox et al. 1973). He proposed that the NFS could have formed in a thin cover overlying active basement wrench faults (Fig. 15.21a). A clearer tectonic cause for the NFS was argued by Schmidt

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Fig. 15.21 Tectonic scale models for the Najd Fault System. a Two structural models for the Najd (from Davies 1984). The diagram on the left explains upper crustal deformation of Najd as occurring in a cover separated by a decollement from moving basement blocks below. The diagram on the right shows the brittle higher levels of Najd faults passing downwards into a broader ductile shear domain below. b The rigid indenter model for the origin of the Najd Fault System (from Davies 1984). c The terrane collision model for the origin of the Najd fault System according to Stoeser and Camp (1985), involving terrane collisions from the west and the east

et al. (1979) and Davies (1984), who considered a crustal collision between a rigid indenter (with a flat interface) and the rigid-plastic Afif and arc terranes to its west, had produced crustal thickening, accommodated by N-S structures (thrusts and folds) (Fig. 15.21b). When this thickened crust rigidified and could no longer fail on N-S structures, the NFS took up further shortening. The model was inspired by studies of the Indian–Asian collision by Molnar and Tapponnier (1977). This model has remained popular and was also the basis of interpretations of the NE-trending dextral strike-slip faults as conjugates of the sinistral NFS (Davies 1984). In the 1980s, questions were raised about the indentation tectonic model. Agar (1987) concluded that the entire NFS had initiated as dextral faults, before reusing as sinistral shears. He found that the curved ends of Najd faults were more appropriate for propagating dextral faults. Nevertheless, he acknowledged that reversals of slip-sense were found in other wrench systems of progressive continental collisions. Stoeser and Camp’s (1985) model for terrane accretion depicted the NFS as commencing during the Afif terrane collision, and evolving during subsequent collisions with the Midyan and Ar Rayn terrane collisions (Fig. 15.21c). This implied collision from both the NW and

the E. Stern (1985) questioned the indentation model on the basis of the lack of evidence for a continent to the east, or for significant crustal thickening relating to a collision. He pointed out that there were time differences between tectonic collision and Najd faulting, and that the NE-trending faults were older than Najd and were therefore unlikely to be dextral Najd conjugates. Most subsequent studies have found the NE-trending dextral faults to be related to suture zones, later reactivated as dextral shears, though the possibility of these shears being Najd conjugates is still accepted by some (El-Sawy and El-Shafei 2019). Studies of the Qazaz–Ajjaj–Hanabiq shear zones have revealed many details of the style of Najd. The dextral Hanabiq was formerly thought to be a conjugate, though this has been challenged by Duncan et al. (1990) who see it as a northern extension of the Hamisana zone. All three shears were believed by AlKashghari (2017), Baggazi et al. (2019) and El-Fakharani et al. (2019) to have formed in the same event, but not as conjugates. Other structural studies of the NE-trending shears have challenged the notion that these are conjugates. Stern (1985) preferred a continental transform-origin for the NFS in accord with extension tectonics (Fig. 15.22a).

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Fig. 15.22 Two further tectonic models for the origin of the Najd Fault System. a the continental transform model (from Stern 1985), associating Najd movement with extension in the North Eastern Desert of Egypt. b Eastward and westward vergent subduction model for the origin of the NFS (from Robinson et al. 2017)

His evidence for the continental transform model included (1) parallelism between the extension direction in Egyptian Northern Eastern Desert and Sinai, and the NFS strands, (2) time overlap of the main rifting movements (600– 575 Ma) and NFS activity (620–560 Ma), and (3) absence of NFS in northernmost Afro-Arabia. Genna et al. (2002) has also referred to the NFS as transforms that governed crustal thinning in the shield. The NFS was considered by Abdelsalam (1994) to be the final stage of a continuous E-W shortening history, with NW-trending sinistral faults and conjugates forming after collision with the Nile craton and development of prominent N-S shortening zones. Fritz et al. (2013) preferred an origin for the Najd that involved a progressive transfer of westwards orogenic transport to sinistral motions on NW-trending parts of the Nabitah suture as a prelude to the Najd Fault System. Burke and Sengör (1986) viewed the Najd as elements accommodating NW-ward tectonic escape from an indentation zone of crustal thickening in the Tanzanian part of the East African Orogen. Tectonic escape was also proposed by Stern (1994) in relation to the NFS, though the escape direction was northwards. Robinson et al. (2017) ascribe the origins of the NFS to opposing subduction polarity between NE part of the shield (eastward subduction of Hijaz beneath Hail & Afif at 620–600 Ma) and the S part of the shield (westwards subduction of Tathlith beneath Asir

at 636–594 Ma) on either side of the Najd Fault Zone (Fig. 15.22b). This configuration explains transpression followed by extension after subduction ceased. Geophysical evidence bearing on the origin of Najd has come from deep crustal studies of the Najd system (described previously in Sect. 12.2) and surveys of the platform covered areas of the ANS to the north and east of the exposed Arabian shield. Johnson and Stewart (1995) reported flat magnetic signatures of a subsurface crustal block, east of Ar Rayn terrane, that could have collided with the western arc terranes. Al-Husseini’s (2000) interpretation of the Arabian basement extended the Amar arc far northwards and defined it as the boundary between the large “Rayn microplate” which collided with the rest of the arc terranes at 640–620 Ma, generating the Najd system. The microplate may have been bounded to the south by the Ruwah fault zone (Johnson and Kattan 2001). Northwards tectonic escape was concluded by Levin and Park (2000) to be the best explanation for magnetic fabrics in the upper mantle north of the ANS. Stern and Johnson (2010) could distinguish western and eastern crust of the Arabian Peninsula, separated by a major N-S trending magnetic anomaly. The two halves differed in crustal and mantle lithosphere thickness. The eastern crust could constitute a separate Neoproterozoic continental plate that collided with the western part of the shield in Neoproterozoic times.

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Johnson and Woldehaimanot (2003) addressed the fundamental question: did Najd form under transpression or extension and rifting (during tectonic escape)? They concluded that transpression could better explain the folds associated with Najd. Numerous other authors have preferred the E-W convergence related transpression mechanism for Najd (Quick 1991). New evidence in the form of metamorphic barometric studies (Al-Saleh et al. 1998; Abu-Alam et al. 2014; Hassan et al. 2016a, b), microstructural and strain studies (Baggazi et al. 2019; El-Fakharani et al. 2019), and numerical rheological studies (Meyer et al. 2014; Meyer 2015) are consistent with a transpression origin of the Najd.

15.12.5 Brittle Reactivation of the NFS and Seismicity Paleozoic and Mesozoic reactivation of the NFS has been recognized in the younger cover rocks above southeastern extensions of the fault system in Yemen (Andre and Blodget 1984; Andre 1989). Based on high-resolution radiometry, thermal IR, and stereoscopic images, these authors found the Najd to extend a further 200 km to the SE of the exposed Arabian Shield. A sinistral displacement of about 2.5 km and some vertical displacements and en échelon folding of cover rocks were deduced to have occurred during mainly Cretaceous times on the Najd system. The above authors supposed that compressive stresses were responsible for the Najd reactivation, perhaps relating to the opening of the Red Sea. Red Sea spreading has also been suspected to have caused rejuvenation of some NE-trending dextral Najd faults, particularly the Ad Damm Fault (AlWash and Zakir 1992) (Fig. 15.23a). El-Isa and Al Shanti (1989) observed that the main recent seismic activity in the Red Sea margins was at the intersections of the spreading zone with NE trending transforms. These transforms could be traced into the continental inland where they were found to include known Najd NE-trending faults, including the Ad Damm and the Wadi Fatima faults (Al-Saud 2008b; Fnais et al. 2015) (Fig. 15.23a). Despite the great length and probable depth of the NW-trending Najd sinistral faults, these structures have not shown much recent seismic activity, and have had little influence on the NNW trend of the Red Sea (Stern and Johnson 2019). This may be due to flattening out of the Najd with depth, or stitching of the deeper parts of the system by multiple intrusions (Stern 2018; Stern and Johnson 2019). Some reactivation of the NW-trending faults as steep normal faults has occurred, forming grabens, e.g., in the Tabuk area of far northern Saudi Arabia, that remains seismically active (Al-Arifi et al. 2013; Zahran et al. 2016; Rehman et al. 2019; Youssof et al. 2020). The reuse of the NW-trending faults as

385

normal faults is consistent with the regional stress field associated with the Red Sea opening. Tertiary to Recent basaltic volcanics have also erupted along NW-trending Najd faults, e.g. at Harret Lunayyir, north of Jeddah. Al-Arifi et al. (2013) found that the focal solutions for earthquakes in the latter area in 2009 pointed to an E-W south-dipping normal fault.

15.12.6 Najd-Related Mineralization Hydrothermal fluids can effectively penetrate and migrate along deep fractures systems, such as the NFS. Moore and Al Shanti (1973) and Moore (1979b) gave a detailed account of the relations of the Zn, Pb, Ag sulphide-rich quartz veins of the Ad Dawadimi area to the regional Najd faults. They found that the most favorable sites for the base metal vein deposits were in the extensional parts of the dense arrays of secondary faults that formed at the terminations of main sinistral Najd faults. Fracture intersections and pinnate joints were intensely veined (Fig. 15.13c). On the question of timing, they concluded that the veins were essentially post-kinematic, while the numerous dyke intrusions into the extensional fractures were syn-kinematic. The fluids were thought to have been related to syn-kinematic granite intrusions or to metamorphic fluids, though they observed very few metamorphic effects at the level of the veins systems. Hydrothermal Au–Cu–As deposits, commonly referred to as “orogenic gold” have also been closely associated with Najd shear systems in the Arabian and Nubian Shields (Abu-Alam et al. 2018). The well-studied Bulghah deposit (Madani 2011) in the northern part of the Arabian Shield consists of gold-pyrite-arsenopyrite ores that formed during the tectonic extension stage, late in the history of the Najd. It was recognized that the areas bordering major Najd fault zones are generally enriched in gold (Madani 2011). Abu-Alam et al. (2018) presented evidence for the metamorphic origin of the gold transporting fluids that were generated at greenschistamphibolite transitional facies temperatures. The fluids were aqueous-carbonic and were cooled during rapid rise along the Najd faults, resulting in precipitation of the gold at suitable dilational sites. In the Nubian Shield Abd El-Wahed et al. (2016) and Zoheir et al. (2019) described the complex array of main and conjugate Najd faults and associated thrusts of transpressional origin that structurally controlled the gold mineralization in the Sukkari gold deposit. Zoheir et al. (2019) also clarified the lithological controls on gold deposition, and emphasized that gold was commonly deposited at the sheared boundaries of Fe-rich rocks (ultramafic and mafic) with other rocks. They noted that the orogenic gold mineralizing event occurred approximately between the regional transpression and transtension stages.

386

Z. Hamimi and A.-R. Fowler

Fig. 15.23 a Distribution of earthquake epicentres (red dots) for Mw < 3.0 magnitude quakes in the Jeddah region (from Fnais et al. 2015), demonstrating the seismic activity along the Ad Damm fault zone. The parallelism of Red Sea transforms and the NE-trending Najd fault elements are evident. b Model for the relationships between the water table and elements of the Najd Fault System in Egyptian Nubian Shield (from Sultan et al. 2008)

15.12.7 Najd Role in Hydrogeological Systems Regionally extensive fault systems such as the NFS have strong effects on river and wadi drainage patterns that, in turn, influence the hydrogeology of the near-surface aquifers. In arid and semi-arid climates, such as those of Saudi Arabia and Egypt, the water table may be deep enough to lie within the fractured bedrocks beneath the superficial sedimentary cover. These fractured aquifers are even more

strongly influenced by the fault, joint, and dyke patterns in the bedrock. The Neoproterozoic Najd faults have been locally reactivated, as described in the previous section, leading to their reuse as normal faults. The resulting narrow grabens produced are typically filled with coarse-grained permeable sediments (conglomerates, sandstones) that are flanked by breccia zones along the reactivated faults. Such features have profound effects on the local hydrogeology, including recharge rates. A good example of this is the Wadi

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Fatima fault zone studied by Alyamani (2006). These authors found that the NE-trending Fatima faults resulted in basins with distinctive hydrogeological properties, due to focusing on groundwater flow and blocking of flow paths in conjunction with E-W dykes. The compartmentalization of these basins is reflected in dramatic differences in salinity of groundwater from one basin to the next. The hydrogeology of the NE-trending Wadi Fatima zone was also studied by Al-Garni (2009), with similar conclusions. These authors used magnetic and DC resistivity methods to outline subsurface structures affecting groundwater flow. Al-Garni (2007) surveyed a nearby major NW-trending wadi (Wadi Rahjan), and found that the groundwater flow was also affected by E-W dyke barriers. These dykes follow one of the common Najd tension fracture trends. Madani et al. (2019) examined the NE-trending Wadi Yalamlam on the Saudi Red Sea coast and reported three generations of WNW- to ENE-trending dykes that acted as groundwater flow barriers. The oldest (Neoproterozoic) dyke set was foliated and was intruded along WNW-trending steep shear zones. Najd Faults in the Eastern Desert of Egypt have exerted similar controls on the fractured aquifers (Sultan et al. 2008) (Fig. 15.23b). These authors combined field, remote sensing, and geophysical data with a digital elevation model to produce a GIS system valuable for groundwater exploration in this arid region. They found a spatial association of wells with Najd Faults, especially where these faults intersected other faults. Dykes also played a role as groundwater flow barriers. Acknowledgements We would like to thank Ghaleb Jarrar and Peter Johnson for reviewing the chapter and their constructive comments that significantly improve the manuscript. Thanks are also due to Tamer Abu-Alam, Ousama Hamid, Mahmoud Hassan, Wael Hagag, Mohamed Abu Soliman for sending some articles. We are grateful to Peter Johnson for Fig. 15.2 image.

References Abd El-Naby H, Frisch W, Siebel W (2008) Tectonic–metamorphic evolution of the Wadi Hafafit culmination (central Eastern Desert, Egypt). Implication for Neoproterozoic core complex exhumation in NE Africa. Geol Acta 6:293–312 Abd El-Wahed MA (2010) The role of the Najd fault system in the tectonic evolution of the Hammamat molasse sediments, Eastern Desert, Egypt. Arab J Geosci 3:1–26 Abd El-Wahed MA, Harraz HZ, El-Behairy MH (2016) Transpressional imbricate thrust zones controlling gold mineralization in the Central Eastern Desert of Egypt. Ore Geol Rev 78:424–446 Abd-Allah AMA, Ahmed AH, El-Fakharani A, El-Sawy EK, Ali KA (2014) Fatima suture: a new amalgamation zone in the western Arabian Shield, Saudi Arabia. Precamb Res 249:57–78 Abdeen MM, Greiling RO (2005) A quantitative structural study of late Pan-African compressional deformation in the Central Eastern Desert (Egypt) during Gondwana assembly. Gondw Res 8:457–471

387 Abdeen MM, Greiling RO, Sadek MF, Hamad SS (2014) Magnetic fabrics and Pan-African structural evolution in the Najd Fault corridor in the Eastern Desert of Egypt. J Afr Earth Sci 99:93–108 Abdeen MM, Sadek MF, Greiling RO (2008) Thrusting and multiple folding in the Neoproterozoic Pan-African basement of Wadi Hodein area, south Eastern Desert, Egypt. J Afr Earth Sci 52:21–29 Abdelazeem M, El-Sawy KE, Gobashy MM (2013) Analysis of magnetic gradients at North East of Wadi Ar Rika quadrangle, Saudi Arabia, to delineate subsurface linear features and faults. NRIAG J. Astron Geophys 2:27–38 Abdelsalam MG (1994) The Oko shear zone: post-accretionary deformations in the Arabian Nubian Shield. J Geol Soc Lond 151:767–776 Abdelsalam MG, Stern RJ (1996) Sutures and shear zones in the Arabian-Nubian Shield. J Afr Earth Sci 23:289–310 Abu-Alam TS, Abdel El Monsef M, Grosch E (2018) Shear-zone hosted gold mineralization of the Arabian-Nubian Shield: Devolatilization processes across the greenschist-amphibolite facies transition. In: Ferrero S, Lanari P, Goncalves P, Grosch EG (Eds.) Metamorphic Geology: Microscale to Mountain Belts. Geol Soc Lond, Special Publication 478:287–313 Abu-Alam TS, Hassan M, Stüwe K, Meyer S, Passchier C (2014) Multistage tectonism during Gondwana collision: Baladiyah Complex, Saudi Arabia. J Petrol 55:1941–1964 Abu-Alam TS, Stüwe K (2009) Exhumation during oblique transpression: the Feiran-Solaf region, Egypt. J Metam Geol 27:439–459 Abu-Alam TS, Stüwe K, Hauzenberger C (2010) Calc-silicates from Wadi Solaf region, Sinai, Egypt. J Afr Earth Sci 58:475–488 Abuzeid HT (1984) Geology of the Wadi Hamrawin Area, Red Sea Hills, Eastern Desert, Egypt, Ph.D. Thesis, Univ. of S.C., Columbia Agar RA (1984) Geologic map of the Zalm quadrangle, sheet 22F, Kingdom of Saudi Arabia. Saudi Arabian Deputy Ministry for Mineral Resources, Open-File Report DGMR-OF-04-01:120 p Agar RA (1986) The Bani Ghayy Group; sedimentation and volcanism in pull-apart grabens of the Najd strike-slip orogen, Saudi Arabian Shield. Precamb Res 31:259–274 Agar RA (1987) The Najd fault system revisited: a two-way strike-slip orogen in the Saudi Arabian Shield. J Struct Geol 9:41–48 Al-Arifi NS, Fat-Helbary RE, Khalil AR, Lashin AA (2013) A new evaluation of seismic hazard for the northwestern part of Saudi Arabia. Nat Hazards 69:1435–1457 Al-Damegh K, Sandvol E, Barazangi M (2005) Crustal structure of the Arabian plate: New constraints from the analysis of teleseismic receiver functions. Earth Planet Sci Lett 231:177–196 Al-Garni MA (2007) Investigating the Groundwater Occurrence in Wadi Rahjan and Its Potential Contribution to Ain Zubaida Using Magnetic and Electric Methods. Makkah Al-Mukarramah, JKAU, Earth Sci. 18:23–47 Al-Garni MA (2009) Geophysical Investigations for Groundwater in a Complex Subsurface Terrain, Wadi Fatima, KSA: A Case History. Jordan J Civ Eng 3(2):118–136 Al-Husseini M (2000) Origin of the Arabian Plate Structures: Amar collision and Najd rift. GeoArabia 5:527–542 Ali-Bik MW, Sadek MF, Ghabrial DS (2014) Late Neoproterozoic metamorphic assemblages along the Pan-African Hamisana Shear Zone, Southeastern Egypt: Metamorphism, geochemistry and petrogenesis. J Afr Earth Sci 99:24–38 AlKashghari WA (2017) Tectonic styles of Hanabiq, Ajjaj, and Qazaz shear zones: implication of oblique transpression in the Arabian Shield, Saudi Arabia, Ph.D. Thesis. King Abdelaziz University Al-Saleh AM (2012) The Kirsh gneiss dome: an extensional metamorphic core complex from the SE Arabian Shield. J Geosci Arab. https://doi.org/10.1007/s12517-0101-0179-1 Al-Saleh AM, Boyle AP, Mussett AE (1998) Metamorphism and 40 Ar/39Ar dating of the Halaban ophiolite and associated units:

388 evidence for two-stage orogenesis in the eastern Arabian shield. J Geol Soc Lond 155:165–175 Al-Saleh AM, Kassem OMK (2012) Microstructural finite strain analysis and 40Ar/39Ar evidence for the origin of the Mizil gneiss dome, eastern Arabian Shield, Saudi Arabia. J Afr Earth Sci 70:24– 35 Al-Saud MM (2008a) Structural mapping from high resolution aeromagnetic data in west central Arabian Shield, Saudi Arabia using normalized derivatives. Arab J Geosci 1:129–136 Al-Saud MM (2008b) Seismic characteristics and kinematic models of Makkah and central Red Sea regions. Arab J Geosci 1:49–61 AlWash MA, Zakir FAR (1992) Tectonic analysis of the Jeddah Taif area on the basis of LANDSAT satellite data. J Afr Earth Sci 15:293–301 Alyamani MS (2006) Hydrogeological and hydrochemical indications of faults control on groundwater flow and quality in Wadi Fatimah Basin, western part of Saudi Arabia. JKAU Earth Sci. 17:1–25 Andre CG (1989) Evidence for Phanerozoic reactivation of the Najd fault system in AVHRR, TM, and SPOT images of Central Arabia. Photogram Eng Rem Sens 55:1129–1136 Andre CG, Blodget HW (1984) Thermal infrared satellite data for the study of tectonic features. Geophys Res Lett 11:983–986 Avigad D, Gvirtzman Z (2009) Late Neoproterozoic rise and fall of the northern Arabian-Nubian shield: the role of lithospheric mantle delamination and subsequent thermal subsidence. Tectonophys 477:217–228 Baggazi HM, AlKashghari WA, El Fakharani A, Matsah M, El-Shafei MK (2019) Northern Arabian Shield shear zones: strain analysis comparison. Arab J Geosci 12:75. https://doi.org/10.1007/s12517019-4233-3 Bamousa AO (2013) Infracambrian superimposed tectonics in the Late Proterozoic units of Mount Ablah area, southern Asir Terrane, Arabian Shield, Saudi Arabia. Arab J Geosci 6:2035–2044 Berhe SM (1990) Ophiolites in NE and East Africa: implications for Proterozoic crustal growth. J Geol Soc Lond 147:41–57 Bishta AZ, Sonbul AR, Qudsi IZ (2015) Utilizing the image processing techniques in mapping the geology of Al Taif area, central western Arabian Shield, Saudi Arabia. Arab J Geosci 8:4161–4175 Blasband B, White S, Brooijmans P, De Boorder H, Visser W (2000) Late Proterozoic extensional collapse in the Arabian-Nubian Shield. J Geol Soc Lond 157:615–628 Bonovia FF, Chorowicz J (1992) Northward expulsion of the Pan-African of NE Africa guided by a reentrant zone of the Tanzania craton. Geology 20:1023–1026 Bregar M, Bauernhofer A, Pelz K, Kloetzli U, Fritz H, Neumayr P (2002) A late Neoproterozoic magmatic core complex in the Eastern Desert of Egypt: emplacement of granitoids in a wrench-tectonic setting. Precamb Res 118:59–82 Brooijmans P, Blasband B, White SH, Visser WJ, Dirks P (2003) Geothermobarometric evidence for a metamorphic core complex in Sinai, Egypt. Precamb Res 123:249–268 Brown GF (1972) Tectonic Map of the Arabian Peninsula. Saudi Arabian Directorate General of Mineral Resources Map AP-2, Scale 1:4,000,000 Brown GF, Coleman RG (1972) The tectonic framework of the Arabian Peninsula. 24th Int Geol Congr Montreal 3:300–305 Brown GF, Jackson RO (1960) The Arabian Shield, 21st Int. Geol Congr Norden 9:69–77 Brown GF, Schmidt DL, Huffman AC Jr (1989) Geology of the Arabian Peninsula: Shield Area of Western Saudi Arabia. US Geological Survey Professional Paper 560-A, 188 p Burke K, Sengör C (1986) Tectonic escape in the evolution of the continental crust. In: Barazangi M, Brown L (eds), Reflection Seismology, vol. 14. American Geophysical Union Geodynamics Series, pp 41–53

Z. Hamimi and A.-R. Fowler Calvez JY, Alsac C, Delfour J, Kemp J, Pellaton C (1984) Geological evolution of western, central, and eastern parts of the northern Precambrian Shield, Kingdom of Saudi Arabia. King Abdulaziz University, Jiddah, Faculty of Earth Science Bulletin 6:24–48 Calvez JY, Alsac C, Delfour J, Kemp J, Pellaton C (1983) Geologic Evolution of Western, Central, and Eastern Parts of the Northern Precambrian Shield, Kingdom of Saudi Arabia. Saudi Arabian Deputy Ministry for Mineral Resources Open-File Report BRGM-OF-03-17, 57 p Cao S, Neubauer F (2016) Deep crustal expressions of exhumed strike-slip fault systems: Shear zone initiation on rheological boundaries. Earth Sci Rev 162:155–176 Cole JC, Hedge CE (1986) Geochronologic Investigation of Late Proterozoic Rocks in the Northeastern Shield of Saudi Arabia. Saudi Arabian Deputy Ministry for Mineral Resources Technical Record USGS-TR-05-5, 42 p Collins AS, Pisarevsky SA (2005) Amalgamating eastern Gondwana: the evolution of the Circum-Indian Orogens. Earth Sci Rev 71:229–270 Davies FB (1980) Reconnaissance geology of the Duba Quadrangle, sheet 27/35D, Kingdom of Saudi Arabia, scale 1:100,000, Geol. Map GM Saudi Arabia Dir. Gen. Miner. Resour., GM-57, 23 p Davies FB (1982) Quantitative strain analysis of Pan-African granites intruded into a ductile shear zone in the northwestern Al Hijaz, Kingdom of Saudi Arabia, Saudi Arabia Dep. Minist Miner Resour Prof Pap 1:53–65 Davies FB (1984) Strain analysis of wrench faults and collision tectonics of the Arabian-Nubian Shield. J Geol 82:37–53 Delfour J (1970) Le groupe de J’balah une nouvelle unite du bouchlier Arabe. Bur Recher Géol Min Bull 2(4):19–32 Delfour J (1977) Geology of the Nuqrah quadrangle, Sheet 25E, Kingdom of Saudi Arabia: Directorate General of Mineral Resources Geological Map GM 28, Scale 1:250,000, 32 p Delfour J (1979) Pan-African orogeny in the northern part of the Arabian Shield. Bull. Soc. Geol. France S7-XXI(4):449–456 Delfour J (1982) Geologic Map of the Halaban quadrangle, Sheet 23G, Kingdom of Saudi Arabia. Directorate General of Mineral Resources Geologic Map GM 46, Scale 1:250,000, 32 p Divi RS, Zakir FA (2001) Delineation of Tectonic Features Utilizing Satellite Remote Sensing Data: I—The Southern-Half of the Arabian Shield. Gondw Res 4(2):159–161 Duncan IJ, Rivard B, Arvidson RE, Sultan M (1990) Structural interpretation and tectonic evolution of a part of the Najd Shear Zone (Saudi Arabia) using Landsat thematic-mapper data. Tectonophys 178:309–315 El Fakharani A, AlKashghari WA, Baggazi HM, El Shafei MK, Matsah MI (2019) Microstructural analyses of the Najd Fault System in Midyan Terrane, NW Arabian Shield, Saudi Arabia. International Journal of Earth Sciences. https://doi.org/10.1007/ s00531-019-01803-w El-Isa ZH, Al Shanti A (1989) Seismicity and tectonics of the Red Sea and western Arabia. Geophys J 97:449–457 El-Rabaa SM, Al-Shumaimri MS, Al-Mishwat AT (2001) Fate of the Najd fault system in northwestern Saudi Arabia and southwestern Jordan. Gondw Res 4:164–165 El-Ramly MF, Greiling R, Kröner A, Rashwan AA (1984) On the tectonic evolution of the Wadi Hafafit area and environs, Eastern Desert of Egypt. King Abdulaziz University, Jiddah. Bull Fac Earth Sci 6:114–126 El-Sawy EK, El-Shafei M (2019) Geometric and tectonic analysis of Ad-Damm mega-scale fold: implication of Neoproterozoic Transpressive Regime in the west-central Arabian Shield. Arab J Geosci 12(7):224 El-Sawy EK, Masrouhi A (2019) Structural style and kinematic evolution of Al Ji’lani area, Ad Dawadimi terrane, Saudi Arabia. J Afr Earth Sci 150:451–465

15

Najd Shear System in the Arabian-Nubian Shield

El-Shafei MK (2017) Thrust duplex deformation in the volcaniclastic sequence of the Fatima fold-and-thrust belt in the west-central Arabian Shield. J Asian Earth Sci 138:211–229 Emam A, Hamimi Z, El-Fakharani A, Abdel-Rahman E, Gomez-Barreiro J, Abo-Soliman MY (2018) Utilization of ASTER and OLI data for lithological mapping of Nugrus-Hafafit area, South Eastern Desert of Egypt. Arab J Geosci 11:756 Fleck RJ, Coleman RJ, Cornwall HF, Greenwood WR, Hadley DG, Schmidt DL, Prinz WC, Ratte JC (1976) Geochronology of the Arabian Shield, western Saudi Arabia: K-Ar results. Geol Soc Am Bull 87:9–21 Fleck RJ, Greenwood WR, Hadley DG, Anderson RE, Schmidt DL (1980) Rubidium-Strontium Geochronology and Plate-tectonic Evolution of the Southern Part of the Arabian Shield. USGS Professional Paper 1131, 38 p Fleck RJ, Hadley DG (1982) Ages and Strontium Initial Ratios of Plutonic Rocks in a Transect of the Arabian Shield. Saudi Arabian Deputy Ministry for Mineral Resources Open-File Report USGS-OF-03-38, 43 p Flowerdew MJ, Whitehouse MJ, Stoeser DB (2013) The Nabitah fault zone, Saudi Arabia: A Pan-African suture separating juvenile oceanic arcs. Precamb Res 239:95–105 Fnais M, Al-Amri A, Abdelrahman K, Abdelmonem E, El-Hady S (2015) Seismicity and Seismotectonics of Jeddah-Makkah Region, West-Central Saudi Arabia. J Earth Sci 26:746–754 Fossen H (2010) Structural Geology. Cambridge University Press, 480p Fowler A, El-Kalioubi B (2004) Gravitational collapse origin of shear zones, foliations and linear structures in the Neoproterozoic cover nappes, Eastern Desert, Egypt. J Afr Earth Sci 38:23–40 Fowler A, Hamimi Z (2020) Structural and Tectonic Framework of Neoproterozoic Basement of Egypt: From Gneiss Domes to Transpression Belts. In: Hamimi Z, El-Barkooky A, Martínez Frías J, Fritz H, Abd El-Rahman Y (eds) The Geology of Egypt, Regional Geology Reviews. Springer, Switzerland, pp 81–129 Fowler A, Hassen I, Hassan M (2015) Tectonic evolution and setting of the Sa’al Complex, southern Sinai, Egypt: A Proterozoic continental back-arc rift model. J Afr Earth Sci 104:103–131 Fowler A, Khamees H, Dowidar H (2007) El Sibai gneissic complex, Central Eastern Desert, Egypt: folded nappes and syn-kinematic gneissic granitoid sheets—not a core complex. J Afr Earth Sci 49:119–135 Fowler A, Osman AF (2009) The Sha’it–Nugrus shear zone separating Central and South Eastern Deserts, Egypt: a post-arc collision low-angle normal ductile shear zone. J Afr Earth Sci 53:16–32 Fritz H, Abdelsalam M, Ali KA, Bingen B, Collins AS, Fowler AR, Ghebreab W, Hauzenberger CA, Johnson PR, Kusky TM, Mace P, Muhongo S, Stern RG, Viola G (2013) Orogen styles in the East African Orogen: a review of the Neoproterozoic to Cambrian tectonic evolution. J Afr Earth Sci 86:65–106 Fritz H, Dallmeyer DR, Wallbrecher E, Loizenbauer J, Hoinkes G, Neymayr P, Khudeir AA (2002) Neoproterozoic tectonothermal evolution of the Central Eastern Desert, Egypt: a slow velocity tectonic process of core complex exhumation. J Afr Earth Sci 34:137–155 Fritz H, Messner M (1999) Intramontane basin formation during oblique convergence in the Eastern Desert of Egypt: magmatically versus tectonically induced subsidence. Tectonophys 315:145–162 Fritz H, Wallbrecher E, Khudeir AA, El Ela A, Dallmeyer DR (1996) Formation of Neoproterozoic metamorphic core complexes during oblique convergence (Eastern Desert, Egypt). J Afr Earth Sci 23:311–329 Genna A, Nehlig P, Le Goff E, Guerrot C, Shanti M (2002) Proterozoic tectonism of the Arabian Shield. Precamb Res 117:21–40

389 Gettings ME, Blank HR, Mooney WD (1986) Crustal structure of southwestern Saudi Arabia. J Geophys Res 91:6491–6512 Ghebreab W (1998) Tectonics of the Red Sea region reassessed. Earth-Sci Rev 45:1–44 Greenwood WR, Hadley DG, Anderson RE, Fleck RJ, Schmidt DL (1976) Late Proterozoic cratonization in southwestern Saudi Arabia. Phil Trans Roy Soc Lond A280:517–527 Greiling RO, Abdeen MM, Dardir AA, El-Akhal H, El-Ramly MF, Kamal El-Din GM, Osman AF, Rashwan AA, Rice AHN, Sadek MF (1994) A structural synthesis of the Proterozoic Arabian-Nubian Shield in Egypt. Geolog Rundsh 83:484–501 Hadley DG (1974) The Taphrogeosynclinal Jubaylah Group in the Mashhad Area, Northwestern Hijaz, Kingdom of Saudi Arabia. Saudi Arabian Directorate General of Mineral Resource Bulletin 10, 18 pp Hadley DG (1976) Geology of the Bi’r Juqjuq quadrangle, Kingdom of Saudi Arabia: Saudi Arabian Directoral General of Mineral Resources. Geologic Map GM-26, scale 1:100,000 Hagag W, Moustafa R, Hamimi Z (2018) Neoproterozoic evolution and Najd-related transpressive shear deformations along Nugrus Shear Zone, South Eastern Desert, Egypt (Implications from field-str uctural data and AMS-technique). Geotectonics 52(1):114–133 Hamimi Z, Abd El-Wahed M, Gahlan HA, Kamh SZ (2019) Tectonics of the Eastern Desert of Egypt: Key to understanding the Neoproterozoic evolution of the Arabian-Nubian Shield (East African Orogen). In: Hamimi Z, Hamoudi M, Djemai S, Zoheir B (eds) Bendaoud A Geology of the Arab World—An Overview. Switzerland, Springer Geology, pp 1–81 Hamimi Z, Abd El-Wahed MA (2020) Suture(s) and Major Shear Zones in the Neoproterozoic Basement of Egypt. In: Hamimi Z, El-Barkooky A, Martínez Frías J, Fritz H, Abd El-Rahman Y (eds) The Geology of Egypt, Regional Geology Reviews. Springer, Switzerland, pp 153–189 Hamimi Z, El-Fakharani A, Abdeen M (2014a) Polyphase deformation History and Strain Analysis of the Post-Amalgamation Depositional Basins in the Arabian-Nubian Shield: Evidence from Fatima, Ablah and Hammamat Basins. J Afr Earth Sci 99:64–92 Hamimi Z, El-Fakharani A, Emam A, Gomez-Barreiro J, Abdelrahman E, Abo-Soliman MY (2018) Reappraisal of the kinematic history of Nugrus shear zone using PALSAR and microstructural data: implications for the tectonic evolution of the Eastern Desert tectonic terrane, northern Nubian Shield. Arab J Geosci 11:494. https://doi.org/10.1007/s12517-018-3837-3 Hamimi Z, El-Sawy KE, El-Fakharani A, Shujoon A, Matsah M, El-Shafei M (2014b) Neoproterozoic Structural Evolution of the NE-trending 620–540 Ma Ad-Damm Shear Zone, Arabian Shield, Saudi Arabia. J Afr Earth Sci 99:51–63 Hamimi Z, El-Shafei M, Kattu G, Matsah M (2013) Transpressional Regime in Southern Arabian Shield: Insights From Wadi Yiba Area, Saudi Arabia. Miner Petrol 107:849–860 Hamimi Z, Hagag W, Kamh S, El-Araby A (2020) Application of remote-sensing techniques in geological and structural mapping of Atalla Shear Zone and Environs, Central Eastern Desert, Egypt. Arab J Geosci 13:414. https://doi.org/10.1007/s12517-020-05324-8 Hamimi Z, Matsah M, El-Shafei M, El-Fakharani A, Shujoon A, Al-Jabali M (2012) Wadi Fatima Thin-Skinned Foreland FAT Belt: A Post Amalgamation Marine Basin in the Arabian Shield. Open J Geol 2:271–293 Hassan M, Abu-Alam TS, Hauzenberger C, Stüwe K (2016a) Geochemical signature variation of pre-, syn-, and post-shearing intrusives within the Najd fault system of western Saudi Arabia. Lithos 263:274–291 Hassan M, Stüwe K, Abu-Alam TS, Klötzli U, Tiepolo M (2016b) Time constraints on deformation of the Ajjaj branch of one of the

390 largest Proterozoic shear zones on Earth: The Najd Fault System. Gondw Res 34:346–362 Hedge CE (1984) Precambrian Geochronology of Part of Northwestern Saudi Arabia. Saudi Arabian Deputy Ministry for Mineral Resources Open-File Report USGS-OF-04-31, 12 p Hoffman PF (1999) The break-up of Rodinia, birth of Gondwana, true polar wander and the snowball Earth. J Afr Earth Sci 28:17–33 Howland AF (1979) ‘Tectonics of the Najd Transcurrent Fault System, Saudi Arabia’ by J.M. Moore (1979): discussion. J Geol Soc Lond 136:441–454 Idris EO (2006). Subsurface Structural Framework of the Tertiary Sedimentary Rocks East of Jeddah, Saudi Arabia. PhD Thesis. Khartoum University, Sudan Jacobs J, Thomas RJ (2004) A Himalayan-type indenter-escape tectonic model for the southern part of the Late Neoproterozoic Early Paleozoic East African-Antarctic Orogen. Geology 32:721– 724 Jacobs J, Fanning CM, Henjes-Kunst F, Olesch M, Paech HJ (1998) Continuation of the Mozambique Belt into East Antarctica: Grenville-age metamorphism and polyphase Pan-African high-grade events in central Dronning Maud Land. J Geol 106:385–406 Jacobs J, Bingen B, Thomas RJ, Bauer W, Wingate MTD, Feitio P (2008) Early Palaeozoic orogenic collapse and voluminous late-tectonic magmatism in Dronning Maud Land and Mozambique: insights into the partially delaminated orogenic rootof the East African-Antarctic Orogen? Geol Soc Lond Spec Publ 308:69–90 Johnson PR (1996) Geochronologic and Isotopic Data for Rocks in the East-Central Part of the Arabian Shield Stratigraphic and Tectonic Implications. Ministry of Petroleum and Mineral Resources, Directorate General of Mineral Resources, Jeddah, Saudi Arabia Johnson PR, Andresen A, Collins AS, Fowler AR, Fritz H, Ghebreab W, Kusky T, Stern RJ (2011) Late Cryogenian-Ediacaran history of the Arabian-Nubian Shield: a review of depositional, plutonic, structural, and tectonic events in the closing stages of the northern East African Orogen. J Afr Earth Sci 61:167–232 Johnson PR, Kattan F (2001) Oblique sinistral transpression in the Arabian shield: the timing and kinematics of a Neoproterozoic suture zone. Precambr Res 107:117–138 Johnson PR, Kattan FH, Wooden JL (2001) Implications of SHRIMP and microstructural data on the age and kinematics of shearing in the Asir Terrane, southern Arabian shield, Saudi Arabia. Gondw Res 4:172–173 Johnson PR, Kattan FK, Al-Saleh AM (2004) Neoproterozoic ophiolites in the Arabian shield: field relations and structure. In: Kusky TM (Ed.), Precambrian Ophiolites and Related rocks, Developments in Precambrian Geology, vol. 13. Elsevier, pp. 129–162 Johnson PR, Stewart ICF (1995) Magnetically inferred basement structure in central Saudi Arabia. Tectonophys 245:37–52 Johnson PR, Woldehaimanot B (2003) Development of the Arabian– Nubian shield: perspectives on accretion and deformation in the northern East African Orogen and the assembly of Gondwana. In: Yoshida M, Windley BF, Dasgupta S (eds), Proterozoic East Gondwana: Supercontinent Assembly and Breakup: Geol. Soc. Lond. Spec. Publ. 206:290–325 Jones RR, Holdsworth RE, Bailey W (1997) Lateral extrusion in transpression zones: the importance of boundary conditions. J Struct Geol 19:1201–1217 Jones RR, Tanner PWG (1995) Strain partitioning in transpression zones. J Struct Geol 17:793–802 Kahal AY, Abd El-Motaal E, Kassem OMK, Al Ghoreiby AHR (2019) Strain analysis and deformation history of Zalm area, Arabian shield, Saudi Arabia. J Afr Earth Sci 150:441–450

Z. Hamimi and A.-R. Fowler Kassem OMK, Hamimi Z (2018) Finite Strain Analysis of the Wadi Fatima Shear Zone in Western Arabia, Saudi Arabia. Geotectonics 52:251–265 Katumwehe AB, Abdelsalam MG, Atekwana EA, Laó-Dávila DA (2016) Extent, kinematics and tectonic origin of the Precambrian Aswa Shear Zone in eastern Africa. Gondw Res 34:241–253 Kellogg KS, Janjou D, Minoux L, Fourniguet J (1986) Geological Map of the Wadi Tathlith Quadrangle, Sheet 20G, Kingdom of Saudi Arabia. Saudi Arabian Deputy Ministry for Mineral Resources Geologic Map GM 103, Scale 1:250,000, 27 p Kemp J, Pellaton C, Calvez, JY, 1980. Geochronological Investigations and Geological History in the Precambrian of Northwestern Saudi Arabia. Saudi Arabian Directorate General of Mineral Resources Open-File Report BRGM-OF-01-1, 120 p Kennedy A, Johnson PR, Kattan FH (2004) SHRIMP Geochronology in the Northern Arabian Shield. Part I. Data Acquisition. Saudi Geological Survey Open-File Report SGS-OF-2004-11 Kennedy A, Johnson PR, Kattan FH (2005) SHRIMP Geochronology in the Northern Arabian Shield. Part II. Data Acquisition 2004. Saudi Geological Survey Open-File Report SGS-OF-2005-10 Kennedy A, Johnson PR, Kattan FH (2009) SHRIMP geochronology in the northern Arabian Shield: part III data acquisition, 2006. Saudi Geological Survey Open File Report, Jeddah Kennedy A, Kozdroj W, Johnson PR, Kattan FH (2011) SHRIMP Geochronology in the Northern Arabian Shield. Part III. Data Acquisition 2006. Saudi Geological Survey Open-File Report SGS-OF-2007-9 Kennedy A, Kozdroj W, Kattan FH, Kodzroj MZ, Johnson PR (2010) SHRIMP Geochronology in the Arabian Shield (Midyan Terrane, Afif Terrane, Ad Dawadimi Terrane) and Nubian Shield (Central Eastern Desert), Part IV. Data Acquisition 2008. Saudi Geological Survey Open-File Report SGS-OF-2010-10, 101 p Kozdrój W, Kennedy AK, Johnson PR, Ziółkowska-Kozdrój M, Kadi K (2018) Geochronology in the southern Midyan terrane: a review of constraints on the timing of magmatic pulses and tectonic evolution in a northwestern part of the Arabian Shield. Internat Geol Rev 60:1290–1319 Kröner A, Greiling R, Reischmann T, Hussein IM, Stern RJ, Durr S, Kruger J, Zimmer M (1987) Pan-African crustal evolution in the Nubian segment of NE Africa. In: Kröner A (Ed.), Proterozoic Lithospheric Evolution, American Geophysical Union Geodynamics Series, vol. 17, pp 235–257 Kröner A, Todt W, Hussein IM, Mansour M, Rashwan AA (1992) Dating of late Proterozoic ophiolites in Egypt and the Sudan using the single grain zircon evaporation technique. Precamb Res 59:15–32 Kusky T, Matsah M (2003) Neoproterozoic Dextral Faulting on the Najd Fault System, Saudi Arabia, preceded sinistral faulting and escape tectonics related to closure of the Mozambique Ocean. In: Yoshida M, Windley BF, Dasgupta S, Powell C (eds), Proterozoic East Gondwana: Supercontinent Assembly and Breakup. Geol Soc, Lond Spec Publ 206:327–361 Kröner A, Stern RJ (2004) Pan-African orogeny. Encyclopedia of geology (2004), vol 1, Elsevier, Amsterdam Lemiszki PJ, Brown LD (1988) Variable crustal structure of strike-slip fault zones as observed on deep seismic reflection profiles. Geol Soc Am Bull 100:665–676 Levin V, Park J (2000) Shear zones in the Proterozoic lithosphere of the Arabian Shield and the nature of the Hales discontinuity. Tectonophys 323:131–148 Li ZX, Bogdanova SV, Collins AS, Davidson A, De Waele B, Ernst RE, Fitzsimons ICW, Fuck RA, Gladkochub DP, Jacobs J, Karlstrom KE, Lul S, Natapov LM, Pease V, Pisarevsky SA, Thrane K, Vernikovsky V (2008) Assembly, configuration, and break-up history of Rodinia: A synthesis. Precamb Res 160:179–210

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Najd Shear System in the Arabian-Nubian Shield

Madani A, Niyazi B, Elfakharani A, Osman H (2019) The Effects of Structural Elements on Groundwater of Wadi Yalamlam, Saudi Arabia Using Integration of Remote Sensing and Airborne Magnetic Survey. Earth Sys Environ 3:301–312 Madani AA (2011) Knowledge-driven GIS modeling technique for gold exploration, Bulghah gold mine area, Saudi Arabia, Egypt. J Rem Sens Space Sci 14:91–97 Makroum F (2017) Structural interpretation of the Wadi Hafafit culmination: a Pan-African gneissic dome in the central Eastern Desert, Egypt. Lithosphere 9:759–773 Matsah MI, Qari MHT, Hegazi AM, Amlas MA, Hamimi Z (2004) The Neoproterozoic Ad-Damm Shear Zone: Dextral transpression in the Arabian Shield, Saudi Arabia. Egypt J Geol 48:215–236 Matsah MIM, Kusky TM (2001) Analysis of Landsat ratio imagery of the Halaban-Zarghat fault and related Jifn basin, NE Arabian shield: implications for the kinematic history of the Najd fault. Gondw Res 4:182 Meyer SE (2015) Branching shear zones in the Arabian-Nubian Shield. PhD Thesis, University of Mainz Meyer SE, Passchier C, Abu-Alam T, Stüwe K (2014) A strike-slip core complex from the Najd fault system, Arabian shield. Terra Nova 26:387–394 Mogren S, Al-Amri AS, Al-Damegh K, Fairhead D, Jassim S, Algamdi A (2008) Sub-surface geometry of Ar Rika and Ruwah faults from Gravity and Magnetic Surveys. Arab J Geosci 1:33–47 Mogren SMA (2004) Geophysical Investigations on the Najd Fault System. PhD Thesis, University of Newcastle upon Tyne Molnar P, Tapponnier P (1977) Relation of the tectonics of eastern China to the India-Eurasia collision: application of slip-line field theory to large-scale continental tectonics. Geology 5:212–216 Moore JM (1979a) Tectonics of the Najd transcurrent fault system, Saudi Arabia. J Geol Soc Lond 136:441–454 Moore JM (1979b) Primary and secondary faulting in the Najd fault system, Kingdom of Saudi Arabia. US Geological Survey Open File Report 79-1661 Moore JM, Al Shanti AM (1973) The use of stress trajectory analysis in the elucidation of part of the Najd Fault System, Saudi Arabia. Proc Geol Assoc Lond 4:383–403 Moore JM, Al Shanti AM (1979) Structure and mineralization in the Najd Fault System, Saudi Arabia. In: Tahoun S (ed) Evolution and Mineralization of the Arabian-Nubian Shield, vol 2. Pergamon, New York, pp 17–28 Nance RD, Murphy JB, Santosh M (2014) The supercontinent cycle: a retrospective essay. Gondw Res 25:4–29 Naylor MA, Mandl G, Sijpesteijn CHK (1986) Fault geometries in basement-induced wrench faulting under different initial stress states. J Struct Geol 8:737–752 Nehlig P, Asfirane F, Genna A, Guerrot C, Nicol N, Salpeteur I, Shanti M, Thiéblemont D (2001) Aeromagnetic map constrains cratonization of the Arabian Shield. Terra Nova 13(5):347–353 Nehlig P, Genna A, Asfirane F (2002) A review of the Pan-African evolution of the Arabian Shield. GeoArabia 7:103–124 Nettle D (2009) A Sequence Stratigraphic, Geochronological and Chemostratigraphic Investigation of the Ediacaran Antaq Basin, Eastern Arabian Shield, Saudi Arabia, Geology and Geophysics. University of Adelaide, p. 83 Och LM, Shields-Zhou GA (2012) The Neoproterozoic oxygenation event: environmental perturbations and biogeochemical cycling. Earth-Sci Rev 110:26–57 Patchett PJ, Chase CG (2002) Role of transform continental margins in major crustal growth episodes. Geology 30:39–42 Quick JE (1991) Late Proterozoic transpression on the Nabitah fault system—implications for the assembly of the Arabian Shield. Precamb Res 53:119–147

391 Rehman F, El-Hady SM, Harbi HM, Fahad Ullah M, Rehman S, Riaz O, Hassan S (2019) Potential Seismogenic Source Model for the Red Sea and Coastal areas of Saudi Arabia. Int J Econ Environ Geol 10(2):26–31 Robinson FA, Bonin B, Pease V, Anderso JL (2017) A discussion on the tectonic implications of Ediacaran late- to post-orogenic A-type granite in the northeastern Arabian Shield, Saudi Arabia. Tectonics 36:582–600. https://doi.org/10.1002/2016TC004320 Samkari A (2015) Structural Evolution of Ad Damm Shear Zone, Western Saudi Arabian Margin and Its Relation to Red Sea Rift System. MSc Thesis, Florida State University Sanderson D, Marchini RD (1984) Transpression. J Struct Geol 6:449– 458 Schmidt DL, Brown GF (1984) Major-element evolution of the late Proterozoic shield of Saudi Arabia: Proceedings, First symposium of Pan-African Crustal Evolution in the Arabian-Nubian Shield IGCP no. 164. King Abdulaziz University, Jiddah 6:1–21 Schmidt DL, Hadley DG, Stoeser DB (1979) Late Proterozoic crustal history of the Arabian shield, southern Najd province, Kingdom of Saudi Arabia. King Abdulaziz University, Jiddah. Inst Appl Geolo Bullet 3(2):41–58 Sengör AMC, Natalin BA (1996) Paleotectonics of Asia: fragment of a synthesis. In: Yin A, Harrison TM (eds) The Tectonics of Asia. Cambridge University Press, New York, pp 486–640 Shackleton RM, Ries AC, Graham RH, Fitches WR (1980) Late Precambrian ophiolitic melange in the east Desert of Egypt. Nature 285:472–474 Shalaby A, Stüwe K, Fritz H, Makroum F (2006) The El Mayah molasse basin in the Eastern Desert of Egypt. J Afr Earth Sci 45:1–15 Squire RJ, Campbell IH, Allen CM, Wilson CJL (2006) Did the Transgondwanan Supermountain trigger the explosive radiation of animals on Earth? Earth Planet Sci Lett 250:116–133 Stacey JS, Agar RA (1985) U-Pb isotopic evidence for the accretion of a continental micro plate in the Zalm region of the Saudi Arabian Shield. J Geol Soc Lond. 142:1189–1203 Stacey JS, Agar RA (1986) U-Pb isotopic evidence for accretion of a continental microplate in the Zalm region of the Saudi Arabian Shield. Report for Ministry of Petroleum and Mineral Resources, Deputy Ministry for Mineral Resources, Saudi Arabia. U.S. Geol. Surv. Open-file Report 86–166 Stacey JS, Stoeser DB, Greenwood WR, Fischer LB (1984) U-Pb zircon geochronology and geologic evolution of the Halaban–Al Amar region of the eastern Arabian Shield, Kingdom of Saudi Arabia. J Geol Soc Lond 141:1043–1055 Stern RJ (1985) The Najd fault system, Saudi Arabia and Egypt: a late Precambrian rift-related transform system? Tectonics 4:497–511 Stern RJ (1994) Arc assembly and continental collision in the Neoproterozoic East African Orogen: implications for the consolidation of Gondwanaland. Ann Rev Earth Planet Sci 22:315–319 Stern RJ (2002) Crustal evolution in the East African Orogen: a neodymium isotopic perspective. J Afr Earth Sci 34:109–117 Stern RJ (2018) Neoproterozoic formation and evolution of Eastern Desert continental crust—The importance of the infrastructuresuperstructure transition. J Afr Earth Sci 146:15–27 Stern RJ, Abdelsalam MG (1998) Formation of juvenile continental crust in the Arabian-Nubian shield: evidence from granitic rocks of the Nakasib Suture, NE Sudan. Geolog. Rundsch 87:150–160 Stern RJ, Johnson PR (2010) Continental lithosphere of the Arabian Plate: a geologic, petrologic, and geophysical synthesis. Earth-Sci Rev 101:29–67 Stern RJ, Johnson PR (2019) Constraining the Opening of the Red Sea: Evidence from the Neoproterozoic Margins and Cenozoic Magmatism for a Volcanic Rifted Margin. Rasul NMA, Stewart ICF (Eds.), Geological Setting, Palaeoenvironment and Archaeology of the Red

392 Sea. Springer Nature Switzerland. https://doi.org/10.1007/978-3319-99408-6_4 Stern RJ, Kröner A, Manton WI, Reischmann T, Mansour M, Hussein IM (1989) Geochronology of the late Precambrian Hamisana shear zone, Red Sea Hills, Sudan and Egypt. J Geol Soc Lond 146:1017–1029 Stoeser DB, Camp VE (1985) Pan-African microplate accretion of the Arabian Shield. Geol Soc Am Bull 96:817–826 Stoeser DB, Stacey JS (1988) Evolution, U-Pb geochronology, and isotope geology of the Pan-African Nabitah orogenic belt of the Saudi Arabian Shield. In: El-Gaby S, Greiling RO (eds) The Pan-African belt of NE African and Adjacent Areas. FriedrichViewig and Sohn, Braunschweig/Wiesbaden, pp 227–288 Stoeser DB, Stacey JS, Greenwood WR, Fisher LB (1985) U/PB zircon geochronology of the southern part of the Nabitah mobile belt and Pan-African Continental collision in the Saudi Arabian Shield. Ministry of Petroleum and Mineral Resources, Deputy Ministry for Mineral Resources, Jiddah, Kingdom of Saudi Arabia. U.S. Geol. Surv. Open-File Report 85–239 Stüwe K, Kadi K, Abu-Alam T, Hassan M (2014) The Najd Fault System of Saudi Arabia. EGU General Assembly 2014, held 27 April–2 May, 2014, Vienna, Austria, id.5945 Sultan M, Arvidson RE, Duncan IJ, Stern R, El Kaliouby B (1988) Extension of the Najd Fault System from Saudi Arabia to the central Eastern Desert of Egypt based on integrated field and Landsat observations. Tectonics 7:1291–1306 Sultan M, Wagdy A, Manocha N, Sauck W, Abdel Gelil K, Youssef AF, Becker R, Milewski A, El Alfy Z, Jones C (2008) An integrated approach for identifying aquifers in transcurrent fault systems: The Najd shear system of the Arabian Nubian shield. J Hydrol 349:475–488 Sultan YM, El-Shafei MK, Arnous MO (2017) Tectonic evolution of Kid Metamorphic Complex and the recognition of Najd Fault System in south East Sinai, Egypt. Int J. Earth Sci. (Geol. Rundsch), 106:2817–2836 Tanner WF (1962) Surface structural patterns obtained from strike-slip models. J Geol 70:101–107 Vail JR (1985) Pan-African (late Precambrian) tectonic terrains and the reconstruction of the Arabian-Nubian Shield. Geology 13:839–842

Z. Hamimi and A.-R. Fowler Vauchez A, Tommasi A (2003) Wrench faults down to the asthenosphere: Geological and geophysical evidence and thermomechanical effects. In: Storti F, Holdsworth RE, Salvini F (Eds.), Intraplate Strike-Slip Deformation Belts: London, Geological Society of London, Special Publication 210, 15–34 Wilcox RE, Harding TP, Seeley DR (1973) Basic wrench tectonics. Am Assoc Petrol Geol Bull 57:74–96 Wipfler EL (1996) Transpressive structures in the Neoproterozoic Ariab-Nakasib Belt, northeast Sudan: evidence for suturing by oblique collision. J Afr Earth Sci 23:347–362 Youssof M, Mai PM, Nobile A, Jónsson S (2020) Non-Volcanic Earthquake Swarm Near the Harrat Lunayyir Volcanic Field. King Abdullah University of Science and Technology, Saudi Arabia. https://doi.org/10.1002/essoar.10503155.1 Zahran HM, Sokolov V, Roobol MJ, Stewart ICF, Youssef SE, El-Hadidy M (2016) On the development of a seismic source zonation model for seismic hazard assessment in western Saudi Arabia. J Seismol 20:747–769 Zoheir BA (2011) Transpressional zones in ophiolitic mélange terranes: potential exploration targets for gold in the South Eastern Desert, Egypt. J Geochem Explor 111:23–38 Zoheir BA, Feigenson M, Zi J-W, Turrin B, Deshesh F, El-Metwally A (2018) Ediacaran (* 600 Ma) orogenic gold in Egypt: age of the Atalla gold mineralization and its geological significance. Int Geol Rev 1–16 Zoheir BA, Johnson PR, Goldfarb RJ, Klemm DD (2019) Orogenic gold in the Egyptian Eastern Desert: Widespread gold mineralization in the late stages of Neoproterozoic orogeny. Gondw Res 75:184–217 Zoheir BA, Moritz R (2014) Fluid evolution in the El-Sid gold deposit, Eastern Desert, Egypt. In Garofalo PS, Ridley JR (eds) Goldtransporting hydrothermal fluids in the Earth’s crust. Geological Society, London, Special Publications, 402. https://doi.org/10.1144/ sp402.3 Zouaghi T, Aboud E, El-Sawy EK (2019) Aeromagnetic constraints on the structural framework of Umm Farwah Fault System and associated Neoproterozoic accretion domains of the Asir Terrane; the Biljurshi area, Arabian Shield. Ann. Geophys. 62, GM569, https://doi.org/10.4401/ag-8022

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Low Angle Normal-Sense Shear Zones, Folds and Wrench Faults During the Post-Amalgamation Stage of the Arabian-Nubian Shield Ahmed S. A. A. Abu Sharib

Abstract

The Neoproterozoic East African Orogen (EAO) preserves one of the finest records of a complete Pre-Cambrian Wilson cycle that started with the fragmentation of Rodinia supercontinent (*870–800 Ma) and opening of the Mozambique Ocean, followed by convergence along easterly and/or westerly dipping subduction zones, and was completed by closure of the ocean basin and collision between a collage of continental fragments that comprised East and West Gondwana. A prolonged (*200 Ma) convergence culminated in arcs suturing and terranes accretion followed by arc-continent terminal collision to form a N–S oriented collision zone (EAO) differentiated into the Arabian-Nubian Shield (ANS) to the north and the Mozambique Belt to the south. Terrane accretion resulted in a substantial crustal thickening and differential uplift, of certain parts of the ANS, followed by erosion and deposition of thick cover sequences in a number of, mostly fault-controlled, depositional basins. The latter records a post-accretion complex array of tectonic—“active” and non-tectonic—“passive” related deformations. Active tectonics are attributed to the crustal shortening that accompanied terminal collision and wrenching whilst passive tectonics could be linked to extensional collapse of parts of the orogen and account for the low angle normal and reverse shear zones. Reactivation of pre-existing accretion-related lineaments, and the formation of wrench-related new shears might have created local and overlapping stress fields that resulted in variably oriented structures deviating drastically from the general stress field (s). The differential uplift across the ANS and the use of common criteria to interpret different tectonic regimes (e.g. extension versus wrenching), wherein the deformation events have overlapping, if not matching, dates, make the idea of a regional tectonic model of the ANS inapplicable. A. S. A. A. Abu Sharib (&) Geology Department, Faculty of Science, Beni-Suef University, Beni-Suef, 62521, Egypt

Moreover, in many parts of the shield, it is very plausible to interpret some post-accretion deformation events from different tectonic regimes perspective. Keywords











East African Orogen Neoproterozoic Extensional collapse Wrenching Differential uplift

16.1

Introduction

During most orogenic events, the deformation style changes from mainly compressional during the early stages to dominantly extensional and/or wrenching during the closing/ terminal stages (e.g. Mercier et al. 1987; Dewey 1988a, b; Dewey et al. 1988). However, in complex orogens, there is evidence of interplay between the three deformation styles. For example, oroclinal loops (e.g. Alpine oroclines) mark the interplay between compressional boundary forces and extension and result in thrust-decorated extensional basins (Dewey 1988b). Moreover, localized shortening and extension during widespread wrenching is very common and results in an array of temporally and spatially related complex structures. Late orogenic extension can be categorized into: tectonic and collapse. Tectonic extension is attributed to: (1) an overall change in the dominant plate boundary forces from compressional to tensional (e.g. the disruption of Pangea: Dewey 1988a), (2) subduction rollback, (3) wrench-related localized extension, and (4) orogen-parallel extension as a consequence of orogen-perpendicular shortening. Extensional collapse occurs when horizontal shortening produces over-thickened crust, wherein the vertical stresses due to load greatly exceed the horizontal stresses. Characteristically, most orogens had undergone substantial uplift, with variable rates, associated with late-to post-orogenic morphotectonic phases of significant extension and magmatism. During orogenesis, it is very common for the two types of extension to interfere and even

© The Author(s), under exclusive license to Springer Nature Switzerland AG 2021 Z. Hamimi et al. (eds.), The Geology of the Arabian-Nubian Shield, Regional Geology Reviews, https://doi.org/10.1007/978-3-030-72995-0_16

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enhance each other. The purpose of this chapter is to shed light on some of the taphrogenic events that took place in the ANS towards the waning stages of the Neoproterozoic evets responsible for the EAO together with the late-to post-orogenic wrenching and shortening.

16.2

Geologic Setting

The Arabian-Nubian Shield (ANS) is a greenschist to amphibolite facies-dominated collisional belt that, together with the granulite-facies-dominated Mozambique Belt (MB) (Fig. 16.1), form the N–S trending East African Orogen (EAO) (Stern 1994; Fig. 16.1). The latter formed following the closure of the Mozambique Ocean basin separating the continental fragments comprising East and West Gondwana *870–800 Ma (Li et al. 2007) and culminated in the collision between these continental fragments (Fig. 16.1). Formation of the EAO started during the late Cryogenian (*850–750 Ma) and ended during the Ediacaran (*600–550 Ma) (Li and Powell 1993; Meert 2002 and references therein; Collins and Pisarevsky 2005; Li et al. 2007; Pisarvesky et al. 2008; Eyal et al. 2014; Abd El-Rahman et al. 2016; Elisha et al. 2017; Abu Sharib et al. 2019) making it, despite the intermittent periods of quiescence of convergence, the longest documented orogeny in Earth history. The ANS is made up of a number of Neoproterozoic arc terranes juxtaposed along *NE–SW-and N– S-trending mega shear zones (Fig. 16.2) (Bentor 1985; Stern 1994, 2002; Stein and Goldstein 1996; Meert 2002; Stoecer and Frost 2006; Johnson et al. 2011; Fritz et al. 2013; Robinson et al. 2017). The former trend marks the accretion of the juvenile arcs (Western arc terranes of Johnson et al. 2011), whereas the latter trend marks the collision of the arc terranes (Eastern and Western of Johnson et al. 2011) against the Sahara metacraton to the west and the Arabian craton to the east (Fig. 16.2) (e.g. Stern 1994; Meert 2002; Johnson et al. 2011, and references therein). The prominence of ophiolitic rocks and fragments bounding some of the arc terranes, the lower grade of metamorphism, and the voluminous amounts of post-orogenic granitoids (Tables 16.1 and 16.2) are characteristic features that distinguish the ANS from the Mozambique Belt.

16.3

Post-Amalgamation Events

In the literature on the EAO, the term post-amalgamation refers to all the deformation events (including extension, wrenching and shortening), magmatic activities and depositional processes that post-dated terrane accretion. In a strict sense, apart from geochronological data that might help in distinguishing pre- and post-amalgamation deformation

events, structures related to post-amalgamation events crosscut and overprint the accretion-related structures. However, in the absence of a clear crosscutting relationship, it is difficult to distinguish between post-amalgamation and accretion-related shortening since both events produced the same type of structures mainly folds and thrusts, with similar orientations. Hence, there is a consensus among researchers interested in ANS geology to use the post-amalgamation depositional basins as a reference point. The latter are variably sized Neoproterozoic terrestrial, shallow marine and mixed terrestrial/shallow marine mostly structurally controlled depositional basins that are well developed in the north-eastern part of the ANS. The Hammamat (Egypt) and the Ablah, Murdama, Hibshi, Jibalah, Bani Ghayy and Fatima (Saudi Arabia) basins are the best-studied basins from the Nubian and Arabian shields, respectively. A common feature of most basins is the basal regional unconformity surface separating them from the underlying newly accreted arc terranes (Abdeen et al. 1992; Rice et al. 1993; Johnson 2003; Johnson et al. 2011). Moreover, whilst some basins are totally localized within a specific terrane, others formed across adjacent terranes or sub-terranes. For example, the Murdama group basins overlie, the Siham and Suwaj subterranes, of the Afif complex terrane (e.g. Johnson 2003), whilst the Bani Ghayy basins lie across the boundary between Afif and Asir terranes. All compressional deformation events that affected these basins will be assigned to the “post-amalgamation” shortening category. However, it should be kept in mind, as will be seen in the next sections, that these post-amalgamation events and processes are, to a great extent, interlinked in space and time in a rather complex manner. For example, wrenching was accompanied with localized zones of shortening and extension. The latter produced a number of pull-apart basins that have been potential depocenters for molasse and/or marine sediments, as well as sites for magmatic activities. On the contrary, late orogenic extension has been interpreted to facilitate wrenching (see below), and of course, played a significant role in the emplacement of the widespread extension-related intrusions.

16.3.1 Taphrogenic Event 16.3.1.1 Low Angle Normal Shear Zones (LANSZ) Normal faults and low angle normal-sense shear zones (detachments) are widespread, but typically are better developed in the Nubian sector of the ANS. LANSZ in Northern ANS Low angle normal shear zones have been recorded from the Eastern Desert and Sinai terranes. In the Central and Southern Eastern Desert of Egypt, the NW-trending Meatiq

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Low Angle Normal-Sense Shear Zones, Folds and Wrench …

395

Fig. 16.1 A general map showing the distribution of the Pan-African belts in the assembled East and West Gondwana continents

and Hafafit gneissic domes, respectively, are bounded to the north and south by NNW- and SSE-dipping normal shear zones, with NNW-and SSE-directed sense of shear, respectively (Rice et al. 1992; Wallbrecher et al. 1993; Fritz et al. 1996, 2000; Fowler and Osman 2009; Andersen et al. 2010). In both cases, the shear zones separate a footwall block comprising amphibolite facies ortho- and para-gneisses from an allochthonous greenschist facies succession of island arc rocks, ophiolitic melange and Hammamat sediments in the hanging-wall. The gneissic domes have been interpreted to represent metamorphic core complexes associated with crustal extension (Sturchio et al. 1983a, b; Fritz et al. 1996, 2002; Neumayr et al. 1996a, 1998; Loizenbauer et al. 2001). According to Neumayr et al. (1998), the N-and S-directed low angle normal faults in the Meatiq area played a significant role in the exhumation and updoming of the Meatiq core complex and were associated with a phase of low grade metamorphism, their M3, that has been dated at 580 Ma. To the west of the Meatiq dome, Fowler and El Kalioubi (2004) interpreted the NW-trending Neoproterozoic Um Esh-Um Seleimat nappe (an eugeoclinal succession comprising intensively foliated molasse-type sedimentary rocks and ophiolitic melange) as an extensional, top-to-the NW and low angle shear zone. Towards the extreme north-eastern tip

of the ANS in the Sinai Peninsula, a widespread subhorizontal, 1.5 km thick and NW-dipping normal shear zone with top-to the NW sense of shear has been recorded in Wadi Kid area (Blasband et al. 1997, 2000). The shear zone has been linked to a high geothermal gradient as evidenced by regional metamorphism at LP and HT (Blasband et al. 1997, 2000; Brooijmans et al. 2003) and has been interpreted to have developed during the exhumation of a metamorphic core complex similar to that in the Eastern Desert. However, as in the case of the Eastern Desert, compressional rather than extensional origin has been suggested for the low angle shear zones (Shimron 1984, 1987; Fowler et al. 2010). Despite the common belief and the growing body of evidence that the low angles shear zones are linked to core complex formation, the nature of the extension that resulted in exhumation of the core complex is rather controversial (see below). LANSZ in Southern ANS Towards the southernmost part of the ANS, extensional shear zones have been recorded from the Bulbul Belt, in southern Ethiopia (Tsige and Abdel Salam 2005). The *100 km long, 1–5 km wide N-trending and Chulul shear zone (southern extension of the Bulbul shear zone) have

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Fig. 16.2 A general map showing the different accreted terranes (arcs) the make up the Arabian-Nubian Shield (ANS) prior to the opening of the Red Sea Rift

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Low Angle Normal-Sense Shear Zones, Folds and Wrench …

397

Table 16.1 Geochronological data of some selected extensional-related dykes from the different terranes of the ANS

Nubian shield

Arabian Shield

Terrane

Eastern Desert

Midyan

Higaz

Locality

Age in Ma

Feiran area, Sinai (Egypt)

591 ± 9 Ma

Lithology

Stern and Manton (1987)

References

NE Sinai (Egypt)

591 ± 12 Ma

Stern and Manton (1987)

N Sinai (Egypt)

600–540 Ma

Kessel et al. (1998)

Eastern Desert (Egypt)

540–595 Ma

Stern and Hedge (1985)

Ajaj shear zone (Saudi Arabia)

573 ± 6 Ma

Kennedy et al. (2011)

Jordan

574 ± 11 Ma 564 ± 10 Ma

Jarrar et al. (1992, 2003)

Jordan

575 ± 6 Ma 545 ± 13 Ma

Jarrar (2001)

Wadi Araba Jordan

550 ± 13 Ma

Jarrar et al. (1991)

Jibalah Group Saui Arabia

577 ± 5 Ma

Kusky and Matsah (2003)

been interpreted to be a low angle oblique slip normal-sense shear with a top-to-southeast sense of shear as indicated by SE-plunging stretching lineation that overprints an earlier E-dipping mylonitic fabric (Tsige and Abdel Salam 2005). The footwall of the shear zone comprises amphibolite facies gneisses and migmatites of the Melka Guba domain, the southern extension of the Alghe terrane, whereas the hanging-wall is occupied by greenschist facies association of island arc, ophiolite and plutonic rocks constituting the Chulul domain, the southern extension of the Bulbul terrane. According to the authors, extension along the Chulul shear zone took place after a prolonged period of convergence that culminated in the collision between the Melka Guba and Chulul domains with the development of the pervasive E-dipping mylonitic foliation. In southern Ethiopia, the Bulbul belt is one of four low grade N-trending belts (others being Moyale, Megado and Kenticha) sandwiched between high grade gneisses and migmatites and dominated by volcano-sedimentary successions, wherein the maficultramafic rocks have been interpreted to be ophiolites marking Neoproterozoic suture zones (Tsige and Abdel Salam 2005, and the references therein). However, the Wilson cycle model proposed for the belts has been challenged by Warden and Horkel (1984), Ghebreab (1992) and Worku and Yifa (1992) who suggested an ensialic model wherein the mafic-ultramafic rocks intruded into the high grade gneisses and migmatites in an intracratonic rift basin that did not evolve into a passive plate margin. In eastern Eretria, Ghebreab and Talbot (2000) recorded late Pan-African sub-horizontal ductile to semi-ductile extensional shear zones with top-to-the northeast tectonic transport in the amphibolite facies gneisses and schists of the Ghedem domain that is separated from the structurally overlying volcano-sedimentary rocks of the Bizen domain

by a 2–3 km wide moderately to gently west-dipping transition zone. The shear zones are attributed to a phase of gravitational collapse that took place late during the second phase of Pan-African deformation (PAD2; Ghebreab and Talbot 2000). According to the authors, the sub-horizontal shear zones together with the flat-lying fabric (PAD2) were exploited during the Cenozoic NE–SW extension of the Red Sea in that they influenced the localization of the detachments and controlled the location of potential subsequent normal faults. The Oligocene dolerite sills were intruded along these reactivated structures. It is noteworthy that many low angle normal-sense shear zones overprint earlier mylonitic fabrics having senses of shear opposite to those shown by the normal shear zones. The latter observation implies reactivation of earlier reverse-sense shear zones as normal-sense shear zones as the stress regime changes from compression to extension.

16.3.2 Wrenching A common feature of many orogenies is the late orogenic wrenching. Typically, it follows terrane accretion and continent–continent collision and decorates the latter with an extensive array of strike-slip faults associated with transpressional and/or transtensional strain. For example, in the Tibetan plateau, Himalaya, the N–S shortening between Eurasia and India micro-plates during the mid-Eocene (*45 Ma) ended up with continent–continent collision that was followed during the Miocene by a significant wrenching that produced widespread conjugate strike-slip faulting (Dewey 1988b). Similarly, the EAO witnessed a significant period of late orogenic wrenching along with what is known as the Najd Fault System (NFS) that affects mainly the northern and north-eastern parts of the orogen

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A. S. A. A. Abu Sharib

Table 16.2 Geochronological data of some selected post-orogenic granitoids from the different terranes of the ANS

Nubian shield

Terrane

Locality

Age in Ma

Lithology

References

Gebel Gattar

576 ± 6 Ma, 594 ± 3

Younger granite

Youssef (2005), Moussa et al. (2008)

Um Had

596.3 ± 1.7, 595

Granite

Andersen et al. (2009), Abu Sharib et al. (2019)

Arieki

593 ± 3.1

Granite

Andresen et al. (2009)

Central Eastern Desert

550–540

Granite

Lundmark et al. (2011)

CED and SED (Egypt)

620–530

K-rich granites

Schmidt et al. (1979), Habib et al. (1985), Greiling et al. (1994)

Sinai (AL Suite)

608–580

Alkaline and peralkaline granite and monzodiorite

Be’eri-Shlevin et al. (2009)

Sinai

579–594

monzogranite, syeno— granite and alkali granite

Ali et al. (2009)

Sabaloka Central Sudan

506 ± 4 591 ± 5

High-K, calc-alkaline granodiorite shoshonitic granite

Abdel Salam and Stern (1985)

Ethiopia

606 ± 1 613 ± 1, 612 ± 6

Granodiorite granite

Miller et al. (2003), Avigad et al. (2007)

Afif

Abanat suite

570–585

peralkaline to peraluminous granites

Cole and Hedge (1986)

Ar Rayn

alkali granite suite

607 ± 6 and 583 ± 8

alkali granite

Doebrich et al. (2007)

Midyan

Araba complex (Jordan)

600–560

Eastern Desert

Haya

Arabian shield

cutting across the ANS in a NW–SE direction. The NFS is a brittle to ductile NW–SE striking crustal scale shear zone (1100 km long and up to 400 km wide) dominated by left-lateral strike-slip faults having a total net displacement of some 240–300 km (Fig. 16.3) (Brown and Jackson 1960; Delfour 1970; Brown 1972; Brown and Coleman 1972; Moore 1979; Schmidt et al. 1979; Davies 1984; Stern 1985; Johnson et al. 2011). Subordinate conjugate NE-trending right-lateral shears are documented (Moore 1979; Davies 1980; Johnson et al. 2011, and the references therein). The NW- and NE-trending faults are arranged in a parallel, en echelon and sinusoidal curved geometry (Moore 1979). Braided fault zone is very common when curved faults join or intersect together (Moore 1979). The inferences that the NFS extends south-eastward across the covered basement rocks of the Arabian plate into south Yemen, eastern Arabia, India and Iran (Brown 1972; Moore and Al-Shanti 1979; Stern and Johnson 2010; Al-Husseini 2000), southwards in the southern ANS and Mozambique Belt in Kenya and Madagascar (Raharimahefa and Kusky 2010), and north-eastward into Jordan El-Rabaa et al. 2001) with a total length of more than 2000 km (Moore 1979) make the NFS

Jarrar et al. (2003)

one of the greatest transcurrent fault zone on Earth (e.g. Johnson et al. 2011). Examples of NFS in the Arabian part of the shield include the sinistral NW-trending Qazaz-Ar Rika and the Halaban-Zarghat, and the dextral NE-trending Ad Damm shear zones (Johnson et al. 2011). Examples from the Eastern Desert terrane of the Nubian shield include the NW-trending Kharit-Hodein, Nugrus and Atalla shear zones, and the NE to ENE-trending Qena-Safaga and Mubarak-Barramiya conjugate shear belts (Hamimi et al. 2019). Based on superposition and overprinting criteria, the NFS affects and displaces the late Cryogenian-Ediacaran sutures that formed during terrane accretion, and hence, it is attributed to a late Ediacaran post-accretion event. As is the case in all strike-slip fault systems (e.g. Sylvester 1988), the NFS was accompanied with an array of secondary structures such as normal faults, thrusts, folds and oblique-and strike-slip faults. Along the course of the NFS, diverse zones of transpression, transtension and pure strike-slip faulting were produced (e.g. Fritz et al. 1996; Abd El-Wahed 2010; Abd El-Wahed et al. 2016; Stern 2017). The extensional domains controlled sediment accumulation within some post-accretion depositional basins and the sites of

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Low Angle Normal-Sense Shear Zones, Folds and Wrench …

extrusion of some volcanics and dyke swarms (Fig. 16.4) (see below). Time of Activity of the Najd Fault System (NFS) Generally speaking, the NFS is a late Proterozoic crustal scale transcurrent fault that has been active during the interval 620– 540 Ma (Fleck et al. 1976; Stern 1985). Criteria used to constrain this age range include (1) depositional age of sediments accumulated within syn-Najd basins and (2) crystallization age of syn- and post-shear igneous rocks. Compared to syn-shear plutons, which are slightly to intensively deformed and elongated parallel to strands of the Najd Fault System, post-shear plutons are non-deformed, circular in map-shape and discordant to the main fabric of the country rock. Strands of the NFS have been constrained between the pre-shear 567 ± 86 Ma Abu Aris granite, in central Arabia, and the post-shear 563 ± 71 Ma Tukhfah granite, in northeast Arabia, (Fleck and Hadley 1982). Elongated syn-shear plutons in Arabia are dated at 625–575 Ma (Duyverman et al. 1982; Hedge 1984). In the southern part of the shield, 577– 529 Ma felsic dykes cut across a set NW–SE-trending strands of the Najd Fault System (Fleck et al. 1979). The Jibalah group was deposited in pull-apart basins that formed due to movement along elements of the NFS (Hadley 1974; Husseini 1989; Al-Husseini 2000; Johnson 2003). The minimum depositional age of the group is constrained by 540 ± 18 Ma andesitic flows within the group (Brown 1972). Deposition in the Al Jifn basin (Jibalah Group basins) is bracketed between 625 ± 4 Ma and 576 ± 5 Ma based on the crystallization ages of rocks below the basin and a felsite dyke cutting through the basin-fill, respectively (Matsah and Kusky 2001). In Jabal Jibalah area, the maximum age of faults activity is constrained by the crystallization ages of 574 ± 28 (Calvez et al. 1984), and 567 ± 6 and 581 ± 1 (Brown et al. 1989) on granites forming basement to the group. The Al Junaynah Group, which unconformably overlies the *640 Ma granite, and was deposited along the N–S-trending Nabitah fault/shear zone, has been folded and dextrally sheared implying post 640 Ma fault activity. Similarly, to the west of the Nabitah fault and within the Asir terrane, the Ablah Group (641 ± 4–613 ± 7 Ma: Agar 1986; Doebrich et al. 2004) was deposited in N–S-trending marine basins that are intruded by Najd-related A-type granitoids dated at 617 ± 17 Ma and 605 ± 5 (Moufti 2001). In conclusion, from the geochronological data presented, it is clear that the last *100 Myr (*630–530 Ma) of the East African Orogeny was characterized by intense wrench tectonics manifested by the Najd Fault System that cuts mainly through the north and north-eastern parts of the Arabian shield and continues northwestwards into the central Midyan terrane of the Nubian shield.

399

16.3.3 Post-Amalgamation Shortening Based on overprinting and superposition criteria and the relative timing and style of deformation, the post-amalgamation structures linked to shortening are placed into two main categories: convergence—or transpressionrelated. This subdivision takes cognizance of the following facts and relationships: (1) some depositional basins have been multiply deformed; thus, the rocks in these basins have been affected by more than one phase of folding; (2) the rocks in some depositional basins are lie with unconformity on deformed rocks in older basins; (3) in some basins, the folds are arranged in an en echelon pattern indicative, interpreted to be the result of late shearing.

16.3.3.1 Convergence-Related Shortening In the literature on the ANS, the convergence-related shortening refers to all E–W to ENE–WSW shortening events that accompanied the Ediacaran terminal collision between the continental fragments of East and West Gondwana. Structures related to this tectonic event are recorded in basins of the Arabian and Nubian sectors of the ANS. Basins of the former include Murdama Group (Agar 1988; Kattan and Harire 2000), Bani Ghayy (Johnson 2003), Ablah and Fatima (Hamimi et al. 2012, 2014). Example of the Nubian basins includes Wadi Himur in the Southern Eastern Desert. In the Arabian basins, the shortening produced mesoscopic gently plunging open to isoclinal and intrafolial upright, N– S-trending folds with a well-developed vertical to steeply E-dipping axial planar foliation. Local W-verging overturned folds are not uncommon. Linear fabric elements include variably oriented mineral and stretching lineations. In Wadi Himur in the Allaqi-Heiani suture, slices of siliceous marble intercalated with thin conglomerate beds are in thrust contact with a thrust nappe comprising arc volcanics and ophiolitic serpentinite along the western side and ophiolitic amphibolite along the eastern side of the wadi. The shortening produced meso-to macroscopic, gently NNW-plunging, NNW–SSE-trending and asymmetric folds. 16.3.3.2 Transpression-Related Shortening The deformation style, type of strain and arrangement of compressional ductile structures related to this shortening event indicate that the prevailing tectonic regime was regional-scale transpression, wherein bulk horizontal shortening components were related to regional shearing. Shear-related shortening can be categorized into easterly and northerly directed. N–S to NNW–SSE Shortening Preserved in Fatima basin, the northerly directed shortening produced a group of kinematically and geometrically related

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Fig. 16.3 A schematic diagram shows the differentially uplifted Arabian-Nubian Shield crosscut by the NW–SE-trending Najd Fault System (NFS)

thrusts and folds. This deformation event is manifested by fold-initiated thrusts and thrust-related NW to NNW- and occasionally N-verging overturned folds with gently NE to ENE plunging axes (Hamimi et al. 2014). E–W to ENE–WSW to NE–SW Shortening Structures related to this shortening event are recorded in Ablah and Maslum (Murdama) basins, in eastern and southern Arabia, and in Hammamat basins (in wadis Queih, Hodein and Um Gheig), the Central Eastern Desert, Egypt. In wadi Yiba, Ablah basin, the shortening produced N to NNW-oriented thrusts and W-verging, thrust-related

overturned folds (Hamimi et al. 2014). In the southwestern margin of Maslum basin (Murdama Group basins), it produced en echelon NW–SE-trending folds that have been attributed to ductile shearing along the Ar Rika fault, one of the strands of the Najd Fault System (Johnson 2003). In some of the Hammamat basins, for example, in wadi Queih, this event produced NE-verging folds and the SW-dipping thrusts that accompanied a positive flower structure formed as a consequence of movement along Najd-related NW– SE-trending sinistral strike-slip faults (Abdeen et al. 1992; Abdeen and Warr 1998; Abdeen 2003; Abdeen and Greiling 2005). Shear sense criteria indicate sinistral shear in the

16

Low Angle Normal-Sense Shear Zones, Folds and Wrench …

401

Fig. 16.4 A tectonic model shows pull-apart basins formed between overstepped and overlapped Najd strands that cut across the differentially uplifted Arabian-Nubian Shield. The carbonates and clastics are restricted to the subsided and uplifted parts of the shield, respectively

Hammamat and Maslum basins and dextral shear in the Ablah basin. In wadi Um Gheig, the Hammamat Group sedimentary rocks are affected by SW-dipping oblique slip reverse faults associated with open to tight horizontal to gently plunging NW–SE-trending, and NE-verging folds with axial planes ranging from vertical to inclined to horizontal. The inclined axial planes are parallel and sub-parallel to the NE-directed reverse faults (Abdeen 2003). In wadi Hodein area, Hamimi et al. (2014) mapped a postamalgamation sedimentary succession composed of acidic volcanics intercalated with conglomerate that is folded about meso- and macroscopic NW–SE-trending, NW- and SE-plunging folds parallel to major NNW–SSE-trending sinistral shear zones. Kinematic indicator associated with deformed pebbles in conglomerate is consistent with SSE-directed tectonic transport (Abdeen et al. 2008).

16.3.3.3 Superposition and Fold Interference Pattern Between Compressionand Transpression-Related Folds Although the rocks in most of the post-amalgamation basins are multiply deformed, only a few preserve more than one fold generation: to mention, the Murdama, Ablah, Fatima and some of the Hammamat basins. At meso- and macroscopic scales, superposition of the differently oriented folds that formed during the convergence- and transpressionrelated shortening produced conspicuous fold interference patterns. Kilometer scale dome and basin interference pattern (e.g. Ramsay 1967; Ramsay and Huber 1987) formed due to the superposition of NW- and NE-trending folds have been documented from the rocks of the Hammamat Group exposed between wadis Shihimiyya and Hammamat (Fowler and Osman 2001, Fig. 2). The same pattern has been

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A. S. A. A. Abu Sharib

recorded at the macroscopic scales in the rocks of the Hammamat Group in wadis Hammamat, Zeidun and Queih (Abdeen and Greiling 2005, Fig 2). In the Ablah Basin, superposition between the NW-verging transpressional and N–S-to NNW–SSE-trending compressional folds produced a well-developed interference pattern (Hamimi et al. 2014, Fig. 17d). It is noteworthy that in Fatima and Ablah basins, Hamimi et al. (2014) recorded a post-transpression local phase of folding, their F3, with structures plunging that is moderately to steeply SE-SSE and E. Superposition of F3 on the pre-transpression-F1 folds were also noted (Hamimi et al. 2014, Fig. 17e).

16.4

Discussion

16.4.1 Extension: Motives, Evidence, and Deformation vs. Collapse Extension There is a general agreement that during the Ediacaran, the ANS experienced significant N–S-to NW–SE extension (Stern et al. 1984, Stern 1985, 1988; Johnson et al. 2011, and references therein). However, the nature and cause of such extension have been and are still controversial. Moreover, the reasons that account for some extensional features in one segment of the shield turned out to be inapplicable in the other segment (see below). Extension in the ANS has been interpreted to be the result of three main mechanisms: extensional deformation; extensional collapse; and wrench-related extension. However, the possibility that more than one mechanism could have acted together in some areas cannot be excluded.

16.4.1.1 Extensional Deformation Extensional deformation refers to all structures formed as a consequence of crustal scale external stresses. At a regional scale, the stresses are most commonly related to relative plate motion either directly across a divergent plate boundary or indirectly when the stresses are the resultant of interaction of a number of plates. Extensional stresses when concentrated within a plate form a continental rift (e.g. the Gulf of Suez Rift: Patton et al. 1994; Khalil and McClay 2001) that may fully evolve and result in continental break-up and the formation of an ocean basin floored by oceanic crust (e.g. the Neoproterozoic splitting of Rodinia). Evidence of Extensional Deformation In addition to the low angle normal shear zones and detachments that have been dealt with in Sect. 16.3.1, other evidence in support of the extension includes dykes, extension-related depositional basins, post-orogenic A-type

granites and bimodal volcanics (e.g. Blasband et al. 2000; Johnson et al. 2011). NE–SW-trending post-tectonic dykes are widespread in the northern part of the ANS, particularly in the Eastern Desert and Midyan terranes, and represent unequivocal evidence of northerly extension (e.g. Stern, 1984, 1985; Genna et al. 2002). They have been recorded from NE Egypt and Sinai (Stern 1984; Stern and Gottfried 1986; Eyal and Eyal 1987; Stern and Manton 1987; Greiling et al. 1994; Abdel-Karim and El-Baroudy 1995) from Jordan (Jarrar 2001, Jarrar et al. 1992, 2004), and from Saudi Arabia (Dodge 1979; Clark 1985; Genna et al. 2002). Extensive geochronological data obtained from bimodal, composite, felsic, mafic and pegmatitic dykes from the Midyan and Eastern Desert terranes (Table 16.1) constrain the period of extension between *600 and *540 Ma. Sediment fill, geometry and orientation of, and syn-sedimentary structures within, some post-amalgamation depositional basins have been taken as evidence of late Ediacaran extension. For example, in the Eastern Desert, Egypt, NE–SW-oriented molasse-type Hammamat basins, some of which are bounded by NE–SW-trending or striking normal faults, have been interpreted to be extension-related (Grothaus et al. 1979; Stern 1985; Fritz et al. 1996). Similarly, the fault-bounded Saramuj Conglomerate Group, Wadi Araba, SW Jordan, comprises 600–550 Ma molasse-type sedimentary rocks, likely equivalent to the Hammamat Group deposited in extension-related NE–SW striking grabens (Jarrar 2001). Moreover, the group is crosscut by NE– SW-trending dykes implying that the dyking and sedimentary basin formation could have occurred during a single extension event. In Oman, at the northernmost tip of the ANS, Husseini (1989) attributed the fault-bounded NE– SW-oriented 620–580 Ma evaporite- and clastic-filled basins to NE–SW extension, thereby extending the upper time boundary of extension to 620 Ma. Extension-related post-orogenic magmatic activity was very common in the ANS. The mainly epi-to mesozonal intrusions are of variable dimensions and shapes, i.e. dykes, sills, stocks, plutons and batholith. A-type granites are very widespread (Stern and Hedge 1985; Beyth et al. 1994; Moghazi et al. 1998; Garfunkel 1999; Jarrar et al. 2003; Mushkin et al. 2003; Moussa et al. 2008). The spectacular and voluminous exposures of alkaline granite led Stoeser (1986) to state that the ANS contains one of the largest alkaline granite fields in the world. The plutons being late-to post-tectonic in age are isotropic and typically discordant to the main rock fabric of the metamorphosed country rocks. Exceptions include highly deformed igneous bodies that are either shear-related or have been affected by a later shearing. Geochemically, the granitoids are enriched in LILE and have mostly alkaline, peralkaline, peraluminous and subordinate calc-alkaline geochemical signature indicative of a

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post-orogenic within-plate tectonic setting. Compositionally, they comprise medium to high-K granodiorite, granite, shoshonitic granite and shoshonite in increasing order of abundance. The granitoids are interpreted to be mantle-derived and intruded in an attenuated crust with minimal contribution from older crustal components (Gillespie and Dixon 1983; Stern and Hedge 1985; Pegram et al. 1980; Beyth et al. 1994; Greiling et al. 1996). Extension-related calc-alkaline magmatism is common in orogens with a long history of subduction (Hooper et al. 1995). This accounts for the association of post-orogenic calc-alkaline and alkaline intrusions in the ANS. Extension-related bimodal volcanics are widespread throughout the ANS. In the At Tuwawiyyah Formation at the base of the Murdama group (650 Ma; Cole 1988) in the Maslum basin, bimodal volcanics of rhyolite, andesite, dacite and basalt composition are interbedded and inter-finger with a dominantly epiclastic succession of sandstone and conglomerate (Bois et al. 1975; Johnson 1996). Geochemically, the volcanic rocks have calc-alkalic and high-K calc-alkalic affinity, and their bimodal character implies extension-related sedimentation of the basal clastic units of the Murdama group (Johnson 2003). In the Hadha basin, the Bani Ghayy Group comprises bimodal rhyolitic, rhyodacitic, andesitic and basaltic volcanics interbedded with a clastic-dominated succession of conglomerate, greywacke, sandstone and siltstone (Johnson 2003). SHRIMP U-Pb zircon dating of the rhyolite yielded a crystallization age of 650 Ma (Johnson 2003), consistent with the maximum age of the Haml batholith (650–600 Ma) that intrudes the group. The bimodality of the volcanic rocks together with the fanglomerate depositional environment of the conglomerate, led Agar (1986) to interpret the group as having been deposited in fault-bounded (grabens) basins. In the northeast part of the Hibshi basin, bimodal volcanics of rhyolite and basalt composition are exposed.

16.4.1.2 Possible Models for Extensional Deformation In the ANS, two possible models may account for the extensional deformation: orogen-parallel extension and orogen perpendicular shortening, and wrench-related extension. Orogen-Normal Shortening and Orogen-Parallel Extension A common geologic phenomenon during many orogenies is the switch in tectonic regime upon indentation from dominantly compressional during orogen-normal shortening to dominantly extensional during orogen-parallel extension. For examples, the Tauern Window, Eastern Alps (Royden and Burchfield 1989; Ratschbacher et al. 1989; Scharf et al. 2013; Favaro et al. 2015) and the Tibetan plateau, Himalaya

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(Tapponnier et al. 1986; Royden et al. 1997). Characteristically, the extension results in tectonic unroofing, and the formation of a system of normal faults and related, variably sized, rifts are oriented perpendicular to the direction of extension (Molnar and Tapponnier 1978; Tapponnier et al. 1981, 1986; Favaro et al. 2015). In the ANS, the late-to post-orogenic E–W to ENE–WSW striking normal faults that bound the metamorphic core complexes (gneissic domes), and the ENE–WSW to NE–SW-oriented molasse basins, and dyke swarms can be related to the orogen-parallel extension (e.g. Stern 1985; Shalaby et al. 2006; Johnson et al. 2011). Moreover, a short-lasted (605– 595 Ma) orogen-parallel extension was proposed as a tectonic model to explain the mechanism of exhumation and formation of the metamorphic core complexes (Fig. 16.6) (Wallbrecher et al. 1993; Fritz et al. 1996; Loizenbauer et al. 2001). Preservation of the extension-related structures in the northern and north-eastern parts of the ANS, i.e. towards the periphery of the orogenic belt, strengthens the latter conclusion. Wrench-Related Extension Major transcurrent faults are very commonly associated with local domains of extension (e.g. Sylvester 1988, and references therein) that are located between overstepping divergent strike-slip faults (Fig. 16.5), at the releasing bends along curved strike-slip faults, and where the fault terminates along a horsetail splay (e.g. Sylvester 1988, and references therein). Depositional basins of variable sizes form in these extensional domains and range from sag bonds to pull aparts (Crowell 1974a, b; Schubert 1980; Garfunkel 1981; Mann et al. 1983; Şengör et al. 1985). The Jibalah group, Midyan terrane of the Arabian shield, is an excellent example of wrench-related basins. The Al Kibdi basin is a pull apart basin formed between two sinistral left-stepping strands of the Najd Fault System, whereas the Al Jifn basin is formed at the releasing bend of the Halaban-Zaraghat fault (Hadley 1974; Husseini 1989; Al-Husseini 2000).

16.4.1.3 Non-Tectonic Gravitational Collapse-Related Extension Gravitational instability and resultant collapse is a very common phenomenon in many orogenic belts (i.e. Canadian Cordillera; Olivier and Teyssier 2001; Tibetan Plateau of the Himalaya: Dewey 1988; Dewey et al. 1993; Variscides: Henk 1997; Basin and Range Province: Dewey 1988b, Braun and Beaumont 1989; Gondwana: Yang et al. 2019). Typically, it is a late orogenic process that follows an early stage of collision, thrust stacking and crustal thickening that ends up with orogenic uplift (Coney and Harms 1984; Platt 1986; Dewey 1988b). Orogenic collapse is generally controlled by the amount and rate of uplift, which over a period

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Fig. 16.5 a, b Schematic diagrams show the potential of formation of alternating domains of compression and extension between overstepped convergent and divergent strike-slip faults, respectively, during initial a and progressive b wrenching. The domains of compression and extension develop into thrusts and pull aparts, respectively. Progressive wrenching would lead to formation of pull apart basins and exhumation of the structurally lower metamorphic core complexes

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of millions of years and an uplift rate of 3–6 mm/yr which is very common in orogenic belts (Saini et al. 1978; Cliff et al. 1985; Copeland et al. 1987; Butler and Prior 1988) would result in a substantial topography. As a corollary, collapse takes place to re-equilibrate the overthickned crust back to a normal (*20–30 km) thickness (Dewey 1988a, b; Dewey et al. 1993). An elevation of 3 km has been taken as a threshold above which mountain belts start collapsing (Dewey 1988b). Factors that cause orogenic uplift include (1) crustal underplating, (2) lithospheric thinning due to mantle delamination and/or decoupling, and hotspot jetting (Dewey 1988b; Platt and England 1993) followed by (3) extension-related magmatic activity in the thinned crust (e.g. Lister and Davis 1989; Lister and Baldwin 1993; Warren and Ellis 1996) followed by the upward ascent of mantle material at the root of the orogen (Fritz et al. 1996, 2002; Neumayr et al. 1998; Bregar et al. 2002). Mechanically, orogenic collapse occurs when (1) the vertical compression caused by the weight of the over-thickened crust overcomes the horizontal compression acting across on it (i.e. rV > rH). In this case, the mountain belt would collapse under its own weight, (2) plate boundary forces switch from compressional to extensional (Variscides: Henk 1997), (3) there is a drastic decease in the strength and viscosity of the crust caused by rheological modifications wherein thermal relaxation, radiogenic (advective) heating and high geothermal gradient culminate in crustal melting and magma generation at middle to upper crustal levels (i.e. Vanderhaeghe and Teyssier 2001). Late orogenic collapse has been reported from different parts of the ANS. In wadi Kid area, South Sinai, Blasband et al. (1997, 2001), attributed the exhumation of the metamorphic core complex (see Sect. 16.3.1.1), together with the NE–SW-trending dykes and oriented depositional basins to collapse-related extension reminiscent to the Mesozoic and Early Cenozoic orogenic collapse in North America Cordillera (e.g. Sturchio et al. 1983a). They further concluded that gravitational instability and collapse were active during the late stages of Pan-African orogeny in the ANS. In a similar way in the Eastern Desert, Greiling et al. (1994) interpreted a *20 Ma period (575–595 Ma) of extensional collapse wherein normal faults, molasse basins and core complexes were formed. However, according to Greiling et al. (1994), the extensional collapse was not the last tectonic event to affect the Pan-African where they attributed the widespread NNW-directed thrusts and folds recorded in the greenschist facies-dominated eugeoclinal succession to a later phase of NNW–SSE shortening. The northerly directed thrusts were also overprinted by a phase of transpression. Fowler and El Kalioubi (2004) interpreted the top-to-the NW low angle shear zone in the Neoproterozoic cover nappe in the Eastern Desert, Egypt, as a collapse-related tectonic nappe emplaced

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by gliding-spreading mechanism to be a consequence of crustal thickening following the accretionary tectonics. The younger granites in the Eastern Desert have been interpreted to be extensional collapse-and/or rift-related intrusions (Greiling et al. 1994; Farahat et al. 2007; Moussa et al. 2008). According to Greiling et al. (1994), the extensional collapse event can be bracketed between 595 Ma and 575 Ma. Acceptance of the extensional collapse model to interpret the post-amalgamation extension-related features in the ANS is dependent on answering two main questions: Was there a sign of significant uplift that exceeded the threshold value? and was there taphrogenic extension-related significant magmatism? Regarding the question on uplift, it is important to establish whether the process of uplift and consequently the collapse was a regional event throughout the whole ANS or restricted to specific parts of the shield. The second question has been dealt with in Sect. (16.4.1.1) and will not be discussed further here. Was There a Significant Uplift Throughout the ANS? Following England and Molnar (1990), the term uplift refers to the displacement of rocks in an opposite direction to gravity. There is a growing body of evidence that during the late Cryogenian-Ediacaran, significant crustal thickening and uplift linked to arc-accretion took place in different parts of the ANS. Crustal stacking and thickening have been bracketed between 630 Ma and 600 Ma (Greiling et al. 1996; Kröner and Stern 2004) and between 597 Ma and 584 Ma (Fritz et al. 2002; Abdel-Naby et al. 2008; Abu El-Enen et al. 2016). In the text that follows, crustal thickening and uplift will be dealt with from three perspectives: structural, metamorphic and depositional. From the structural point of view, in Wadi Kid area, South Sinai, Egypt, Blasband et al. (2000) identified accretion-related upright isoclinal NE–SW-trending folds, which they interpreted as evidence of crustal thickening and hence uplift. Generally speaking, a high metamorphic grade of upper amphibolite facies conditions recorded in many metamorphic terranes has been related to crustal thickening during orogenesis (Spear et al. 1991). All high grade gneissic domes that have been interpreted as metamorphic core complexes in the Central and Southern Eastern Desert (El-Gaby et al. 1990; Wallbrecher et al. 1993; Fritz et al. 1996; Loizenbauer et al. 2001; Abd El-Naby et al. 2008) and Sinai (Blasband et al. 1997, 2000; Brooijmans et al. 2003; Abu-Alam and Stüwe 2009; Abu El-Enen and Whitehouse 2013) have been metamorphosed at upper amphibolite facies conditions. The PT conditions have been estimated for some of these core complexes. In the Meatiq dome, Neumayr et al. (1998) estimated metamorphic conditions of 610–690 C at 6–8 kbar. Moreover, pressure in excess of 8 kbar has been interpreted based on kyanite relics (Neumayr et al. 1998). In

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the South Eastern Desert, collision-related crustal stacking and thickening have been documented in the Hafafit metamorphic core complex, and based on geothermobarometry and pseudosection calculations, a volcano-sedimentary succession has been metamorphosed at upper amphibolite facies conditions of 570–675 C and 9–13 kbar (Abu El-Enen et al. 2016). According to these authors, the metamorphism and deformation were related to the collision between East and West Gondwana with the metamorphic peak reached at an average crustal depth of 33 km, followed by a phase of isothermal decompression that they related to the rapid exhumation accompanied wrenching. Age data from Fritz et al. (2002) and Abd El-Naby et al. (2008) indicate the core complex was exhumed between 593 Ma and 580 Ma. In a similar fashion, Abu-Alam and Stüwe (2009) and Abu El-Enen (2011) interpreted the Feiran-Solaf metamorphic core complex, central Sinai, to have formed during a single PT path with a metamorphic peak estimated at temperature of *700 C and pressure of 7–9 kbar, followed by a phase of isothermal decompression. According to the authors, a transpressive tectonic regime accounts for the concurrent isothermal decompression and the associated shortening. Additionally, migmatitic biotite and hornblende gneisses of the Feiran Solaf core complex recorded upper amphibolite facies conditions of 640–700 C and 4–5 kbar (Eliwa et al. 2008). The type of post-amalgamation sedimentary basin, the nature of basin-fill and the nature of contact between the basin and the underlying basement rocks provides a direct clue for significant uplift. The fault-bounded Hammamat sedimentary basins are dominantly filled with molasse-type sedimentary facies (Akaad and Noweir 1969, 1980; Grothaus et al. 1979; Eliwa et al. 2006; Shalaby et al. 2006; Abd El-Wahed 2010) indicative of rapid uplift and erosion of the source area (e.g. Mitchell and Reading 1978; Miall 1978). In wadis Hammamat (the type locality) and Kareem in the Central Eastern Desert, the basins are filled, respectively, with 4000 and *7000 m thick succession of clastic rocks composed dominantly of polymictic conglomerate, sandstone, greywacke and siltstone (Akaad and Noweir 1969 1980; Fritz and Messner 1999; Abd El-Wahed 2009). Moreover, as a sign of sedimentation in association with tectonism, particularly extensional, some of the Hammamat basins were interpreted to be intermontane basins (Fritz et al. 1996; Fritz and Messner 1999; Abd El-Wahed 2010, and references therein) bounded by normal faults as well as intra-basinal syn-sedimentary normal faults (Grothaus et al. 1979; Fritz et al. 1996). All post-amalgamation basins contain thick piles of sedimentary, with subordinate volcanics, rocks the base of which rest unconformably on basement rocks composed of the accreted arc terranes (Johnson 2003; Johnson et al. 2011; Hamimi et al. 2014, and references therein) pointing to the autochthonous character of the

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basins. Sediments shed from the uplifted and eroded basement rocks were deposited as detritus in these basins. Johnson (2003) inferred uplift and erosion in the range of 10–15 km prior to deposition of the 8000 m thick, detrital-dominated, Murdama Group which lies unconformably on greenschist—to amphibolite-, and locally, granulite-facies rocks of the Suwaj and Siham sub-terranes of the Afif complex (Cole 1988). In the framework mineralogy triangular diagram for provenance discrimination, the sandstone (Zaydi Formation) of the Maslum basin (Murdama group basins) plots in the magmatic arc field (Greene 1993; Johnson 2003). The latter result supports the notion that the underlying highly eroded arc terranes were a significant source of detritus for the post-amalgamation basins. The Hibshi formation (Hibshi basin) is a thick,  5000 m, succession of volcanic, volcaniclastic and epiclastic rocks, the base of which unconformably overlies the Ha’il basement terrane on the north (Johnson 2003). The basin has been interpreted to be a fault-controlled continental basin (Williams et al. 1986). The formation occupies a faulted syncline structure, which led Johnson (2003) to interpret it as representing a major subsidence. Rocks of the Antaq and Al Kibdi basins (Jibalah basin) unconformably overlie the Suwaj terrane on the west and Siham arc, respectively (Johnson 2003). The Jurdhawiyah Group, in the Safih basin, oversteps the erosive contact between the Murdama Group and the underlying arc rocks of the Suwaj sub-terrane (Johnson 2003). Similarly, the Fatima Group (Fatima basin) rests unconformbly on >757 Ma basement rocks comprising high grade gneisses, amphibolites, schist and andesite (Hamimi et al. 2014). By contrast, significant carbonate units of considerable thickness have been recorded in some basins. Variably sized limestone lenses (tens of meters to 300 m thick) and a succession of 1000 m thick of carbonaceous and stromatolitic limestone of the Farida formation are recorded in, respectively, the Marghan and Maslum basins of the Murdama group basins (Johnson 2003). Limestone (1000 m thick) and 150–800 m thick carbonate sequence including oolitic, stromatolitic limestone and dolomite (Kattan and Harire 2000) are recorded in the Mujayrib and Hadha basins, respectively, of the Bani Ghayy group basins. Moreover, the greywacke and siltstone of the Mujayrib basin (Bani Ghayy group basins) bear the sedimentological features characteristic of a Bouma sequence (Agar 1986). The above-mentioned features (carbonate units and Bouma sequence) are indicative of deposition in a sub-aqueous shallow marine conditions (Agar 1986; Wallace 1986) and imply that parts of the eastern Arabian shield have been submerged under marine water prior to or during the deposition of the Murdama and Bani Ghayy groups (e.g. Johnson 2003). In a similar fashion, the Fatima Group, Jeddah terrane, is differentiated into three units (Nebert et al. 1974): a shallow marine-platform carbonate-dominated

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middle unit (Basahel et al. 1984) that is composed essentially of stromatolitic fossiliferous limestone, and sandwiched between two clastic-dominated units composed of sandstone, siltstone and conglomerate (Hamimi et al. 2014). Similarly, the Ablah group, Asir terrane, is composed of two marble units (lower and upper) interbedded with two clastic units (Hamimi et al. 2012, 2014). The carbonate-dominated units are interpreted as having been deposited in shallow water, sub-to intertidal environments (Hamimi et al. 2014). It is noteworthy that there is a strong evidence of a second cycle of erosion and uplift after deposition of the Murdama and Bani Ghayy groups and prior to the deposition of the younger Jibalah, Hibshi and Jurdahwiyah groups. This is indicated by (1) folding of the Murdama and Bani Ghayy groups, (2) presence of an erosive surface between the folded Murdama group and the overlying Jurdhawiyah group and (3) the exhumation of rapidly cooled epizonal granites beneath the unconformities at the base of the post-Murdama/Bani Ghayy Jibalah, Hibshi and Jurdahwiyah basins (Agar 1986; Cole 1988; Al-Saleh et al. 1998). In conclusion, it is clear that, the ANS preserves strong evidence of periodic uplift during the Ediacaran. Rapid erosion of the uplifted amalgamated arc terranes provided vast amounts of detritus that filled the molasse-type Hammamat basins and constituted the clastic-dominated basal units of the Arabian basins. The Arabian shield probably experience at least two cycles of uplift separated by a period of rapid erosion, subsidence and deposition of the Murdama, Bani Ghayy, Fatima and Ablah groups in shallow marine environment. Intermittent and insignificant small cycles of uplift may account for the clasticdominated units interbedded with the shallow marine carbonate facies. However, deposition of the clastic and non-clastic successions in shallow marine environments cannot be excluded. On the other hand, the persistence of continental facies and scarcity of proper marine deposits within the molasse-type Hammamat basins imply that substantial pars of the Nubian shield had significant height prior to and during deposition of the Hammamat sediments (Figs. 16.3 and 16.4). Preservation of evidence of subsidence in the Arabian basins relative to those in the Nubian part of the shield can be interpreted in terms of differential uplift with greater uplift rates in the latter basins relative to the former ones (Fig. 16.3). Conversely, differential subsidence may also account for the drastic change in lithologies in both shields. In the latter case, compared to the Arabian shield, the subsidence was not that efficient in the Nubian region which remained high enough to ensure that no sea water covered the shield. The significant uplift, particularly, in the Nubian section, the fault-bounded (grabens) basins and the widespread coeval extensionrelated magmatic activities imply that gravitational extensional collapse model could have been active and explains

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the extensional features observed mostly in the post-amalgamation Hammamat basins in the Nubian shield, and to a much lesser extent, in some of the basins in the Arabian shield.

16.4.2 Wrenching One of the robust and widely accepted mechanisms to account for late orogenic wrenching is the “extrusion” or “escape tectonics” that was introduced by Molnar and Tapponnier (1975) to explain the widespread late tertiary (25–5 Ma) intracontinental right-lateral strike-slip faults in the southern part of the Tibetan plateau following terminal collision between the Indian and Eurasian plates at the beginning of the Miocene (*25 Ma). The model proposes that following collision is one of the continents (India) acted as an indentor facilitating the “escape” of the immediately adjacent sector of Eurasia eastwards towards the oceanic free-face. This escape is being accommodated along a system of orogen-parallel strike-slip and transform faults (Molnar and Tapponnier 1975). Hence, it is the position of the rigid indentor that determines the loci of crustal thickening and concentration of the strike-slip/transform faults, whereas the facing direction of the oceanic free-face determines the movement direction of the continental block and, consequently, the dominant sense of shear along the wrench/strike-slip faults. The same model was used by Dewey et al. (1986) to account for the westward escape of Anatolia from the Arabian/Eurasian collision zone. Similarly, in the ANS, the wrenching along the Najd Fault System has been interpreted in the light of the “escape” tectonics model (e.g. Johnson et al. 2011; Hamimi et al. 2019, and references therein). In that model, the terminal collision and indentation between East and West Gondwana were followed by tectonic escape of one of the continental plates and the widespread development of NW–SE striking sinistral strike-slip faults. According to Fleck (Fleck et al. 1979), Moore (1979) and Schmidt et al. (1979), the rigid indentor is positioned to the east (east of the Idsas suture), and hence, the NFS was formed as a consequence of the south-eastward movement of the Nubian (south of the Central Eastern Desert, Egypt)/south Arabian part of the shield relative to north Arabia and the Northern Eastern Desert of Egypt. However, geochronological (Glennie et al. 1974; Stacey et al. 1984) and isotopic (Duyverman et al. 1982) data, and the absence of proper evidence for a significant crustal thickening and high grade metamorphic rocks indicative of continent–continent collision preclude the existence of a continental rigid indentor east of Idsas suture (e.g. Stern 1985). It is more appropriate, for the aforementioned reasons, that during the terminal collision between East and West Gondwana, the Tanzanian craton acted as the rigid

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indentor (Bonavia and Chorowicz 1992) allowing the north-westerly escape of the ANS towards the Paleo-Tethys ocean (Fig. 16.1) (Stern 1994, Kusky and Matsah 2003). Based on: (1) the lack of significant uplift accompanied the Najd faulting, (2) the lack of the characteristic crustal thickening, particularly in the Arabian part of the shield, that accompanies the continent–continent collision, (3) the overlapping dates between the extension (rift)-related magmatism (see Sect. 16.4.1.1) and Najd wrenching, (4) synchrony between Najd faulting and suturing, and (5) the spatial and temporal relation between the Najd faults and the deposition of extension-related post-amalgamation sedimentary succession, Stern (1985) and Al-Husseini (2000) proposed an alternative kinematic model to account for the Najd faulting. In that model, the NFS was formed as a consequence of late Ediacaran NW–SE directed crustal scale extension within the newly formed EAO continental crust. The extended crust progressively developed into, mostly, preferentially oriented continental rifts, localized mainly to the north and north-eastern parts of the ANS (Stern 1985, 1994) that had acted as depocentres for the postamalgamation volcano-sedimentary successions and controlled the loci of the extension-related magmatic activities.

16.4.3 Shortening: Convergenceand Transpression-Related As documented in Sect. 16.3.3, the post-amalgamation shortening includes all collision- and wrench-related structures formed in response to directed tectonic-related compression. In terms of tectonics, we refer to this category as active tectonics-related structures to discriminate it from the non-tectonic category in which “contraction” structures might have formed passively (see below).

16.4.3.1 Active Tectonics-Related Structures There is agreement that during the Ediacaran, the terminal collision between East and West Gondwana took place along an E–W to ENE–WSW direction, and that pure shear was the main tectonic regime (Johnson et al. 2011; Fritz et al. 2013, and references therein). Coaxial ductile deformation accompanied this tectonic event include the N–S-to NNW– SSE-trending upright and locally W-verging folds in Murdama, Ablah, Bani Ghayy and Fatima group basins in the Arabian shield, and in a few localities in the Eastern Desert terrane of the Nubian shield. Despite the dominant pure shear regime, some collision-related transpressional structures have been recorded (e.g. Fatima basin; Hamimi et al. 2014). These structures are oriented E–W to ENE–WSW, which implies a shift of *90o from the dominant stress field (i.e. N–S versus E–W shortening). Such disturbances in stress field direction are mostly attributed to reactivation of

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favourably oriented, inherited, accretion-related and lineaments (faults) producing local stress fields. For example, in the Fatima basin, the accretion-related NE–SW-oriented wadi Fatima dextral shear zone (pre-dates deposition of the post-amalgamation Fatima Group) was dextrally reactivated during the E–W terminal collision and produced a localized N–S oriented shortening component that gave rise to the earliest structures preserved in the Fatima basin. Other possible examples of collision-related reactivation of pre-existing basement faults include the reactivation of the contemporary N–S-trending arc-arc, Oko and Hamisana, and arc-continent, Keraf, suture zones. The transpression structures in these shear zones; dextral in the Hamisana and sinistral in the Oko and Keraf (Abdelsalam 1994; Abdelsalam and Stern 1996; Johnson et al. 2011) overprint structures related to an earlier phase of collision-related E–W compression. It is noteworthy that some of the suture zones in the ANS have reactivation dates that are very comparable and/or overlap with some of the movements along the Najd faults (e.g. Johnson et al. 2011). In a similar way, reactivation of variably oriented basement structures would result in local stress fields with shortening components that apparently deviate from the dominant stress field. Other factors that could have contributed to the deviation of the stress field from the dominant E–W direction include deformation partitioning (e.g. Abu Sharib and Bell 2011; Bell and Sanislav 2011; Abu Sharib 2015) and reorientation of some structures during later events. For more than half a century, documented research has demonstrated that the Najd Fault System is a crustal scale, brittle to ductile shear zone dominated by NW–SE-trending sinistral strike-slip faults (Fig. 16.4) with a minor set of conjugate NE–SW-trending dextral strike-slip faults (Brown and Jackson 1960; Delfour 1970; Brown 1972; Brown and Coleman 1972; Moore 1979; Schmidt et al. 1979; Davies 1984; Stern 1985; Johnson et al. 2011). As in the case of major transcurrent faults, the NFS produced a complex array of superimposed structures that have classical en echelon arrangements (Sylvester 1988, and references therein). These structures include as follows: the Riedel (R) shears (synthetic strike-slip faults; loosely referred to as the Najd trend), conjugate Riedel (R′) shears (antithetic strike-slip faults), P-shears (synthetic strike-slip faults), Y-shears (synthetic strike-slip faults parallel to the principal displacement zone), in addition to folds and reverse faults, and normal faults that, respectively, form perpendicular and parallel to the incremental shortening axis. For a detailed description of strike-slip fault systems, their mechanics, and associated structures, the reader is referred to Sylvester (1988). Complexities of the Najd-related shortening arise from the fact that discrepancies in stress fields cause localized shortenings that produce very variably oriented structures, which require

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careful study and analysis to ensure they are not interpreted as due to overprinting rather than of a single major shear event. Moreover, generally speaking, the compression accompanied movement along the dominant NW-trending Najd strands is oriented in an almost E–W direction (e.g. Abdeen and Abdelghaffar 2011). The latter observation implies that in places, the Najd-and collision-related structures share a common shortening direction and requiring sound field data including unequivocal kinematic indicators to avoid an erroneous interpretation. Discrepancies in the stress fields could be attributed to: (1) local shortening components produced between two overlapping faults or along a restraining bend of an anastomosing fault during movements along the R, R′, Y and P-shears, (2) local domains of interfering shortening produced due to superposition of the shears, (3) local shortening that arise from the reactivation of favourably oriented pre-existing basement lineaments, and (4) the sequential development of shears (e.g. earlier dextral followed by later sinistral) could result in cases where early formed shear-related structures have structures related to the later shearing superimposed. For example, it is widely accepted that the Najd regime started with dextral shear and culminated with dominant sinistral movement. Shear-related structures characteristic of movement along the Najd strands include the en echelon arrangement of folds in the sedimentary rocks filling many of the post-amalgamation depositional basins (see Sect. 16.3.3.2) and distinctive NW–SE sinistral faults. In profile view, some of these faults form typical flower structures comprising a system of folds and thrusts (Abdeen et al. 1992; Abdeen 2003; Abdeen and Greiling 2005). Examples of reactivated basement faults include the N–S-trending Nabitah and Um Farwah shear zones. The Nabitah shear zone marks the suture between the Tathlith and Asir terranes that formed during the Nabitah orogeny (*680–640 Ma) and shows consistent dextral shearing (Johnson et al. 2011). Um Farwah, located within Asir terrane to the west of the Nabitah fault, is an ophiolitedecorated and accretion-related shear zone that has been reactivated during Najd activity. It delineates and intersects the boundary between the Ablah Group and the Asir arc terrane, at its northern and the central parts, respectively (Moufti 2001). In the Ablah Group, folds and thrusts orientation as well as fold vergence are consistent with sinistral sense of shearing along the reactivated Um Farwah shear zone (Johnson et al. 2011; Hamimi et al. 2014). The reactivation enhanced crustal melting and emplacement of the 617 ± 17 and 605 ± 5 Ma A-type granitoids (Moufti 2001) that intrude the Ablah Group. Other examples include the reactivation of the NW–SE to N–S-trending Hulayfah-Ad Dafinah-Ruwah suture zone (Johnson et al. 2011).

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16.4.4 The Eastern Desert Low Angle to Sub-Horizontal Shear Zone: A Shortening-(Tectonic) versus Extensional Collapse(Non-Tectonic) Related Fabric! A characteristic feature of the Neoproterozoic rocks of the Central and Southern Eastern Desert and Sinai, Egypt, is the presence of a number of tectonic windows exposing upper amphibolite facies gneissic (dominantly granitoid ortho-gneisses and quartzofeldspathic para-gneisses) domes that are elongated with their long axes oriented in a NW–SE direction and have been widely accepted as metamorphic core complexes (see Sect. 16.3.1). A major low angle to sub-horizontal mylonitic shear zone separates the high grade gneissic domes from a structurally overlying carapace of greenschist facies eugeoclinal units that represent thrust sheets (Ries et al. 1983; Sturchio et al. 1983a, b; Habib et al. 1985; Blasband et al. 1997, 2000; Fowler and Osman 2001) separated by sheared contacts (Ries et al. 1983). The eugeoclinal tectono-stratigraphic succession includes island arc associations (Stern 1981; Habib 1987; Abdel-Kari 1994), ophiolitic melange rocks (serpentinites, metagabbros and metavolcanics; El-Sharkawy and El-Bayoumi 1979; Nasseef et al. 1980; Shackleton et al. 1980; Ries et al. 1983; Ali et al. 2020) and Hammamat molasse sediments (Akaad and Noweir 1969; Grothaus et al. 1979; Messner 1996; Naim et al. 1996; Kamal El-Din et al. 1996; Abd El-Rahman et al. 2010; Abu Sharib et al. 2019). These are the “Pan-African Nappes” of Ries et al. (1983), Habib et al. (1985), and El Gaby et al. (1988), the “eugeoclinal thrust sheet/complex” of Andresen et al. (2009) and the “Eastern Desert Shear Zone” of Andresen et al. (2010). West of the Meatiq gneissic dome, the Hammamat Group rocks and the eugeoclinal succession is imbricated and folded together (Fowler and Osman 2001; Loizenbauer et al. 2001). The low angle shear zones are very crucial in the interpretation of the tectonic evolution of the Precambrian rocks in Egypt. However, despite the consensus regarding the tectonic transport direction with very consistent top-to-the NW to NNW shear sense criteria (Ries et al. 1983; Habib et al. 1985; Shackleton 1986, 1994; Greiling et al. 1994; Fritz et al. 1996; Greiling 1997; Loizenbauer et al. 2001; Shalaby et al. 2005; Abd El-Wahed 2008; Andresen et al. 2009; Abd El-Wahed and Kamh 2010; Abd El-Wahed 2014; Abdeen et al. 2014), the nature and kinetics of the shear zones are controversial. They have been interpreted as compression—(Ries et al. 1983; Sturchio et al. 1983), extension—(Fritz et al. 1996, 2002), and transcurrent-related (Loizenbauer et al. 2001) high strain zones, or as a result of interplay between two or more tectonic regimes (Loizenbauer et al. 2001). The variable tectonic models of the low angle shear zones have been used

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primarily to account for the exhumation of the spatially associated gneissic domes (op.cit.). NNW-directed thrusts and associated E–W to ENE– WSW-trending NW-verging folds affect the molasse-type Hammamat Group rocks in the Um Had and Queih basins of the Central Eastern Desert (Ries et al. 1983; Abdeen et al. 1992; Greiling et al. 1994; Abdeen et al. 1996; Naim et al. 1996; Abdeen and Warr 1998; Fowler and Osman 2001). Similarly, E–W-trending folds have been recorded in the Hammamat basins in wadis Karim, Messar, Um Gheig, and Zeidun, Central Eastern Desert (Greiling et al. 1994; Abdeen and Greiling 2005). Moreover, the Hammamat Group rocks in wadi Muweih have been affected by NNW-directed thrusts and associated foliation (Ries et al. 1983). At lower stratigraphic levels of the Hammamat succession, the deformation intensifies and results in the mylonitic and phyllonitic foliation (Ries et al. 1983; Fowler and Osman 2001). In wadi Hammamat, the type locality, a similar conclusion was reached by Messner (1996) who noted an increase in the degree of shearing and metamorphism towards the base of the Hammamat Group. From the metamorphic point of view, the low angle shear zones record a regional phase of retrogression (Fowler and Osman 2001). The debate regarding the kinetics of the low angle mylonitic shear zones can be settled if the compelling question of whether the shear zones are accretion-or post-accretion-related tectonic feature is convincingly answered. Answering this question would be of a great significance in interpreting the tectonic evolution of this essential part of the ANS. Except for Greiling et al. (1994), all the authors who advocated for compression as the prime factor in the formation of the shear zones, interpreted the NNW to NW-ward directed thrusts and the associated folds as accretion-related structures. As previously explained, the shear zone affects the sedimentary rocks of the Hammamat Group as a component of the eugeoclinal succession. Moreover, the molasse sequence has been recorded to be imbricated and folded together with the eugeoclinal succession (e.g. Fowler and Osman 2001; Loizenbauer et al. 2001). Similarly, in the Allaqi-Heiani suture zone, thin slices of post-amalgamation metasedimentary successions comprising marble, quartzite lenses and thin bands of conglomerate are tectonically interleaved and imbricated with an ophiolitic assemblage of amphibolite, serpentinites and island arc rocks (Abdeen and Abdelghaffar 2011). The tectono-stratigraphic units have been affected by accretion-related N–S shortening that produced the NNE-dipping low angle thrusts and associated SSW-vergent folds implying SSW-ward tectonic transport (Abdeen and Abdelghaffar 2011). The latter conclusion clearly contradicts the general consensus in the literature that deposition of the molasse-type Hammamat Group post-dated arc-accretion. Moreover and in terms of relative timing, in the kid

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metamorphic core complex, Sinai, Blasband et al. (2000) have dated the arc-collision event as well as the top-to-NW sub-horizontal shear zone. The former event has been dated at 750–650 Ma that is consistent with the published age range for arcs accretion (800–700 Ma) in different parts of the ANS (e.g. Bentor 1985; Stern and Hedge 1985; Stoeser and Camp 1985; Kröner et al. 1992; Stern 1993; Greiling et al. 1994; Abdelsalam and Stern 1996), whereas activity of the shear zone has been dated at 620–580 Ma. The geochronological results reveal explicitly that arc-accretion and shear zone formation are two irrelevant tectonic events that are temporally separated by at least 20 Ma. Moreover, the geochronological results have been augmented by field evidence where accretion-related folds are traversed by shear zone-related foliations (Blasband et al. 2000). In addition to that, the N–S to NNW–SSE shortening that produced the shear zones and associated structures does not match the E– W to ENE–WSW-directed compression that accompanied the terminal collision and/or the Najd transpression. Consequently, for the above-mentioned reasons, it is argued that the low angle shear zones represent a distinct late post-amalgamation extension-related event (Fig. 16.6) that was coeval with the other extension-related features (see Sect. 16.1) not related to transtensional domains along the Najd faults. In conclusion, the eugeoclinal nappe with the basal sub-horizontal dominantly top-to-the N to NNW-directed shear zones can be interpreted as a spreading nappe (e.g. Merle 1989; Fowler and El-Kalioubi 2004) with strain increasing dramatically downwards towards the basal contact with the gneissic domes, as well as along the internal shear zones (gliding surfaces) separating the different tectono-stratigraphic units of the eugeoclinal successions. The reverse sense of movement along these shear zones implies, contrary to Fowler and El Kalioubi (2004), a pure spreading rather than gliding-spreading mechanism. The nappe emplacement could have been synchronous with or immediately after the post-orogenic extension-related magmatism, which decreased significantly the shear strength of the crust and facilitated the spreading (e.g. Vanderhaeghe and Teyssier 2001). The overturned, intrafolial and thrust-related folds could have formed due to crumpling of the moving crustal blocks in the direction of spreading in a similar way overturned, and slump-related folds are formed in sliding soft sediment. The change in fold style from open to tight and from upright to overturned to recumbent may be attributed to the strain gradient and the degree of shearing, which in turn are directly related to its position within the spreading nappe. The strain gradients within a spreading nappe increases from top-to bottom and from rear to front (Merle 1989). The dominant top-to-the-N to NNW tectonic transport implies that the general slope responsible for the nappe emplacement was facing north. Moreover, the local

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Fig. 16.6 Tectonic model shows a complete Wilson cycle of the Arabian-Nubian Shield. a Break-up of Rodinia supercontinent and opening of the Mozambique Ocean. b Convergence through ocean–ocean and ocean–continent subductions and intrusion of syn-orogenic granitoids. c Suturing of arcs (terminal collision) and uplift of the accreted terrane. d Arc-continent collision and further uplift of the accreted terranes. e Extension (collapse-and/or deformation-related), intrusion of post-orogenic granitoids, crustal thinning and exhumation of the structurally lower metamorphic core complexes

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opposite (S-to SSE-ward) sense of shear along the thrusts and shear zones, and fold vergence most probably accompanied local irregularities in the paleo-slope and/or spreading in opposite directions around topographic highs (gneissic domes?). Accordingly, we interpret the NNW to N-directed thrusts, the low angle shear zones and E–W to ENE– WSW-trending folds as “passive” non-tectonic structures that formed as a consequence of gravitational instability and collapse of the orogenic belt during the late Ediacaran.

16.4.5 Inversion Tectonics and Basins Inversion Inversion tectonics refers to the reversal of slip direction along faults when the associated tectonic regime changes from extensional to compressional or vice versa (Cooper and Williams 1989). It is positive when normal faults are reactivated as reverse faults and negative when the opposite change takes place (Williams et al. 1989; Hayward and Graham 1989). Well-documented examples of inversion have been recorded in many orogenic belts: the Alps (Ziegler 1989; De Graciansky et al. 1989), the Variscides (Coward et al. 1989) and the Rockies (Powel and Williams, 1989). Reactivation of basin-bounding faults during positive inversion leads to closure and uplift of the basin in what is known as “basin inversion”. In the ANS, post-amalgamation basins are mostly fault-bounded (Johnson 2003; Johnson et al. 2011, and references therein), implying components of extension perpendicular to the strike of the bounding faults. Basin-fill sequences folded about axes trending parallel to the bounding faults are indicative of basin closure, inversion and uplift. For example, the Idayri Basin in the, Jurdhawiyah group basins, is bounded by the E–W trending Shara fault (Johnson 2003, Fig. 8) and is affected by E–W-trending anticlines, synclines and chevron folds (Johnson 2003). The Mujayrib Basin (part of the Bani Ghayy group basins) is a NNW-trending graben affected by N–S-trending folds having steep east-dipping axial planes parallel to steeply east-dipping reverse faults defining the eastern margin of the basin (Johnson 2003). Similarly, the Hammamat Group sedimentary rocks are affected by NW– SE-trending folds, and juxtaposed against a nappe complex (Um Gheig area: Abdeen 2003; Hamimi et al. 2014), Dokhan volcanics (wadi Atalla: Fowler and Osman 2001) and pelitic metasedimentary succession (wadi Um Had: Abu Sharib et al. 2019) by NE-and/or SW-dipping thrusts. In the case of the Hammamat basins, the parallelism between the direction of folds and thrusts and that of basins elongation imply that the bounding thrusts most probably may have been the original basin-bounding faults that were reactivated during inversion. The fact that the basin-fill sequences are folded about axes oriented parallel to the basin elongation

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direction reveals the same inversion axis, i.e. the axis of compression and extension did not change during inversion. The E–W oriented Jurdhawiyah basin unconformably overlying the folded rocks in the N–S oriented Murdama basin indicates a complex tectono-sedimentary history that involved basin opening, closure and uplift, and reopening with a 90o switch in stress field direction.

16.4.6 Complications of Tectonic Interpretations: A Regional Single Versus Multi-models In the previous sections, it has been demonstrated that the geochronological data for the dyke swarms, A-type granitoids, some of the post-amalgamation depositional basins, and the activity of faults in the Najd system are temporally within error. The synchrony of the extension-related magmatic activities (preferentially oriented dykes, bimodal volcanics, and A-type granites) and the activity of the Najd Fault System, as well as the spatial and temporal association of the depositional basins and extension-related bimodal volcanics, led Stern (1985) to link the formation of the Najd Fault System to late Proterozoic crustal scale N–S-directed continental extension, without significant crustal thickening. The manifestation of the extension is the E–W oriented grabens that controlled the deposition of Hammamat Group in Egypt and Hibshi and Jurdhawiyah groups in Arabia. On the other hand, Blasband et al. (1997, 2000) used the same criteria to confirm a late Proterozoic extensional collapse of a thickened crust and exhumation of core complex in the Nubian shield. Additionally, differential uplift across the ANS allowed the temporal deposition of clastic- and shallow marine-dominated successions in post-amalgamation basins, and hence the extensional collapse model, which could have been applicable to some parts of the shield, is inapplicable to other parts. Moreover and among the disagreements, the NFS has been interpreted as a crustal scale, indentationrelated, transcurrent fault system formed, as a consequence of tectonic escape (Johnson et al. 2011, and references therein) reminiscent to the intensive E–W oriented strike-slip fault system in Tibet and southern China (Dewey 1988; Dewey et al. 1988). For the above-mentioned reasons, many of the post-accretion features, including deformations, may fit into different tectonic models rather than being linked to a unique regional single tectonic model.

16.5

Conclusions

Post-accretion (arcs and terranes suturing) depositional basins in the ANS preserve a record of compression, extension and wrenching tectonic regimes that produced a

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complex array of temporally and/or spatially overlapping structures. Active tectonics involved east-west-directed shortening that accompanied the pure shear-related terminal collision between east and west Gondwana, and simple shear-related transpression due to subordinate collisionrelated reactivation of pre-existing basement faults and dominant activity along the Najd Fault System. Passive tectonics is attributed to the extensional gravitational collapse in the Nubian part of the shield and involved the spreading of a tectonic nappe along low angle to sub-horizontal reverse-sense shear zone and the formation of low angle normal-sense shear zones. Displacement along the Najd Fault System is attributed either to indentation and tectonic escape or to north-south-directed crustal scale continental extension. The Najd system is associated with a major component of north-south-directed extension, and hence, the wrenching and extension share common criteria including extension-related preferentially oriented dyke swarms, igneous (plutonic and volcanic) activity and deposition of volcano-sedimentary successions within faultbounded basins. The post-accretion deformation events across the ANS during the Ediacaran are the end result of the interplay between more than one tectonic regime that involved shortening, extension and wrenching in addition to a localized non-tectonic extensional gravitational collapse.

References Abd El-Naby H, Frisch W, Siebel W (2008) Tectonic–metamorphic evolution of the Wadi Hafafit culmination (central Eastern Desert, Egypt). Implication for Neoproterozoic core complex exhumation in NE Africa. Geologica Acta 6:293–312 Abd El-Rahman Y, Polat A, Fryer B J, Dilek Y, El-Sharkawy M, Sakran S (2010) The provenance and tectonic setting of the Neoproterozoic Um HassaGreywacke Member, Wadi Hammamat area, Egypt: Evidence from petrography and geochemistry. J Afr Earth Sci 58:185–196 Abd El-Wahed MA (2008) Thrusting and transpressional shearing in the Pan-African nappe southwest El-Sibai core complex, Central Eastern Desert, Egypt. J Afr Earth Sci 50:16–36 Abd El-Wahed MA (2010) The role of the Najd fault system in the tectonic evolution of the Hammamat molasse sediments, Eastern Desert, Egypt. Arab J Geosci 3:1–26 Abd El-Wahed MA (2014) Oppositely dipping thrusts and transpressional imbricate zone in the Central Eastern Desert of Egypt. J Afr Earth Sci 100:42–59 Abd El-Wahed MA, Harraz HZ, El-Behairy MH (2016) Transpressional imbricate thrust zones controlling gold mineralization in the Central Eastern Desert of Egypt. Ore Geol Rev 78:424–446 Abd El-Wahed MA, Kamh SZ (2010) Pan-African dextral transpressive duplex and flower structure in the Central Eastern Desert of Egypt. Gondwana Res 18:315–336 Abd El-Wahed MA (2009) The role of the Najd fault system in the tectonic evolution of the Hammamat molasse sediments, Eastern Desert, Egypt. Arabian Journal of Geosciences. https://doi.org/10. 1007/s12517-008-0030-0

413 Abdeen MM (2003) Tectonic history of the Pan-African orogeny in Um Gheig area, central Eastern Desert, Egypt. Egypt J Geol 47(1):239– 254 Abdeen MM, Greiling RO, Sadek MF, Hamad SS (2014) Magnetic fabrics and Pan-African structural evolution in the Najd Fault corridor in the Eastern Desert of Egypt. J Afr Earth Sci 99:93–108 Abdeen MM, Abdelghaffar AA (2011) Syn—and post-accretionary structures in the Neoproterozoic central Allaqi-Heiani suture zone, Southeastern Egypt. Precambr Res 185:95–108 Abdeen, M.M., Dardir, A.A., Greiling, R.O., 1992. Geological and structural evolution of the Wadi Queih Area (N Quseir), Pan-African basement of the Eastern Desert of Egypt. Zentralblatt fur Geologie und Palaeontolgie Teil I, 2653–2659 Abdeen MM, Greiling RO (2005) A quantitative structural study of late Pan-African compressional deformation in the Central Eastern Desert (Egypt) during Gondwana assembly. Gondwana Res 8:457– 471 Abdel-Karim AM (1994) Metavolcanics and Older Granites of the Central Eastern Desert of Egypt: a geochemical approach. Bull Fac Sci Zagazig University 16:338–361 Abdeen MM, Sadek MF, Greiling RO (2008) Thrusting and multiple folding in the Neoproterozoic Pan-African basemenet of Wadi Hodein area, south Eastern Desert, Egypt. J Afri Earth Sci 52:21–29 Abdeen MM, Warr LN (1998) Late orogenic basin evolution, deformation and metamorphism in the Pan-African Basement, Wadi Queih, Eastern Desert of Egypt. In: Naim GM, Hussein AA, Greiling RO, Warr LN (eds) International Cooperation Scientific Series, vol 42, 272 pp Abdel-Karim AM, El-Baroudy AF (1995) Petrology and geochemistry of felsic dyke swarms in SW Sinai, Egypt. Arab Gulf J Sci Res 13:239–250 Abdelsalam MG (1994) The Oko shear zone: post-accretionary deformations in the Arabian Nubian Shield. J Geol Soc, London 151:767–776 Abdelsalam MG, Stern RJ (1996) Sutures and shear zones in the Arabian-Nubian Shield. J Afr Earth Sc 23:289–310 Abdeen MM, Greiling RO, Dardir AA (1996) Two phases of Pan-African thrusting in the Proterozoic basement, North Eastern Desert of Egypt. In: Abstr Geol Surv Egypt Cent Conf, Cairo, Egypt, pp 5–6 Abu El-Enen MM, Abu-Alam TS, Whitehouse MJ, Ali K, Okrusch M (2016) P-T path and timing of crustal thickening during amalgamation of East and West Gondwana: a case study from the Hafafit Metamorphic Complex, Eastern Desert of Egypt. Lithos 263:213– 238 Abu El-Enen MM (2011) Geochemistry, provenance, and metamorphic evolution of Gabal Samra Neoproterozoic metapelites, Sinai, Egypt. J Afr Earth Sci 59:269–282 Abu El-Enen MM, Whitehouse MJ (2013) The Feiran-Solaf metamorphic complex, Sinai, Egypt: geochronological and geochemical constraints on its evolution. Precambr Res 239:106–125 Abu Sharib ASAA (2015) Changes in relative plate motion during the Isan Orogeny(1670–1500 Ma) and implications for Pre-Rodinia Reconstruction. Acta Geol Sin 90(1):88–105 Abu Sharib ASAA, Bell TH (2011) Radical changes in bulk shortening directions during orogenesis: significance for progressive development of regional folds and thrusts. Precambrian Res 188:1–20 Abu Sharib ASAA, Maurice AE, Abd El-Rahman YM, Sanislav IV, Schulz B, Bakhit BR (2019) Neoproterozoic arc sedimentation, metamorphism and collision: Evidence from the northern tip of the Arabian-Nubian Shield and implication for the terminal collision between East and West Gondwana. Gondwana Res 66:13–42 Abu-Alam TS, Stüwe K (2009) Exhumation during oblique transpression: the Feiran-Solaf region, Egypt. J Metamorphic Geol 27:439– 459

414 Agar RA (1986) The Bani Ghayy group; sedimentation and volcanism in ‘‘pullapart’’ grabens of the Najd strike-slip orogen, Saudi Arabian Shield. Precambr Res 31:259–274 Agar RA (1988) Geologic Map of the Zalm Quadrangle, Sheet 22 F, Kingdom of Saudi Arabia. Saudi Arabian Deputy Ministry for Mineral Resources Geologic Map GM 89, scale 1:250,000, p 41 Akaad MK, Noweir AM (1969) Lithostratigraphy of the Hammamat-Umm Seleimat district, Eastern Desert of Egypt. Nature 223:284–285 Akaad MK, Noweir AM (1980) Geology and lithostratigraphy of the Arabian Desert orogenic belt of Egypt between latitudes 25350 and 26300 N. King Abdulaziz University, Jiddah. Institute of Applied Geology Bulletin 3(4):127–134 Al-Husseini M (2000) Origin of the Arabian Plate Structures: Amar collision and Najd rift. GeoArabia 5:527–542 Ali RAM, Maurice AE, Pitcairn IK, Ahmed AH, Azer MK, Boskabadi A, Bakhit BR, Shahien MG (2020) Neoproterozoic and Cretaceous mantle oxidation states: Controls and heterogeneity through time. Lithos 356–357: Al-Saleh AM, Boyle AP, Mussett AE (1998) Metamorphism and 40Ar/39Ar dating of the Halaban ophiolite and associated units: evidence for two-stage orogenesis in the eastern Arabian shield. J Geol Soc, London 155:165–175 Andresen A, Abu El-Rus MA, Myhre PI, Boghdady GY, Corfu F (2009) U–Pb TIMS age constraints on the evolution of the Neoproterozoic Meatiq Gneiss Dome, Eastern Desert, Egypt. Int J Earth Sci. https://doi.org/10.1007/s00531-007-0276 Andresen A, Augland LE, Boghdady GY, Lundmark AM, Elnady OM, Hassan MA (2010) Structural constraints on the evolution of the Meatiq gneiss domes (Egypt), East-African Orogen. J Afr Earth Sc 57:413–422 Avigad D, Stern RJ, Beyth M, Miller N, McWilliams MO (2007) Detrital zircon U–Pb geochronology of Cryogenian diamictites and Lower Paleozoic sandstone in Ethiopia (Tigrai): age constraints on Neoproterozoic glaciation and crustal evolution of the southern Arabian–Nubian Shield. Precambrian Res 154(1–2):88–106 Basahel AN, Bahafzalah A, Omara S, Jux U (1984) Early Cambrian carbonate platform of the Arabian shield. Neues Jahrbuch für Geologie und Paläontologie, Monatshefte (N. Jb., Geol. Paläont. Mh.), 113–128 Bell TH, Sanislav IV (2011) A deformation partitioning approach to resolving the sequence of fold events and the orientations in which they formed across multiply deformed large-scale regions. J Struct Geol 33:1206–1217 Bentor YK (1985) The crustal evolution of the Arabo-Nubian massif with special reference to the Sinai Peninsula. Precambr Res 28:1–74 Beyth M, Stern RJ, Altherr R, Kröner A (1994) The late Precambrian Timna igneous complex, southern Israel: evidence for comagmatictype sanukitoid monzodiorite and alkali granite magma. Lithos 31:103–124 Blasband B, Brooijmans P, Dirks P, Visser W, White S (1997) A Pan-African core complex in the Sinai, Egypt. Geologie en Mijnbouw 76:247–266 Blasband B, White S, Brooijmans P, De Broorder H, Visser W (2000) Late Proterozoic extensional collapse in the Arabian-Nubian Shield. J Geol Soc, London 157:615–628 Bois J, Leca X, Shanti M (1975) Geology and Mineral Exploration of the Al Khushaymiyah Quadrangle, 22/44A. French Bureau de Recherches Géologiques et Minières Technical Record 75 JED 14, p 57 Bonavia FF, Chorowicz J (1992) Northward expulsion of the Pan-African of NE Africa guided by a reentrant zone of the Tanzania craton. Geology 20:1023–1026

A. S. A. A. Abu Sharib Braun J, Beaumont C (1989) Physical explanation of the relation between flanks uplift and the breakup unconformity at rifted continental margins. Geology 17:760–764 Brown GF, Coleman RG (1972) The tectonic framework of the Arabian Peninsula. 24th Int Geol Congr Montreal 3:300–305 Bregar M, Bauernhofer A, Pelz K, Kloetzli U, Fritz H, Neumayr P (2002) A late Neoproterozoic magmatic core complex in the Eastern Desert of Egypt: emplacement of granitoids in a wrench-tectonic setting. Precambr Res 118:59–82 Brooijmans P, Blasband B, White SH, Visser WJ, Dirks P (2003) Geothermobarometric evidence for a metamorphic core complex in Sinai, Egypt. Precambr Res 123:249–260 Brown GF, Jackson RO (1960) The Arabian shield. XIX international geological congress, Pt. 9:69–77 Brown FB, Schmidt DL, Huffman Jr, AC (1989) Geology of the Arabian Peninsula: Shield Area of Western Saudi Arabia. US Geological Survey Professional Paper 560-A, 188 p Brown GB (1972) Tectonic Map of the Arabian Peninsula. Saudi Arabian Directorate General of Mineral Resources Map AP-2, Scale 1:4,000,000 Butler RWH, Prior DJ (1988) Tectonic controls on the uplift of the Nanga Parbat Massif Pakistan Himalayas. Nature 333:247–250 Clark MD (1985) Late Protrozoic crustal evolution of the Midyan region, northwestern Saudi Arabia. Geology 13:611–615 Cliff RA, Droop GTR, Rex DC (1985) Alpine metamorphism in the south-east Tauern Window, Austria: 2. Rates of heating, cooling and uplift. J Metamorph Geol 3:403–415 Cole JC (1988) Geology of the Aban al Ahmar Quadrangle, Sheet 25F, Kingdom of Saudi Arabia. Saudi Arabian Deputy Ministry for Mineral Resources Geoscience Map GM 105, Scale 1:250,000, 45 p Cole JC, Hedge CE (1986) Geochronologic investigation of the Late Proterozoic rocks in the north-eastern shield of Saudi Arabia. Saudi Arabian Deputy Ministry for Mineral Resources Tech. Rec. USGS-TR- 05-5, 42 p Collins AS, Pisarevsky SA (2005) Amalgamating eastern Gondwana: the evolution of the Circum-Indian Orogens. Earth-Sci Rev 71:229–270 Coney PJ, Harms TA (1984) Cordilleran metamorhpic core complexes: Cenozoic extensional relics of Mesozoic compression. Geology 12:550–554 Copeland P, Harrison TM, Kidd WSF, Xu R, Zhang Y (1987) Rapid early Miocene acceleration of uplift in the Gangdese Belt Xizang (southern Tibet), and its bearing on accommodation mechanisms of the Indian-Asia collision, Earth Planet. Sci Lett 86:240–252 Calvez JY, Alsac C, Delfour J, Kemp J, Pellaton C (1984) Geological evolution of western, central, and eastern parts of the northern Precambrian Shield, Kingdom of Saudi Arabia. King Abdulaziz University, Jiddah, Faculty of Earth Science Bulletin 6, 24–48 Crowell JC (1974a) Origin of late Cenozoic basins in southern California. In: Dickinson WR (ed) Tectonics and sedimentation. Society of Economic Paleontologists and Mineralogists, Special Publication vol 22, pp 190–204 Crowell JC (1974b) Sedimentation along the San Andreas Fault, California. In: Dott RH, Shaver RH (eds) Modern and ancient geosynclinal sedimentation. Society of Economic Paleontologists and Mineralogists, Special Publication, 19, pp 292–303 Davies FB (1980) Reconnaissacne geology of the Duba Quadrangle, sheet 27/35D, King dome of Saudi Arabia, scale 1:100,000, Geol. Map GM Saudi Arabia Dir. Gen. Miner Resour GM-57, 23 pp Davies FB (1984) Strain analysis of wrench faults and collision tectonics of the Arabian-Nubian Shield. J Geol 82:37–53 Delfour J (1970) Le groupe de j’balah une nouvelle unite du bouchlier Arabe. Bureau de Recherches Géologiques et Minières Bulletin 2 (4):19–32

16

Low Angle Normal-Sense Shear Zones, Folds and Wrench …

Dewey JF (1988a) Lithospheric stress, deformation disruption deformation and tectonic cycles: The disruption of Pangea and the closure of Tethys Spec. Publ Geol Soc London 37:23–40 Dewey JF (1988b) Extensional collapse of orogens. Tectonics 7:1123– 1139 Dewey JF, Saroglu MR, Hempton WSF, Kidd F, Sengor AMC (1986) Shortening of continental lithosphere: The tectonics of eastern Anatolia-Z young collisional zone. Spec Publ Geol Soc London 19:3–36 Dewey FJ, Shackleton RM, Chang C, Yitin S (1988) The tectonic evolution of the Tibetan Plateau. Philos Trans R Soc London A. 327:379–413 Dodge FCW (1979) The Uyaijah ring structure, Kingdom of Saudi Arabia. US Geological Survey Professional Papers, 774-E Doebrich JL, Al-Jehani, Siddiqui AA, Hayes TS, Wooden JL, Johnson PR (2007) Geology and metallogeny of the Ar Rayn terrane, eastern Arabian shield: evolution of a Neoproterozoic continental-margin arc during assembly of Gondwana within the East African Orogen. Precambrian Res 158:17–50 Doebrich JL, Zahony SG, Leavitt JD, Portacio JS, Siddiqui AA, Wooden JL, Fleck RJ, Stein HJ (2004) Ad Duwayhi, Saudi Arabia. Geology and geochronology of a Neoproterozoic intrusion-related gold system in the Arabian shield. Economic Geology 99:713–741 Duyverman HJ, Harris NBW, Hawkesworth CJ (1982) Crustal accretion in the Pan-African: Nd and Sr isotope evidence from the Arabian Shield. Earth Planet Sci Lett 59:315–326 El-Gaby S, List FK, Tehrani R (1988) Geology, evolution and metallogenesis of the Pan-African Belt in Egypt. In: El Gaby S, Greiling R (eds) The Pan-African Belt of NE Africa and Adjacent Areas, Tectonic Evolution and Economic Aspects. Vieweg, Braunschweig, Wiesbaden, pp 17–68 El-Gaby S, List FK, Tehrani R (1990) The basement complex of the Eastern Desert and Sinai. In: Said R (ed) The Geology of Egypt. Balkema, Rotterdam, pp 175–184 Elisha B, Katzir Y, Kylander-Clark AK (2017) Ediacaran (620 Ma) high-grade regional metamorphism in the northern Arabian Nubian Shield: U-Th-Pb monazite ages of the Elat schist. Precambrian Res 295:172–186 Eliwa HA, Abu El-Enen MM, Khalaf IM, Itaya T, Murata M (2008) Metamorphic evolution of Neoproterozoic metapelites and gneisses in Sinai, Egypt: insights from petrology, mineral chemistry and K-Ar age dating. J Afr Earth Sci 51:107–122 Eliwa HA, Kimura J-I, Itaya T (2006) Late Neoproterozoic Dokhan Volcanics, N Eastern Desert, Egypt: Geochemistry and petrogenesis. Precambr Res 151:31–52 El-Rabaa SM, Al-Shumaimri MS, Al-Mishwat AT (2001) Fate of the Najd fault system in northwestern Saudi Arabia and southwestern Jordan. Gondwana Res 4:164–165 El-Sharkawy MA, El-Bayoumi RM (1979) The ophiolites of Wadi Ghadir area Eastern Desert, Egypt. Ann Geol Surv Egypt 9:125–135 Eyal M, Be’eri-Shlevin Y, Eyal Y, Whitehouse MJ, Litvinovsky B (2014) Three successive Proterozoic island arcs in the Northern Arabian-Nubian Shield: Evidence from SIMS U-Pb dating of zircon. Gondwana Res 25(2014):338–357 Eyal Y, Eyal M (1987) Mafic dyke swarms in the Arabian-Nubian Shield. Israeli J Earth Sci 36:195–211 Farahat ES, Mohamed HA, Ahmed AF, El Mahallawi MM (2007) Origin of I—and A-type granitoids from the Eastern Desert of Egypt: implications for crustal growth in the northern Arabian-Nubian shield. J Afr Earth Sci 49:43–58 Favaro S, Schuster R, Handy MR, Scharf A, Pestal G (2015) Transition from orogen-perpendicular to orogen-parallel exhumation and cooling during crustal indentation—Key constraints from 147Sm/144Nd and 87Rb/87Sr geochronology (Tauern Window, Alps). Tectonophysics 665:1–16

415 Fleck RJ, Greenwood WR, Hadley DG, Anderson RE, Schmidt DL (1979) Rubidium-Strontium geochronology and plate-tectonic evolution of the southern part of the Arabian Shield. US Geological Survey Saudi Arabian Mission Report 245:105 Fleck RJ, Hadley DG (1982) Ages and Strontium Initial Ratios of Plutonic Rocks in a Transect of the Arabian Shield. Saudi Arabian Deputy Ministry for Mineral Resources Open-File Report USGS-OF-03-38, 43 p Fleck RJ, Coleman RJ, Corn HF, Greenwood WR, Hadley DG, Schmidt DL, Prinz WC, Ratte JC (1976) Geochronology of the Arabian Shield, western Saudi Arabia’ K-Ar results. Geol Soc Am Bull 87:9–21 Fowler A, El Kalioubi B (2004) Gravitational collapse origin of shear zones, foliations and linear structures in the Neoproterozoic cover nappes, Eastern Desert, Egypt. J Afri Earth Sci 38:23–40 Fowler A, Hassen IS, Osman AF (2010) Neoproterozoic structural evolution of SE Sinai, Egypt: II. Convergent tectonic history of the continental arc Kid Group. J Afr Earth Sc 58:526–546 Fowler A, Osman AF (2009) The Sha’it–Nugrus shear zone separating Central and S Eastern Deserts, Egypt: a post-arc collision low-angle normal ductile shear zone. J Afr Earth Sc 53:16–32 Fowler TJ, Osman AF (2001) Gneiss-cored interference dome associated with two phases of late Pan-African thrusting in the Central Eastern Desert, Egypt. Precambrian Research 108:17–43 Fritz H, Abdelsalam M, Ali KA, Bingen B, Collins AS, Fowler AR, Ghebreab W, Hauzenberger CA, Johnson PR, Kusky TM, Mace P, Muhongo S, Stern RG, Viola G (2013) Orogen styles in the East African Orogen: a review of the Neoproterozoic to Cambrian tectonic evolution. J Afr Earth Sci 86:65–106 Fritz H, Dallmeyer DR, Wallbrecher E, Loizenbauer J, Hoinkes G, Neymayr P, Khudeir AA (2002) Neoproterozoic tectonothermal evolution of the Central Eastern Desert, Egypt: a slow velocity tectonic process of core complex exhumation. J Afr Earth Sc 34:137–155 Fritz H, Messner M (1999) Intramontane basin formation during oblique convergence in the Eastern Desert of Egypt: magmatically versus tectonically induced subsidence. Tectonophysics 315:145– 162 Fritz H, Wallbrecher E, Khudeir AA, Abu El Ela F, Dallmeyer DR (1996) Formation of Neoproterozoic metamorphic core complexes during oblique convergence (Eastern Desert, Egypt). J African Earth Sci 23:311–323 Fritz H, Wallbrecher E, Loizenbauer J, Hoinkes G, Dallmeyer DR, Khudeir AA (2000) Crustal evolution during oblique convergence, transpression and extrusion within the central Eastern Desert of Egypt: a slow velocity process. J Afr Earth Sc 33:33–34 Garfunkel Z (1981) Internal structure of the Dead Sea leaky transform (rift) in relation to plate kinematics. Tectonophysics 80:81–108 Garfunkel Z (1999) History and paleogeography during the Pan-African orogen to stable platform transition: reappraisal of the evidence from the Elat area and the northern Arabian-Nubian Shield. Israeli J Earth Sci 48:135–157 Genna A, Nehlig P, Le Goff E, Guetrrot C, Shanti M (2002) Proterozoic tectonism of the Arabian Shield. Precambrian Res 117:21–40 Ghebreab W (1992) The geological evolution of the Adola Precambrian Greenstone Belt, southern Ethiopia. J Afri Earth Sci 14:457–469 Ghebreab W, Talbot CJ (2000) Red Sea extension influenced by Pan-African tectonic grain in eastern Eritrea. J Struct Geol 22:931– 946 Gillespie JG, Dixon TH (1983) Isotope systematics of some granitic rocks from the Egyptian Shield. Precambrian Res 20:63–77 Glennie KW, Boeuf MG, Hughes Clarke MH, Moody-Stuart M, Pilaar WF, Reinhart BM (1974) Geology of the Oman Mountains, vol. 1, 432 pp., Verhandelin gen van het Kononklijk Nederlands geo logisch mijnbouwkundig Genootschap, Amsterdam

416 Greene RC (1993). Stratigraphy of the Late Proterozoic Murdama group, Saudi Arabia. US Geological Survey Bulletin 1976, 59 p Greiling RO (1997) Thrust tectonics in crystalline domains: the origin of a gneiss dome. Proceedings of the Indian Academy of Science (Earth and Planetary Sciences) 106:209–220 Greiling RO, Abdeen MM, Dardir AA, El-Akhal H, El-Ramly MF, El-Din K, Osman GM, Rashwan AF, Rice AA, Sadek AHN (1994) A structural synthesis of the Proterozoic Arabian-Nubian Shield in Egypt. Geologische Rundshau 83:484–501 Greiling, R.O., Naim, G.M., Hussein, A.A., (Eds.), 1996. Excursion across the Pan-African, Neoproterozoic basement Qena–Quseir Guide Book. Centennial of the Egyptian Geological Survey 1896– 1996. Egyptian Geological Survey and Mineral Authority, Cairo, Egypt Grothaus, B., Eppler, D., Ehrlish, R., 1979. Depositional environments and structural implications of the Hammamat Formation, Egypt. Annals of the Geological Survey of Egypt IX, 564–590 Habib ME (1987) Arc ophiolites in the Pan-African basement between Meatiq and Abu Furad, Eastern Desert. Assiut University, Egypt. Bulletin of the Faculty of Science, pp 241–283 Habib ME, Ahmed AA, El Nady OM (1985) Tectonic evolution of the Meatiq infrastructure, Central Eastern Desert. Egypt. Tectonics 4:613–627 Hadley, D.G., 1974. The Taphrogeosynclinal Jubaylah Group in the Mashhad Area, Northwestern Hijaz, Kingdom of Saudi Arabia. Saudi Arabian Directorate General of Mineral Resource Bulletin 10, 18 p Hamimi Z, Abd Elahed M, Gahlan H, Kamh S (2019) Tectonics of the Eastern Desert of Egypt: key to understanding the Neoproterozoic evolution of the Arabian–Nubian Shield (East African Orogen). In: Bendaoud A, Hamimi Z, Hamoudi M, Djemai S and Zoheir B (eds) The Geology of the Arab World-AnOverview, 1–81 pp. https://doi.org/10.1007/978-3-319-96794-3_1 Hamimi Z, El-Fakharani A, Abdeen M (2014) Polyphase deformation history and strain analysis of the post-amalgamation depositional basins in the Arabian-Nubian Shield: Evidence from Fatima, Ablah and Hammamat basins. In: Fowler AR, Greiling RO and Abdeen MM (eds) Special issue: Arabian-Nubian Precambrian basement geology—progress and developments, J Afr Earth Sci 99 (1):64–92 Hamimi Z, Matsah MI, El-Shafei M, El-Fakharani A, Shujoon A, Al-Gabali M (2012) Wadi Fatima thinskinned foreland FAT belt: a post amalgamation marine basin in the Arabian Shield. Open J Geol 2:271–293. https://doi.org/10.4236/ojg.2012.24027 Hayward AB, Graham RH (1989) Some geometrical characteristics of inversion. Geological Society, London, Special Publications 1989, 44:17–39. https://doi.org/10.1144/gsl.sp.1989.044.01.03 Hedge CE (1984) Precambrian geochronology of part of northwestern Saudi Arabia, dom of Saudi Arabia, King U.S. Geol. Surv. Open File Rep, 84–381, 12 pp Henk A (1997) Gravitational orogenic collapse vs plate-boundary stresses: a numerical modelling approach to the Permo-Carboniferous evolution of CentralEurope. Geol Rundsch 86(1):39–55 Hooper PR, Bailey DG, McCarley Holder GA (1995) Tertiary calc-alkaline magmatism associated with lithospheric extension in the Pacific NW. J Geophys Res 100:10303–10319 Husseini MI (1989) Tectonic and depositional model of Late Precambrian-Cambrian Arabian and adjoining plates. Am Assoc Petrol Bull 73:1117–1131 Jarrar G (2001) The youngest Neoproterozoic mafic dyke suite in the Arabian shield. Mildly alkaline dolerites from S Jordan—their geochemistry and petrogenesis. Geol Mag 138:309–323 Jarrar G, Stern RJ, Saffarini G, Al-Zubi H (2003) Late- and post-orogenic Neoproterozoic intrusions of Jordan: implications

A. S. A. A. Abu Sharib for crustal growth in the northernmost segment of the East African Orogen. Precambrian Res 123:295–320 Jarrar G, WachendorL H, Zellmer H (1991) The Saramuj Conglomerate: evolution of a Pan-African molasse sequence from southwest Jordan. Neues Jahrbuch Geologie Palaontologie Monatshefle 6:335–356 Jarrar G, Wachendorf H, Saffarini G (1992) A late Proterozoic bimodal volcanic/subvolcanic suite from Wadi Araba, SW Jordan. Precambr Res 56:51–72 Jarrar G, Saffarini G, Baumann A, Wachendorf H (2004) Origin, age and petrogenesis of Neoproterozoic composite dikes from the Arabian-Nubian Shield, SW Jordan. Geol J 39:157–178 Johnson PR (1996) Proterozoic Geology of Western Saudi Arabia, East-central Sheet. Saudi Arabian Deputy Ministry for Mineral Resources Open-File Report USGS-OF-96-4, scale 1:500,000, p 71 Johnson PR, Andresen A, Collins AS, Fowler AR, Fritz H, Ghebreab W, Kusky T, Stern RJ (2011) Late Cryogenian-Ediacaran history of the Arabian-Nubian Shield: a review of depositional, plutonic, structural, and tectonic events in the closing stages of the northern East African Orogen. J Afr Earth Sci 61:167–232 Johnson PR (2003) Post-amalgamation basins of the NE Arabian Shield and implications for Neoproterozoic tectonism in the northern East African Orogen. Precambr Res 123:321–337 Kamal El-Din GM, Rashwan AA, Greiling RO (1996) Structure and magmatism in a molasse-type basin Wadi Hammamat-Wadi Atalla areas. In: Greiling RO, Naim GM, Hussein AA (eds) Excursion across the Pan-African, Neoproterozoic basement Qena Quseir Guide Book. Egyptian Geological Survey and Mining Authority, Cairo, Egypt, pp 11–18 Kattan FH, Harire AR (2000) Geology of the Ad Duwayah-B’ir Warshah Mineral Belt: Saudi Geological Survey Open-File Report USGS-OF-99-6, 33 p Khalil SM, McClay KR (2001) Tectonic evolution of the NW Red Sea– Gulf of Suez rift system. In: Wilson RCL., Whitmarsh RB, Taylor B, Froitzheim N (eds) Non-Volcanic Rifting of Continental Margins: A Comparison of Evidence from Land and Sea.Geological Society, London, Special Publications: London, pp 453–473 Kröner A, Todt W, Hussein IM, Mansour M, Rashwan AA (1992) Dating of late Proterozoic ophiolites in Egypt and the Sudan using the single grain zircon evaporation technique. Precambr Res 59:15– 32 Kröner A, Stern RJ (2004) Pan-African orogeny. Encyclopedia of Geology 1:1–12 Kusky T, Matsah M (2003) Neoproterozoic Dextral Faulting on the Najd Fault System, Saudi Arabia, preceded sinistral faulting and escape tectonics related to closure of the Mozambique Ocean. In: Yoshida M, Windley BF, Dasgupta S, Powell C (eds) Proterozoic East Gondwana: Supercontinent Assembly and Breakup, vol 206. Geological Society. London, Special Publication, pp 327–361 Li ZX, Powell CM (1993) Late Proterozoic to early Paleozoic paleomagnetism and the formation of Gondwanaland. In: Findlay RH, Unrug R, Banks MR, Veevers JJ (eds) Gondwana Eight. Balkema, Rotterdam, pp 9–21 Li ZX, Bogdanova SV, Collins AS, Davidson A, De Waele B, Ernst RE, Fitzsimons ICW, Fuck RA, Gladkochub DP, Jacobs J, Karlstrom KE, Lu S, Natapov LM, Pease V, Pisarevsky SA, Thrane K, Vernikovsky V (2007) Assembly, configuration, and break-up history of Rodinia: a synthesis. Precambrian Res 160:179– 210 Lister GS, Baldwin SL (1993) Plutonism and the origin of metamorphic core complexes. Geology 21:607–610 Lister GS, Davis GA (1989) The origin of metamorphic gneiss domes and detachment faults formed during Tertiary continental extension in the northern Colorado River region, USA. J Struct Geol 7:65–93

16

Low Angle Normal-Sense Shear Zones, Folds and Wrench …

Loizenbauer J, Wallbrecher E, Fritz H, Neumayr P, Khudeir AA, Kloetzli U (2001) Structural geology, single zircon ages and fluid inclusion studies of the Meatiq metamorphic core complex: Implications for Neoproterozoic tectonics in the Eastern Desert of Egypt. Precambr Res 110:357–383 Lundmark AM, Andresen A, Hassan MA, Augland LE, Abu El-Rus MA, Baghdady GY (2011) Repeated magmatic pulses in the East African Orogen of Central Eastern Desert, Egypt: an old idea supported by new evidence. Gondwana Research Mann P, Hempton MR, Bradley DG, Burke K (1983) Development of pull-apart basins. J Geol 91:529–554 Matsah MIM, Kusky TM (2001) Analysis of Landsat ratio imagery of the alabanZarghat fault and related Jifn basin, NE Arabian shield: implications for the kinematic history of the Najd fault. Gondwana Res 4:182 Meert JG (2002) A synopsis of events related to the assembly of eastern Gondwana. Tectonophysics 363:1–40 Mercier JL, Armijo R, Tapponnier P, Carey-Gaithardis EC, v TL (1987) Change from late Tertiary compression to Quaternary extension in southern Tibet during the India-Asia collision. Tectonics 6:275–304 Merle X (1989) Strain models within spreading nappes. Tectonophysics 165:57–71 Messner M (1996) Basin analysis of Hammamat type molasse basins, Eastern Desert, Egypt. Proceedings of the Geological Survey of Egypt Centennial Conference 1996, Cairo, Egypt, pp 125–126 Miall AD (1978) Tectonic setting and syndepositional deformation of molasse and other non-marine paralic sedimentary basins. Can J Earth Sci 15:1613–1632 Miller CF, McDowell SM, Mapes RW (2003) Hot and cold granites? Implications of zircon saturation temperatures and preservation of inheritance. Geology (2003) 31(6):529–532. https://doi.org/10. 1130/0091-7613(2003)0312.0.CO;2 Mitchell AHG, Reading HG (1978) Sedimentation and tectonics. In: Reading HG (eds) Sedimentary Environments and Facies. Elsevier, New York, pp 439–476 Moghazi AM, Andersen T, Oweiss GA, El-Bouseily AM (1998) Geochemical and Sr–Nd–Pb isotopic data bearing on the origin of Pan-African granitoids in the Kid area, SE Sinai, Egypt. J Geol Soc, London 155:697–710 Molnar P, Tapponnier P (1975) Cenozoic tectonics of Asia: Effects of a continental collision. Science 189:419–426 Molnar P, Tapponnier P (1978) Active tectonics of Tibet. J Geophys Res 83:5361–5375 Moore JM (1979) Tectonics of the Najd transcurrent fault system, Saudi Arabia. J Geol Soc London 136:441–454 Moore JM, Al-Shanti AM (1979) Structure and mineralization in the Najd Fault System, Saudi Arabia. In: Tahoun S (ed) Evolution and Mineralization of the Arabian-Nubian Shield, 2. Pergamon, New York, pp 17–28 Moufti MRH (2001) Age, geochemicstry, and origin of peraluminous A-type granitoids of the Ablah-Shuwas pluton, Ablah graben Arabian shield. Acta Mineralogica-Petrographica, Szeged 2001 (42):5–20 Moussa EMM, Stern RJ, Manton WI, Ali KA (2008) SHRIMP zircon dating and Sm/Nd isotopic investigations of Neoproterozoic granitoids, Eastern Desert, Egypt. Precambr Res 160:341–356 Mushkin A, Navon O, Halitz L, Heimann A, Hartmann G, Stein M (2003) The petrogenesis of A-type magmas from the Amram Massif, southern Israel. J Petrology 44:815–832 Naim GM, Hussein AA, Dardir AA, Abdeen MM, Greiling RO, Warr LN (1996) The Wadi Queih area lithology and structure in a molasse-type basin. In: Greiling RO, Naim GM, Hussein AA (eds) Excursion across the Pan-African Neoproterozoic basement

417 Qena-Quseir Guide Book. Egyptian Geological Survey and Mining Authority, Cairo, Egypt, pp 65–78 Nasseef AO, Bakor AR, Hashad AH (1980) Petrography of possible ophiolitic rocks along the Qift-Quseir, Eastern Desert, Egypt. IAG Bull Jeddah 3:157–168 Nebert K, Alshaibi AA, Awlia M, Bounny I, Nawab ZA, Sharief OH, Sherbini OA, Yeslam AH (1974) Geology of the area north of Wadi Fatima, Kingdom of Saudi Arabia. Center for Applied Geology, Jeddah, Bulletin 1:31p Neumayr P, Mogessie A, Hoinkes G, Puhl J (1996a) Geological setting of the Meatiq metamorphic core complex in the Eastern Desert of Egypt based on amphibolite geochemistry. J Afr Earth Sc 23:331– 345 Neumayr P, Hoinkes G, Puhl J, Mogessie A (1996b) Polymetamorphism in the Meatiq basement complex (Eastern Desert, Egypt): P–T variations and implications for tectonic evolution. Geological Survey of Egypt, Centennial Symposium Abstracts, pp 139– 141 Neumayr P, Hoinkes G, Puhl J, Mogessie A, Khudeir AA (1998) The Meatiq Dome, (Eastern Desert, Egypt) a Precambrian metamorphic core complex: petrological and geological evidences. J Metamorphic Geol 16:259–279 Olivier V, Teyssier C (2001) Crustal-scale rheological transitions during late-orogenic collapse. Tectonophysics 335(1):211–228. https://doi.org/10.1016/S0040-1951(01)00053-1 Pisarevsky, S.A., Murphy, J.B., Cawood, P.A., Collins, A.S., 2008. Late Neoproterozoic and Early Cambrian palaeogeography: models and problems. In: Brito Neves,B.B., Trouw, R.A.J., de Wit, M.J., Pankhurst, R.J. (Eds.), Western Gondwana: PreCenozoic Geology Bordering the S Atlantic. Geological Society, London, Special Publication, pp. 9–32 Platt JP (1986) Dynamics of orogenic wedges and the uplift of high pressure metamorphic rocks. Bull Geol Soc Am 97:1037–1053 Raharimahefa T, Kusky TM (2010) Temporal evolution of the Angavo and related shear zones in Gondwana. Constraints from LAMC-ICP-MS-U–PB zircon ages of granitoids and gneiss from central Madagascar. Precambr Res 182:30–42. https://doi.org/10. 1016/j.precamres.2010.06.013 Ramsay JG, Huber MI (1987) The techniques of modern structural geology. Academic Press I & II:700 Ramsay JG (1967) Folding and Fracturing of Rocks. McGraw-Hill, New York, p 531 Ratschbacher L, Frisch W, Neubauer F, Schmid SM, Neugebauer J (1989) Extension in compressional orogenic belts: the Eastern Alps. Geology 17:404–407 Rice AHN, Greiling RO, Dardir AA, Rashwan AA, Sadek MF (1992) PanAfrican extensional structures in the area S of the Hafafit Antiform, Eastern Desert of Egypt. Zentralblatt fur Geologie und Palaeontolgie Teil I 11:2641–2651 Rice AHN, Osman AF, Sadek MF, Abdeen MM, Ragab AI (1993) Preliminary comparison of six late—to post-Pan-African molasse basins, E. Desert, Egypt. In: Thorweihe U, Schandelmeier H (eds) Geoscientific Research in Northeast Africa. Balkema, Rotterdam, pp 41–45 Ries AC, Shackleton RM, Graham RH, Fitches WR (1983) Pan African structures, ophiolites and mélange in the Eastern Desert of Egypt. A traverse at 26N. J Geolo Soc, London 140:75–95 Robinson FA, Bonin B, Pease V, Anderson JL (2017) A discussion on the tectonic implications of Ediacaran late—to post-orogenic A-type granite in the northeastern Arabian Shield, Saudi Arabia. Tectonics 36:582–600. https://doi.org/10.1002/2016TC004320 Royden LH, Burchfield BC (1989) Are systematic variations in thrust belt style related to plate boundary processes? (The western Alps versus the Carparthians). Tectonics 8:51–61

418 Royden LH, Burchfield BC, King RW, Wang E (1997) Surface deformation and lower crustal flow in Eastern Tibet. Science 276:788–790. https://doi.org/10.1126/science.276.5313.788 Patton TL, Moustafa AR, Nelson RA, Abdine SA (1994) Tectonic evolution and structural setting of the Suez Rift. In Interior Rift Basins. AAPG Memoir; American Association of Petroleum Geologists: Tulsa, OK, USA, vol 59, pp 7–55 Pegram WJ, Sando TW, Hodges KV (1980) Rare-earth element geochemistry and Nd-isotopic composition of some Egyptian Younger Granites. Geol Soc Am Abstr Programs 12:497 Plat JP, England PC (1993) Convective removal of lithosphere beneath mountain belts: thermal and mechanical consequences. American J Sciences 293:307–336 Saini HS, Nagpaul KK, Sharma KK (1978) Fission-track ages from the Pandoh-Baggi area, Himachal Himalaya, and their tectonic interpretations. Tectonophysics 46:87–98 Scharf A, Handy MR, Favaro S, Schmid SM, Bertrand A (2013) Modes of orogen-parallel stretching and extensional exhumation in response to microplate indentation and roll-back subduction (Tauern Window, Eastern Alps). Int J Earth Sci (Geol Rundsch). https://doi. org/10.1007/s00531-013-0894-4 Schmidt DL, Hadley DG, Stoeser DB (1979) Late Proterozoic crustal history of the Arabian shield, southern Najd province, Kingdom of Saudi Arabia. King Abdulaziz University, Jiddah. Institute of Applied Geology Bulletin 3(2):41–58 Schubert C (1980) Late Cenozoic pull-apart bains, Bocono, fault zone, Venezuelan Andes. J Struct Geol 2:463–468 Şengör AMC, Gorur N, Saroglu F (1985) Strike-slip faulting and related basin formation in zones of tectonic escape: Turkey as a case study. In: Biddle KT, Christie-Blick N (eds) Strike-slip deformation, basin formation and sedimentation. Soc Econ Mineral Paleont Spec Publ 37:227–264 Shackleton RM, Ries AC, Graham RH, Fitches WR (1980) Late Precambrian ophiolite melange in the Eastern Desert of Egypt. Nature 285:472–474 Shackleton RM (1986) Precambrian collision tectonics in Africa, Spec. Publ. Geol Soc London 19:329–349 Shackleton RM (1994) Review of Late Proterozoic sutures, ophiolite melanges and tectonics of eastern Egypt and NE Sudan. Geol Rundsch 83:537–546 Shalaby A, Stu¨we K, Makroum F, Fritz H (2005) A Pan-African crustal scale conjugate shear zone system in the Eastern Desert, Egypt: the Wadi Mubarak belt. Precam. Res. 136:27–50 Shalaby A, Stüwe K, Fritz H, Makroum F (2006) The El Mayah molasse basin in the Eastern Desert of Egypt. J Afr Earth Sc 45:1– 15 Shimron AE (1984) Evolution of the Kid Group, SE Sinai Peninsula: thrusts, mélange and implications for accretionary tectonics during Late Proterozoic of the Arabian-Nubian Shield. Geology 12:242– 247 Shimron AE (1987) Pan-African metamorphism and deformation in the Wadi Kid region, SE Sinai Peninsula: evidence from porphyroblasts in the Umm Zariq Formation. Israel J Earth Sci 36:173–193 Spear FS, Kohn MJ, Florence FP, Menard T (1991) A model for garnet and plagioclase growth in pelitic schists: implications for thermobarometry and P-T path determinations. J Metamorph Geol 8:683– 696 Stacey JS, Stoeser DB, Greenwood WR, Fischer LB (1984) U-Pb zircon geochronology and geologic evolution of the Halaban–Al Amar region of the eastern Arabian Shield, Kingdom of Saudi Arabia. Journal of the Geological Society, London 141:1043–1055 Stein M, Goldstein SL (1996) From plume head to continental lithosphere in the Arabian Nubian shield. Nature 382:773–778 Stern RJ (2017) Neoproterozoic formation and evolution of Eastern Desert continental crust—the importance of the infrastructure-

A. S. A. A. Abu Sharib superstructure transition. J Afr Earth Sci (in press). https://doi.org/ 10.1016/j.jafrearsci.2017.01.001 Stern RJ (1981) Petrogenesis and tectonic setting of Late Precambrian ensimatic volcanic rocks, Central Eastern Desert of Egypt. Precambr Res 16:195–230 Stern RJ (1985) The Najd fault system, Saudi Arabia and Egypt: a late Precambrian rift-related transform system? Tectonics 4:497–511 Stern RJ (1993) Tectonic evolution of the late proterozoic East African Orogen: constraints from crustal evolution in the Arabian-Nubian Shield and the Mozambique Belt. In: Thorweihe U, Schandelmeier H (eds) Geoscientific Research in Northeast Africa. Balkema, Rotterdam, pp 73–74 Stern RJ (1994) Arc assembly and continental collision in the Neoproterozoic East African Orogen: implications for the consolidation of Gondwanaland. Annual Reviews of Earth and Planetary Sciences 22:315–319 Stern RJ (2002) Crustal evolution in the East African Orogen: a neodymium isotopic perspective. J Afr Earth Sc 34:109–117 Stern RJ, Gottfried D (1986) Petrogenesis of a Late Precambrian (575– 600 Ma) bimodal suite in the N Eastern Desert of Egypt. Contrib Miner Petrol 92:492–501 Stern RJ, Gottfried D, Hedge CE (1984) Late Precambrian rifting and crustal evolution in the Northeastern Desert of Egypt. Geology 12:168–172 Stern RJ, Hedge CE (1985) Geochronologic and isotopic constraints on late Precambrian crustal evolution in the Eastern Desert of Egypt. Am J Sci 285:97–172 Stern RJ, Johnson PR (2010) Continental lithosphere of the Arabian Plate: a geologic, petrologic, and geophysical synthesis. Earth Sci Rev 101:29–67 Stern RJ, Manton WI (1987) Age of Feiran basement rocks, Sinai— implications for late Precambrian crustal evolution in the northern Arabian-Nubian Shield. J Geol Soc, London 144:569–575 Stoeser DB, Camp VE (1985) Pan-African microplate accretion of the Arabian shield. Geol Soc Am Bull 96:817–826 Stoeser DB (1986) Distribution and tectonic setting of plutonic rocks of the Arabian Shield. J Afri Earth Sci 4:21–46 Stoeser DB, Frost CD (2006) Nd, Pb, Sr, and O isotopic characteristics of Saudi Arabian shield terranes. Chem Geol 226:163–188 Sturchio NC, Sultan M, Batiza R (1983) Geology and origin of Meatiq Dome, Egypt: a Precambrian metamorphic core complex? Geology 11:72–76 Sturchio NC, Sultan M, Sylvester P, Batiza R, Hedge C, El Shazly EM, AbdelMaguid A (1983b) Geology, age, and origin of the Meatiq Dome—implications for the Precambrian stratigraphy and tectonic evolution of the Eastern Desert of Egypt. King Abdulaziz University, Faculty of Earth Sciences Bulletin 6:127–143 Sylvester AG (1988) Strike-slip faults. Bull Geol Soc Am 100:1666– 1703 Tapponnier P, Mattauer M, Proust FC (1981) Cassaigneau Mesozoic ophiolites, sutures, and large-scale tectonic movements in Afghanistan Earth and Planetary Science Letters 52:355–371 Tapponnier P, Peltzer G, Armijo R (1986) On the mechanism of the collision between India and Asia. In: Coward MP, Reis AC (eds) Collision tectonics. Geological Society of London Special Publication 19, 115–157 Tsige L, Abdelsalam MG (2005) Neoproterozoic-Early Paleozoic gravitational tectonic collapse in the southern part of the Arabian-Nubian Shield. The Bulbul Belt of southern Ethiopia. Precambr Res 138:297–318 Vanderhaeghe O, Teyssier C (2001) Crustal scale rheological transition during late-orogenic collapse. Tectonophysics 335:211–228 Wallace CA (1986) Lithofacies and Depositional Environment of the Maraghan Formation, and Speculation on the Origin of Gold in Ancient Mines, an Najadi area, Kingdom of Saudi Arabia. Saudi

16

Low Angle Normal-Sense Shear Zones, Folds and Wrench …

Arabian Deputy Ministry for Mineral Resources Open-File Report USGS-OF-06-6.19 p Wallbrecher E, Fritz H, Khudeir AA, Farahat F (1993) Kinematics of pan-African thrusting and extension in Egypt. Geosci Res Northeast Afr 27–30 Warden AJ, Horkel AD (1984) The geological evolution of the northeast branch of the Mozambique belt (Kenya, Somalia and Ethiopia). Mitt Osterr Geol Gesch 77:161–184 Warren RG, Ellis DJ (1996) Mantle underplating, granite tectonics, and metamorphic P-T-t paths. Geology 24(7):663–666 Williams GD, Powell CM, Cooper MA (1989) Geometry and kinematics of inversion tectonics. Geological Society, London,

419 Special Publications 1989; v. 44; p. 3–15. https://doi.org/10.1144/ gsl.sp.1989.044.01.02 Williams PL, Vaslet D, Johnson PR, Berthiaux A, LeStrat P, Fourniguet J (1986) Geologic map of the Jabal Habashi quadrangle, sheet 26F, Kingdom of SaudiArabia. Saudi Arab Dep Minist Miner. Resour Geosci Map GM-98 A Worku H, Yifa K (1992) The tectonic evolution of the Precambrian metamorphic rocks of the Adola belt, southern Ethiopia. J Afr Earth Sci 14:37–55 Ziegler PA (1989) Geodynamic model for Alpine intra-plate compressional deformation in western and central Europe. In: Cooper MA, Williams GD (eds) Inversion tectonics. Geol Soc Spec Publ 44:63–85

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Arc Accretion and Calc-Alkaline Plutonism Finalizing the Subduction Stage in the Arabian–Nubian Shield, with Emphasis on the Gneissic Core Complexes Adel A. Surour

Abstract

Role of arc(s) formation in the development of the Neoproterozoic Arabian–Nubian Shield (ANS) is very important, particularly from the geotectonic perspective. For almost two decades, several models have been suggested for the ANS evolution in which the theory of plate tectonics is accredited and the old geosynclinal theory is abandoned. In this respect, the models include (1) infrastructural model, (2) Turkic-type orogenic model, (3) hot-spot model and (4) arc accretion model. In this chapter, basics of arc assembly/arc accretion are presented aiming to understand the evolution of the ANS in a reasonable way. Ages of arc collisions are discussed with an emphasis on the mechanism and structure(s) of the arc collision stage. Owing to their intimate relationship with the formation of arc systems, their distribution, age, field characteristics (including rock associations), geochemistry and petrogenetic models are given. The chapter sheds light on the role of arc formation when the so-called core or gneissic complexes are developed, particularly in the Egyptian Eastern Desert (e.g. at Hafafit, Beitan, Khud’a, El-Shelul and Sibai). Such complexes are widely accepted in recent works to represent syn-arc collision massive plutonic injections along thrusts. This took place during the arc collisions opposing an old concept that considers them as pre-Pan-African windows to the old continent (i.e. infrastructural unit). A synopsis on the terranes including gneissic rocks is given, which deals with the metamorphic complexes in the Sinai Peninsula, in addition to some notes for the aerial distribution and

A. A. Surour Department of Geological Sciences, Faculty of Science, Galala University, 43511 New Galala City, Egypt A. A. Surour (&) Department of Geology, Faculty of Science, Cairo University, 12613 Giza, Egypt e-mail: [email protected]; [email protected]

age of accretions in the Neoproterozoic shield of Saudi Arabia. Keywords













Calc-alkaline plutonism Subduction stage Core complexes Petrogenetic models Infracrustal units

17.1

Introduction

Accretion is defined as a tectonic process in which the tectonic plate at or near a subduction zone gains materials on the edge of continent. These materials include a diverse range of sedimentary and igneous rocks at seamounts, oceanic crust and volcanic arcs (Sattler 1995; van der Pluijm 2004; Hawkesworth et al. 2019; Ferreira et al. 2020). An accretionary wedge, sometimes known as an accretionary prism, is developed when the sedimentary materials on the plate descend upon subduction followed by late accumulation near the continent edge (Fig. 17.1). According to van der Pluijm (2004) accumulated oceanic sediments, volcanic island arcs and seamounts are amalgamated on the upper plate at different times (Fig. 17.2). Based on a classification scheme by Cawood et al. (2009), accretionary orogens are distinguished into retreating orogens and advancing orogens, and they form at intra-oceanic and continental margins that characterize boundaries of convergent plates. For example, the modern western Pacific represents a retreating orogen that is characterized by a long-term extension concerning the subduction site. On the other hand, the Andes is a typical example of advancing orogens that develop by advanced plate overriding plate towards the down going plate. This results in foreland folding, thrust belts and crustal thickening (Charrier et al. 2013; Armijo et al. 2015). Age of accretionary structures can be as old as Archean, Proterozoic and Phanerozoic (e.g. Lee et al. 2007; Rollinson 2010; Stern 2018; Ferreira et al. 2020). Crystalline basement rocks in the

© The Author(s), under exclusive license to Springer Nature Switzerland AG 2021 Z. Hamimi et al. (eds.), The Geology of the Arabian-Nubian Shield, Regional Geology Reviews, https://doi.org/10.1007/978-3-030-72995-0_17

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422 Fig. 17.1 Block diagram showing position of accretionary prism with respect to volcanic arc and subduction zone. https:// commons.wikimedia.org/w/ index.php?curid=49035989

Fig. 17.2 Stages of plate accretion through time

A. A. Surour

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ANS formed by arc accretion in northeast Africa and in the Arabian Peninsula are mostly Neoproterozoic in age. According to Chaumillon and Mascle (1997), the Mediterranean Ridge (the African Plate subduction beneath the Eurasian Plate) is a typical example of accreted terranes with accretionary wedges in modern subduction plate boundaries. In the scope of the Turkic-type orogeny suggested by Sengor and Natal’in (1996), the roles of various subduction– accretion processes, in connection with continental margins development, are very important especially when it comes to crustal growth and its propagation with time. Also, Sengor and Natal’in (1996) believed that during ancient accretion, fore-arc accretion of the oceanic sedimentary slab was the responsible process in which the continental margins gained juvenile materials. However, in several active intra-oceanic and continental subductions, substantial prisms of accretion can be missing. This is owed either to subducted sediments paucity as a sediment-starved trench or to significant subduction erosion like the cases of Tonga, Mariana and Chile. (e.g. Chile, Tonga, Mariana). According to Collins (2002), accretionary orogens can be extensional and they represent magmatism of juvenile suprasubduction zone and back-arc basins inversion. Such a setting enables crustal material addition at some convergent margins. The archipelago orogenesis of Hsu (2003) belongs to the same setting. The accretion of arc-back-arc occupying a lower setting, with respect to plate, represents a characteristic mechanism to create a new continental crust that maintains the crustal loss balance at margins of convergence. Regardless the age and location all over the world, metamorphic complexes expose segments of high-pressure subduction-accretionary complexes that are built during island arc-continental margin convergence. They form a structural pile made up of metamorphosed (mostly serpentinized peridotites) and sedimentary rocks derived from accreted arcs and oceans so that ophiolite complexes are included in parts (Escuder-Viruete et al. 2013). According to Chekhovich (2006), island arcs can be ensiamatic and they are found in the lithosphere and upper crust levels. Subduction jamming prevails and the edifice of the island arcs becomes an attachment to the continent, if the accretion is lithospheric. During accretion at the crustal level, the lithospheric subduction underlying an island arc is developed further, and this helps a suprasubduction igneous belt to form at the continental margin. According to Cawood et al. (2009), cratonization takes with plate convergence so that a region of back- and magmatic arcs develops in zones where mechanical and thermal weakening dominate. The present chapter aims to review the types/mechanisms of arc accretion and the generation of calc-alkaline plutonism in the ANS. It includes the presentation of typical examples in the entire shield, in addition to a discussion about age and major tectonics with a special emphasis on the age and origin of gneissic dome cores, particularly in the Nubian Shield.

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17.2

Age of Arc Collisions

In the western Arabian Peninsula (Arabian Shield) and in northeast Africa (Nubian Shield), vast areas of Neoproterozoic crystalline rocks form the so-called Arabian–Nubian Shield (ANS) of Neoproterozoic age, with some limited exposures dated back to the Archean (Fig. 17.3). According to Whitehouse et al. (1998, 2001a); Abdelsalam et al. (2002); Johnson and Woldehaimanot (2003), some western and south-eastern parts of the ANS are occupied by pre-Neoproterozoic crust. Evidence from geochronology and isotopes suggesting crust dated back to the Archean and Paleoproterozoic in the eastern extremity of the Arabian Shield in the Kingdom of Saudi Arabia are given by Whitehouse et al. (2001a). The principle rock units in the ANS, from a lithostratigraphic point of view, comprise a variety of rock associations and groups that include ophiolites, ophiolitic mélange (dominated by low-grade metamorphosed volcano-sedimentary schistose rocks), complexes of metagabbro–diorite, arc-related calc-alkaline granitoids and equivalent volcanics and active continental margin unmetamorphosed sediments, fresh arc gabbros and post-collisional to within-plate alkaline granites and dyke rocks (El-Bialy 2020). According to this author, the alkaline granites represent evolved calc-alkaline rocks during a late stage of collision that form from a silicic magma characterized by emplacement after the phase of crustal thickening and before the phase characterized by onsetting of extension in the crust some 630–590 Ma ago. Around 610–580 Ma ago, the so-called Dokhan volcanic formed as a higk-K mafic to felsic rock series, and contemporaneous to the event of escape tectonics and extension of the crust during the collisional stage, A-type granites formed at 610–580 Ma ago. El-Bialy (op.cit.) gave a summary for the evolution of the ANS, both tectonically and magmatically, as follows: (1) the *870–670 Ma period that witnessed the development of volcano-sedimentary pile and intrusive rocks as this period was characterized by accreted terrains amalgamation at the continental block of Gondwana particularly at its eastern part; (2) immense thickening of the crust at *650– 640 Ma contemporaneous to the collision between the newly formed or juvenile accreted ANS crust with the pre-Neoproterozoic continental margin of West Gondwana or the so-called Saharan Metacraton, and this occurred in the form of two distinct sutures, namely arc-arc and arc-continental; (3) the post- 630–550 Ma) post-collisional stage that witnessed extensional collapse (extension and thinning) of the ANS juvenile crust and this lasted until *550 Ma. The previously mentioned stages of tectonics and magmatism comprise arc formations, amalgamation, stabilization at the post-collisional sage and finally extensive erosion and the development of voluminous peniplains when

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Fig. 17.3 Neoproterozoic terranes map of the Arabian–Nubian Shield (Johnson et al. 2011). NED = North-Eastern Desert, CED = Central Eastern Desert, SED = South-Eastern Desert (Stern and Hedge 1985)

the early Paleozoic began to form as stated by Bentor (1985). Generally, and based on data cited in Bentor (1985); Stern (1994); Stein and Goldstein (1996); Genna et al. (2002); Johnson and Woldehaimanot (2003); Stoeser and Frost (2006); Doebrich et al. (2007); Cox et al. (2012), the ANS represents a set of volcanic arcs mostly juvenile, that associate remnants of ophiolite formed by amalgamation as the Gondwana is assembled. Ediacaran terrane accretion in the

ANS was discussed in detail by Cox et al. (2012). Eyal et al. (2014) believed in a Proterozoic island arcs model for the Arabian–Nubian Shield, particularly its northern extremity, based on zircon isotopic U/Pb data. The ANS formed when the Rodina broke-up of Rodinia and initial rifting dominated at 870–800 Ma contemporaneous to the Mozambique Ocean’s closure and as the east and west Gondwana fragments collided at 650–600 Ma ago

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as stated by Jacobs and Thomas (2004). Vail (1985); Stoeser and Camp (1985) were the first to define a well-established model of arc assembly in terms of arc accretion for the ANS, and such a model was then adopted by Stern (1994). In the suggested model, the orogeny in East Africa resulted in the juvenile crust around and within the Mozambique Ocean. According to Stern (1994); Stern and Johnson (2010), the formation of ANS juvenile crust took place in the Mid-Cryogenian in the time span from 800 to 690. Generally, arc-related rocks in Saudi Arabia, Yemen, Jordan, Egypt, and Sudan are dominated by fore-arc ophiolites formed in suprasubduction zones (SSZ), metavolcanics and calc-alkaline plutonism, in which the latter resulted in considerable amounts of tonalite and granodiorite. In Egypt, and based on their stratigraphic position and composition, the metavolcanics are distinguished into OMV (standing for the older metavolcanics) that mostly include spilitized pillowed metabasalt of the ophiolite representing fragmented crust of the ancient ocean, and YMV (standing for the younger metavolcanics) that are arc-related but distinctly non-ophiolitic with spatial distribution in the central and southern Egyptian Eastern Desert, CED & SED (Stern 1981, Figs. 17.4 and 17.5a) as well as its northern part, NED (Bühler et al. 2014). OMV and YMV show occasional time overlapping, some of the OMV basalts in few cases appear younger than YMV rock varieties as documented by some authors, e.g. Ali et al. (2009); Andresen et al. (2009). For this reason, and on a stratigraphic basis, Bühler et al. (2014) recommended a new subdivision scheme for the metavolcanics in the Egyptian Eastern Desert considering precise and confirmed geochronological and stratigraphic criteria. The latter was established based on the Wadi Malaak succession (WMS) non-ophiolitic volcano-sedimentary pile metamorphosed in the greenschist facies conditions and crop out between Gebel El Kharazah and Gebel Kifri in the NED. The LA-ICP-MS U-Pb zircon age combined with 40Ar/39Ar geochronological data of biotite and hornblende data given by Bühler et al. (2014) documented for the first time ca. 720 Ma rocks of arc volcanism in the ANS. In such a configuration, the ca. 750 Ma granitoids is underlain by the Wadi Malaak volcano-sedimentary succession (WMS) in a nonconformable relationship overlies, but the WMS itself is overlain by ca. 617 Ma along an angular unconformity with the Dokhan Volcanics and the molasse-type sediments of the Hammamat Group. Other examples of precise age-dating documentation of magmatic rocks in the Mid-Cryogenian of the northern ANS are represented by arc metadacite as a part of the YMV at the Allaqi-Heini suture in the SED that overlies gabbro of the ophiolitic assemblage formed at 730 ± 6 Ma, and according to Ali et al. (2010) it yields a 733 ± 7 Ma SHRIMP U-Pb zircon age. In the CED, plagiogranite at the Ghadir area gave a 746 ± 19 Ma Pb–Pb zircon evaporation age by Kröner et al. (1992), and similarly

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Anderson et al. (2009) assigned a 736.5 ± 1.2 Ma TIMS U– Pb zircon age for the ophiolitic gabbro at the Fawakhir-Bir El Sid area. Therefore, it can be concluded that the island arc formation around 870 Ma is the earliest stage of the ANS tectonic evolution and formation of ophiolite as suggested by Pallister et al. (1988). In addition, the latter authors discussed the details the ANS accretion in the crust, particularly in Saudi Arabia, based on combined field, U-Pb ages, and Pb-isotope systematic.

17.3

Syn-Collisional (Arc-Related) Older Granitoids

According to Stern (1979), the older granitoids (GI) constitute about 27% of the basement outcrop in the Egyptian Eastern Desert, which are distinguished by Hussein et al. (1982) into GI, GII and GII (Fig. 17.5b). According to Hussein et al. (1982); Stern and Hedge (1985), subduction at 850–610 Ma resulted in the emplacement of I-type granitoids of calc-alkaline composition. The arc-related granitoids are termed as the Older Granites (OG) that were formed at three events as recognized by Hassan and Hashad (1990). The first event is known as the 850–800 Ma Shaitian event, the second is the 760–710 Ma Hafafit event and the Meatiq event at 630–610 Ma represent the third event. According to Loizenbauer et al. (2001); Andresen et al. (2009); Eliwa et al. (2014); El-Shazly and Khalil (2016), age of older granitoids never exceeds 770 Ma based on U–Pb age of zircon. These rocks show distinct variability in the mineralogical composition being represented basically by quartz diorite, granodiorites and tonalities, in addition to fewer monzogranites (El-Bialy 2004). A sharp contact between the older granitoids and the country rocks is sharp and non-reactive. Offshoots and apophyses are common and it is worth to mention here that the contacts are somehow straight, angular and locally crenulated. The granitoids are characterized by chilled margins. According to El-Metwally (1993); Surour and Kabesh (1998); El-Bialy (2004); Asran and Abdel Rahman (2012), the older granitoids host considerable amount of mafic microgranular enclaves. It is not difficult to identify and to distinguish the mafic enclaves and fine-grained xenoliths of mafic rocks, for example, fine-grained biotite schists. Also, El Bialy (2004, 2020) identified rounded shape bodies that lack evidence of thermal metamorphism. Geochemistry of older granitoids (GI) in Egypt was interpreted and summarized by El-Bialy (2020) based on data from literature including his own work, e.g. El-Bialy (2004); Farahat et al. (2007); Bea et al. (2009); El Mahallawi and Ahmed (2012); El-Bialy and Omar (2015). Generally, the Egyptian GI granitoids are medium-K calc-alkaline, metaluminous to slightly peraluminous and magnesian

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Fig. 17.4 Generalized geological map of the Eastern Desert and the Sinai Peninsula showing a complete stratigraphic sequence of Neoproterozoic rock as Tonian-Cryogenian and Ediacaran associations. Major structural elements are shown. The proposed line separating the Keraf suture and the Sahara metacraton is shown. Compilation of the figure is based on data from Stern (1981, 1994), Johnson and Katan (2008), Stern and Johnson (2010), and Johnson et al. (2011)

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Fig. 17.5 a Distribution of metavolcanic rocks in Egypt, mostly YMV and less abundant OMV with the ophiolites (undifferentiated) from El-Bialy (2020) modified after El-Ramly (1972). b Distribution of arc-related and post-collisional granitic rocks in Egypt. Modified by El-Bialy (2020) after Farahat et al. (2011)

(FeO*/FeO* + MgO < 0.8) when it comes to magma type. They are assigned as I-type granites having a broad compositional spectrum ranging from highly silicic to basic (75– 50 wt% SiO2), low K2O and K2O/Na2O (less than 2.5 wt% and 0.7, respectively). Their REE patterns are characterized by both delicate negative and positive anomalies of Eu. In addition, they show relative enrichment in the LILE Ba and Sr, whereas Ga and HFSE such as Nb, Zr and Y are remarkably depleted, and the total REE content is lesser than 100 ppm. From the petrological point of view, two contrasting models of petrogenesis were suggested for the origin of the older granitoids that form in a typical arc setting as follows: (1) high fractional crystallization of lithospheric mantle-derived mafic magmas, with crustal modification through either assimilation or melt hybridization (e.g. El-Bialy 2004; Eliwa et al. 2014; El-Bialy and Omar 2015), (2) fractional crystallization of mafic crustal melts generated by high partial melting degrees of mafic lower crustal

amphibolite protolith (Farahat et al. 2007; El Mahallawi and Ahmed 2012). Structurally, Hamimi and Abdel-Wahed (2020) stated that arc accretion in the ANS result in different types of shear zones that can be distinguished into: syn-accretion- and post-accretion shear zones. The first and second groups are represented by NNE-oriented Hamisana Shear Zone (HSZ) and the NW-trending, most probably Najd-related, shear zones, respectively. The latter group includes some famous shear zones like those in the Nugrus, Hodein-Wadi Kharit and Attala areas, in addition to relatively younger ENE- and E-trending Barramiya-Wadi Mubarak Shear Belt and the Abu Dabbab Shear Zone. Hamimi and Abdel-Wahed (2020) stated that it is widely accepted that there are significant effects of Najd Fault System (NFS) on the ANS and the CED of Egypt. Stern (1985) considered the NFS as the largest Proterozoic Shear System on Earth, representing the youngest major structural element in the Egyptian Eastern

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Desert. The NFS has a great importance due to its major extension, role in the exhumation of metamorphic core complexes and prominence in Gondwana cratonization. In vast areas within the Arabian Shield, primary and secondary faults in the NFS were studied in detailes by Moore (1979) who defined the NFS as a major Proterozoic Fault System, more specifically as strike-slips so it can be considered as a major transcurrent system. According to Fritz et al. (2002); Abd El-Wahed (2014), the major shear zones bounding the gneiss domes are considered as sinistral strike-slip shear zones combined with domes exhumation in the sense of external shears. On the other hand, Andresen et al. (2009) believed that they represent remnants of thrusts in the northwest direction. Nowadays, in the NFS, the NW-trending shear zones of a sinistral sense are genetically linked with deposition of sediments, exhumation of gneiss domains, and emplacement of syn-tectonic granitoids (e.g. Fritz et al. 2002, 2013; Abdeen et al. 2014; Abd El-Wahed et al. 2016; Makroum 2017). In addition, Hamimi et al. (2019) believed in a significant role of shear zones within the NFS in the evolution and exhumation of gneissic core complexes in the Egyptian Eastern Desert, and they believed that the change of shear kinematics within the NFS from transpressive to transtension through time and emplacement of transpressive- and transtension-related granitoids (655 and 645 Ma) is very characteristic. The latter is exclusively post-collisional to typical within-plate and represented by the Younger Granites (GII and GIII) with distinct geochemical signature pertaining alkaline magma in most but some calc-alkaline and peralkaline pulses are also known. Among them, peraluminous granites show evidence of derivation from a pelitic or sedimentary source, i.e. S-type magma, with a very characteristic crystallization of aluminous minerals such as garnet and cordierite.

17.4

Gneissic Domes: Continental Windows (Infrastructure) versus Arc-Related Plutonism and Metamorphism

Presence of pre-Pan-African continental windows is a matter of debate for decades. In the Egyptian Eastern Desert as well as in the Midyan and other terranes of western Saudi Arabia, Mesoproterozoic or older rocks are not present but they are encountered at the Khida terrane in the south (Whitehouse et al. 2001a). El-Gaby et al. (1984) and some other workers (e.g. Hamimi et al. 1994; Hamimi 1996; Khudier et al. 2008) believed that the continental windows (tier 1) represent the infrastructure. It represents the lowest unit structural unit beneath the massive and thick Pan-African suprastructural rocks as exhumed pre-Pan-African crust. Such windows are dominated by granitoid orthogneisses, locally migmatized, and include metabasite, meta-ultramafic rocks and

meta-sediments. The gneisses of tier 1 appear as domes or “basement windows” (e.g. Meatiq, Hafafit, El Shalul gneiss complexes; Fig. 17.6a, b) surrounded by low-grade rocks belonging to tier 2. Generally, the windows are bounded by arc-related shear zones like the case of Nugrus thrust for the Hafafit area (Fig. 17.7). Lundmark et al. (2012) presented precise U-Pb ages of zircon, titanite and polycrase from some plutons in a complicated ortho-gneiss complex at the Hafafit area suggesting a Pan-African age for metamorphism and calc-alkaline magmatism. Similar conclusion was reached for similar gneisses at the Meatiq and El Sibai areas (Andresen et al. 2009; Augland et al. 2012) based on some ID-TIMS data. Lundmark et al. (2012) believed that the subdivisions of granitic rocks in Egypt into “Older” and “Younger”, on field basis only, are not enough. They (op.cit.) proposed that the evolutionary model should include the whole Eastern Desert, both tectonically and magmatically. Selected published age data given by Lundmark et al. (2012) provide robust ages for granitoids from some areas such as El Shelul, Meatiq and El Sibai (Fig. 17.8). They (op.cit.) added that newly presented ages represent magmatism and/or metamorphism as six pulses at (1) 705–680, (2) 660, (3) 635–630, (4) 610–604, (5) 600–590 and (6) 540 Ma. Mostly, the ages define pulses of granitic rocks in the eugeoclinal allochton and its underlying domes and gneissic windows. They represent two tiers characterizing two crustal levels. It is proposed that syn-orogenesis represents pulses 1–3, whereas mid-crustal gneisses exhumation along shear zones in the Eastern Desert represent pulses 4 and 5. The 6th or final pulse overprints the East-African orogeny.

17.5

Arc-Related Gneissic Complexes and Metamorphic History in the Sinai Peninsula, with an Emphasis on the Wadi Kid Environ

In the Sinai Peninsula, metamorphosed volcanic and sedimentary successions crop out in the Wadi Kid, Feiran-Solaf, Taba and Wadi Sa’al areas (Fig. 17.9). The latter figure shows that the metamorphic complexes in Sinai (four, isolated) are confined to post-collisional granitic rocks that occupy the majority of the Sinai Peninsula (up to 70%). This is also applied to the Elat areas well as the northernmost tip of the ANS (Eyal et al. 2014) that are located within post-collisional granitoids that occupy up to 70% of the total area in Sinai (Egypt) and Elat area (southern Israel) and the northernmost part of the Arabian–Nubian Shield. The metamorphic rocks include metasediments, metavolcanic rocks (both mafic and felsic) that are intruded by plutons of granite, diorite and gabbros, and all are subjected to deformation, particularly penetrative. Eyal et al. (2014) suggested

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Fig. 17.6 a Simplified geological map of the Egyptian Eastern Desert showing locations of gneissic domes and possibly migmatized terrains of arc origin. Modified by Lundmark et al. (2012) after Moussa et al. (2008). b Location of regional-scale extensional shear zones in the CED and northern extreme of the SED compiled by Fowler and Hamimi (2020) from Greiling et al. (2014) and Fowler and Osman (2009, 2013)

at least three island arcs during the Meso- to Neoproterozoic times and they needed about 500 My to develop, on the basis of SIMS age data of zircon and titania, from these metamorphic complexes. The Sa’al arc (SMC) needed *100 My to develop (1.03–0.93 Ga); the Feiran-Solaf-Taba arcs (FSMC and TMC) needed *130 My to develop (870– 740 Ma), and only *20 My (640–620 Ma) were enough to develop the Kid arc. Evidence for an older, ca. 1.1 Ga, pre-Sa’al island arc relies on some radiometric data of xenocrystic zircon. The major basement units in the area of Wadi Kid (Fig. 17.10), according to Shimron (1980); El-Metwally et al. (1999); El-Baily (2010, 2013), are distinguished from the oldest to youngest as follows: gneisses, metasediments, metavolcanics (basalt and andesite), metagabbro-diorite complex, older granitoids (tonalite and quartz diorite), Dokhan volcanics (rhyolite and dacite) and younger granites (monzo- and syenogranite). Particularly at the Wadi Melhag,

the Dokhan volcanic is common whereas a peculiar pluton at Wadi-Um Zeriq junction represents the younger granites. Gneissic rocks compose the major metamorphic components of the KMC consisting of foliated and lineated diorite and tonalite in which mafic xenoliths and enclaves are common, and in addition augen structure can be seen in some outcrops. These gneisses at the Wadi Qenaia were assigned by Bentor and Eyal (1987) as the Qenaia Formation that is believed to be of ortho-origin (Brooijmans et al. 2003; Be’eri-Shlevin et al. 2009). Metasediments are widely distributed in south-eastern Sinai, and they are distinguished by El-Metwally et al. (1999) metamorphosed calc-pelites and psammo-pelites. Metamorphosed calc-pelitic rocks include para-amphibolite, in addition to chlorites nd hornblende schists. In most, phyllites and garnet-biotite schists represents the metamorphosed psammo-pelitic rocks. These metasediments constitute the Umm Zariq Formation (Furnes et al. 1985), Heib Formation and most of Malhak Formation

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Fig. 17.7 Simplified structural geology map of the Hafafit gneissic dome and associated plutonic rocks, mostly arc-related calc-alkaline plutonic rocks (Lundmark et al. 2012)

(Shimron 1980). There is a dispute about the total thickness of the four formations that is believed to amount up to 12,000 m by Shimron (1987), whereas Blasband et al. (2000) stated that the thickness does not exceed 1500 m. The belt of Wadi Kid shows a complex metamorphic history that took place during Neoproterozoic time (Brooijmans et al. 2003). The Wadi Kid metamorphic belt comprise a thick pile of volcano-sedimentary rocks that can be distinguished into slightly metamorphosed volcanics (basalt, andesite, high K-dacite, rhyolite and pyroclastics), whereas siliceous mudstone, siltstone, greywacke and conglomerate represent the metamorphosed shallow-marine sedimentary rocks as given by Furnes et al. (1985). The metamorphic and structural history of this belt can be related to the tectonic framework of Sinai Peninsula and consequently to the Neoproterozoic Pan-African Orogeny. Blasband et al. (1997, 2000); Brooijmans et al. 2003) believed that the lithospheric thermal upwelling resulted in the development of the Kid core complex in sense of orogenic collapse. These suggestions are based mostly on the presence of high-temperature metamorphic assemblages and the minimum melting temperature estimations for the granitic suite (Blasband et al. 2000). Shallaly (2006) and El-Bayoumi et al. (2006) described in details the counter-clockwise pressure–temperature–time path of metavolcanics and gneissose diorite. They (op.cit.) related their geothermometric history to some events of deformation. According to Gad and Kusky (2007), anticlinal and synclinal structures trending in the ENE–WSW direction resulted in a fold series in the volcano-sedimentary pile at the area of Wadi Kid. There, grade of metamorphism

increases in the upstream (amphibolite facies), whereas the greenschist facies characterize the downstream. They (op.cit) summarized that the gneiss form elongate asymmetrical NEto NW-plunging folds in the south-western part of Wadi Kid area. They juxtapose the metasediments along a NNW- to NW striking sharp-faulted contact. The core complex at the area of Wadi Kid has several analogues in the ANS, and they are considered all together as extensional ones. Abundant extensional phase(s) at the closure of the Pan-African time might help to understand the ANS tectonics properly. The Neoproterozoic of the ANS has here tectonic phases, namely an oceanic phase represented by ophiolites and remnants of island arc, arc-accretion and the formation of terranes delineated by sutures. This resulted in thickening of the lithosphere. The NW-SE extension took place when the Neoproterozoic came to its end. During this phase, the gneissic core complexes started to develop in addition to the development of late-orogenic basins (extensional) and large zones of strike-slips (Blasband et al. 1997, 2000; Brooijmans et al. 2003). Shallaly (2006) and Abdel Wahed and Shallaly (2007) defined the tectonic model of the area at the Wadi Kid environ in terms of collision and subduction reversal of arc-continent. Finally, Fowler et al. (2010) concluded that the subduction was to the N or NNW based on the arc trend (ENE–WSW to E–W) and style of the arc-related basins. They (op.cit.) believed that the northern part of the Wadi Kid area experienced a four-stage deformation history (D1 to D4). The three deformation events, D1 to D3, in the southern part of the Kid area involve NW–SE to NNE–SSW compressional tectonism. The most acceptable tectonic models argue for an event of arc collision (D1)

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Fig. 17.8 Compiled U–Pb ages from the Hafafit/Marsa Alam, Meatiq, El Sibai and El Shalul areas (a) and probability density plot shows six magmatic/metamorphic peaks (b) from Lundmark et al. (2012)

that thickened the crust. D2 in these models is represented by core complex exhumation as an extensional tectonic event. The HT-LP metamorphism is older than younger granite intrusions that surround the Wadi Kid area. The D2 event was characterized by compression that produced folds and thrusts with top-to-the-SSE transport direction. Garnet and

staurolite porphyroblast-matrix foliation textures indicate high-T growth of porphyroblast as pre- or early-S2 that postdates F2 folding of beds, and therefore these porphyroblasts develop during the D2 event. F2 folding resulted in the increase of pressure during the high-T metamorphism prograde stage.

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Fig. 17.9 Four metamorphic “core” complexes in the Sinai Peninsula. TMC = Taba (Elat in the original script) Metamorphic Complex, KMC = Kid Metamorphic Complex, SMC = Sa’al Metamorphic Complex and FSMC = Feiran-Solaf Metamorphic Complex. Note that Q and UZ denote the Qenaia migmatized gneisses and gneisses of the Umm Zariq Formation within the KMC (from Eyal et al. 2014)

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Fig. 17.10 Geological map of the Wadi Kid area, south-eastern Sinai (from El-Bialy 2013 modified after El-Metwally et al. 1999; El-Bialy 2010)

17.6

Terrane Accretion in the Arabian Shield

Cox et al. (2019) considered the ANS as an ideal example to study the tectonics of island-arc accretion and continental crustal growth owing to the ANS juvenile nature in addition to its low metamorphic grades. The ANS (Fig. 17.1) is the north-east extension of the EAO based on the data cited in Cox et al. (2019), and it is composed of juvenile island arc terranes (e.g. Stern and Johnson 2010; Johnson et al. 2011) with Cryogenian to early Ediacaran accretion (Johnson and Kattan 2008; Johnson et al. 2011) with, at least one, pre-Neoproterozoic crustal block of the so-called Khida subterrane (Whitehouse et al. 2001a). The ANS becomes younger towards its easternmost extremity at the Ar Rayn Terrane (Johnson et al. 2011). According to Johnson and Stewart (1995), more terranes beneath the sedimentary cover in the eastern part of the shield rocks in Saudi Arabia have been identified by magnetic methods. Also, relatively older terranes (840–825 MY) as isolated inliers are reported in Oman by Whitehouse et al. (2016); Alessio et al. (2017). In the ANS, the youngest magmatism associates active subduction, e.g. the Ar Rayn Terrane at the far east of the shield where magma of calc-alkaline composition lasted until *600 Ma (Doebrich et al. 2007). Also in the eastern part of the shield, the Ad Dawadimi Terrane is located (Fig. 17.11), and it is sandwiched between the western Afif Terrane from the Ar Rayn Terrane (Fig. 17.11). An oceanic arc affinity for the latter two was proposed by Stoeser and Frost (2006) on some Pb isotope data presented by (Stacey et al. 1984) in some intrusions belonging to the Ediacaran and Paleozoic. It is believed that the most recent amalgamation of terranes in

Saudi Arabia took place in the eastern extremity of the shield where younger sutures are concealed by the Phanerozoic sediments (Cox et al. 2012). Nevertheless, it is believed that the Ediacaran basins in the Eastern ANS are not identical or analogue to the Omani Huqf basin. Instead, it was a part of the Indian Neoproterozoic passive margin, which is separated by the Mozambique Ocean from the African active margin. Most probably closure did not take place until the late Ediacaran or even the early Cambrian. In general, Neoproterozoic crust formed from different oceanic island-arcs is a common juvenile crust in the entire ANS. This took place as the Mozambique Ocean closed and the Gondwana amalgamated as documented in Johnson and Woldhaimanot (2003); Collins and Pisarvesky (2005); Stern (1994, 2008); Ali et al. (2010); Stern and Johnson (2010). However, in southern Saudi Arabia as well as in Yemen (the Al-Mahfid Terrane), crust older than the Neoproterozoic exists in the Khida Terrane (Stacey et al. 1984; Stoeser et al. 2001; Windley et al. 1996; Whitehouse et al. 2001b). According to Delfour (1979), pelitic schists and meta-greywackes are the principle lithologies of the Abt Formation in the Ad Dawadimi Terrane. Here, it is worth mentioning that no lower and upper contacts of the Abt Formation can be traced or easily identified other than structural contacts to the east and west along major terrane bounding faults. Detailed mapping, lithologically and structurally, along *10 km and 46 km E-W transects indicates the presence of younger members of the formation at the west with schistosity dipping due west, and it consists of a muscovite-biotite-quartz-feldspar schists of greenschist facies metamorphic condition and compositionally immature nature of the protolith sediments. In that terrane, ophiolites

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Fig. 17.11 Terrane map of the Arabian Shield showing major lithostratigraphic units and major structures related to accretion. HZFZ = Halaban-Zargat Fault Zone; AAFZ = Al Amar Fault Zone and ARSZr = Ar Rika Shear Zone. Modified by Cox et al. (2019) after Nettle et al. (2014)

and zones of ophiolitic mélange(s) (Al-Shanti and Gass 1983; Al-Saleh and Boyle 2001; Cox 2011) and a wide variety of granitic rocks (syn- and post-tectonic as mentioned by Stacey et al. 1984), all are present side by side with the Abt Formation. Nawab (1979); Al-Saleh and Boyle (2001); Doebrich et al. (2007) presented two models for the development of the Ad Dawadmi Terrane. The first is a back-arc basin situated between the Ar Rayn and Afif Terranes, or a fore-arc basin at the Ar Rayn Terrane. In the latter, Stoeser and Camp (1985) considered it a microcontinent with an outboard character. The Ad Dawadmi Terrane has similarities with more modern accretionary environment (Stoeser and Frost 2006) owing to the common association of the Abt Formation (immature sediments) and ophiolite mélange zones with forearc affinity (Cox 2011).

17.7

Conclusions

Based on the literature review and data materialized in this chapter, the following conclusions can be fitted:

1. From the geotectonic point of view, the role of arc(s) formation in the Neoproterozoic evolutionary history of the ANS is very important. 2. The old geosynclinal theory is abandoned, and instead the plate tectonics theory enables to formulate several models for the ANS evolution. 3. The models based on the aspects of plate tectonics include (a) infrastructural model, (b) Turkic-type orogenic model, (c) hot-spot model and (d) arc accretion model. 4. Formation mechanism of arc assembly/arc accretion helps to understand the evolution of the ANS, in which ages of arc collisions are dependent on the mechanism and structure(s) of the arc collision stage. 5. The suggested models are based on the intimate relationship with the formation of arc systems, their distribution, age, field characteristics (including rock associations), geochemistry and petrogenetic models. 6. The role of arc formation to develop the so-called core or gneissic complexes, particularly in the Egyptian Eastern Desert (e.g. at Hafafit, Beitan, Khud’a, El-Shelul and Sibai) is highly considered. 7. Most probably the core complexes represent syn-arc collision massive plutonic injections along thrusts, and this took place during the arc collisions opposing an old concept that considers them as pre-Pan-African windows to the old continent (i.e. infrastructural unit). 8. In the eastern extremity of the shield in Saudi Arabia, some terranes witnessed the oldest arc accretion in the entire ANS. In some specific cases in Saudi Arabia and Yemen, windows to an old continent could be considered but ambiguity about age and the geotectonic history still needs a solution.

References Abdeen MM, Greiling RO, Sadek MF, Hamad SS (2014) Magnetic fabrics and Pan-African structural evolution in the Najd Fault corridor in the Eastern Desert of Egypt. J Afr Earth Sci 99:93–108 Abdelsalam MG, Liégeois LP, Stern RJ (2002) The Saharan metacraton. J Afr Earth Sci 34:119–136 Abd El-Wahed MA (2014) Oppositely dipping thrusts and transpressional imbricate zone in the Central Eastern Desert of Egypt. J Afr Earth Sci 100:42–59 Abd El-Wahed MA, Harraz HZ, El-Behairy MH (2016) Transpressional imbricate thrust zones controlling gold mineralization in the Central Eastern Desert of Egypt. Ore Geol Rev 78:424–446 Abdel WM, Shallaly NAM (2007) Arc-continent collision and subduction reversal in the Central Wadi Kid Precambrian rock complex, Southeast Sinai, Egypt. Annals Geol Surv Egypt 29:37–50 Alessio B, Blades M, Murray G, Thorpe B, Collins AS, Kelsey D, Foden JD, Payne S, Al-Khirbash S, Jourdan F (2017) Origin and

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tectonic evolution of the NE basement of Oman: a window into the Neoproterozoic accretionary growth of India? Geol Mag 155:1–25 Ali KA, Azer MK, Gahlan HA, Wilde SA, Samuel MD, Stern RJ (2010) Age constraints on the formation and emplacement of Neoproterozoic ophiolites along the Allaqi-Heiani Suture, South Eastern Desert of Egypt. Gond Res 18:583–595 Ali KA, Stern RJ, Manton WI, Kimura J-I, Khamees HA (2009) Geochemistry, Nd isotopes, and U-Pb SHRIMP zircon dating of Neoproterozoic volcanic rocks from the Central Eastern Desert of Egypt: new Insights into the *750 Ma crust-forming event. Prec Res 171:1–22 Al-Saleh AM, Boyle AP (2001) Structural rejuvenation of the eastern Arabian Shield during continental collision: 40Ar/39Ar evidence from the Ar Ridayniyah ophiolitic mélange. J Afr Earth Sci 33:135–141 Al-Shanti AM, Gass IG (1983) The Upper Proterozoic ophiolite melange zones of the easternmost Arabian shield. J Geol Soc 140:867–876 Andresen A, El-Rus MAA, Myhre PI, Boghdady GY, Corfu F (2009) U-Pb TIMS age constraints on the evolution of the Neoproterozoic Meatiq Gneiss Dome, Eastern Desert, Egypt. Inter J Earth Sci 98:481–491 Armijo R, Lacassin R, Coudurier-Curveur A, Carrizo D (2015) Coupled tectonic evolution of Andean orogeny and global climate. Earth-Sci Rev 143:1–35 Asran AM, Abdel Rahman EM (2012) The Pan-African calc-alkaline granitoids and the associated mafic microgranular enclaves (MME) around Wadi Abu Zawal area, North Eastern Desert, Egypt: geology, geochemistry and petrogenesis. J Biol Earth Sci 2 (E1–E1):6 Augland LE, Andresen A, Boghdady GY (2012) U-Pb ID-TIMS dating of igneous and metaigneous rocks from the El-Sibai area: time constraints on the tectonic evolution of the Central Eastern Desert, Egypt. Inter J Earth Sci 101:25–37 Bea F, Abu-Anbar M, Montero P, Peres P, Talavera C (2009) The 844 Ma Moneiga quartz-diorites of the Sinai, Egypt: evidence for Andean-type arc or riftrelated magmatism in the Arabian-Nubian Shield? Prec Res 175:161–168 Be’eri-Shlevin Y, Katzir Y, Valley JW (2009) Crustal evolution and recycling in a juvenile continent: oxygen isotope ratio of zircon in the northern Arabian Nubian Shield. Lithos 107:169–184 Bentor YK (1985) The crustal evolution of the Arabian-Nubian Massif with special reference to the Sinai Peninsula. Prec Res 28:1–74 Bentor YK, Eyal M (1987) The geology of Southern Sinai, its implication for the evolution of the Arabian-Nubian Massif Volume 1: Jebel Sabbagh Sheet. Israel Acad Sci Human Bull, Jerusalem, 484 p Blasband B, Brooijmans P, Dirks P, Visser W, White S (1997) A Pan-African core complex in the Sinai, Egypt. Geologie en Mijnbouw 73:247–266 Blasband B, White S, Brooijmans P, De Boorder H, Visser W (2000) Neoproterozoic extensional collapse in the Arabian-Nubian Shield. J Geol Soc London 157:615–628 Brooijmans P, Blasband B, White SH, Visser WJ, Dirks P (2003) Geothermobarometric evidence for a metamorphic core complex in Sinai, Egypt. Prec Res 123:249–268 Bühler B, Breitkreuz C, Pfänder JA, Hofmann M, Becker S, Linnemann U, Eliwa HA (2014) New insights into the accretion of the Arabian-Nubian Shield: Depositional setting, composition and geochronology of a Mid-Cryogenian arc succession (North Eastern Desert, Egypt). Prec Res 243:149–167 Cawood PA, Kröner A, Collins WJ, Kusky T, Mooney WD, Windley BF (2009) Accretionary orogens through Earth history. In: Cawood PA, Kröner A (eds) Earth Accretionary Systems in Space and Time. Geol Soc London Spec Publ 318:1–36

435 Chaumillon E, Mascle J (1997) From foreland to forearc domains: New multichannel seismic reflection survey of the Mediterranean ridge accretionary complex (Eastern Mediterranean). Mar Geol 138(3– 4):237–259 Charrier R, Herail G, Pinto L, Garcia M, Riquelme R, Farias M, Munoz N (2013) Cenozoic tectonic evolution in the Central Andes in northern Chile and west central Bolivia: implications for paleogeographic, magmatic and mountain building evolution. Intern J Earth Sci 102(1):235–264 Chekhovich VD (2006) Different levels of ensimatic island-arc accretion. Geotectonics 40:101–110 Collins AS, Pisarevsky SA (2005) Amalgamating eastern Gondwana: the evolution of the Circum-Indian Orogens. Earth Sci Rev 71:229– 270 Collins WJ (2002) Nature of extensional accretionary orogens. Tectonics 21(4):12 Cox JM (2011) Jebel Tays ophiolitic thrust sheet, Saudi Arabia: tectonic setting, petrogenesis and geochronological framework. Saudi Geol Surv Tech Rep SGS-TR-2010-6 Cox GM, Foden J, Collins AS (2019) Late Neoproterozoic adakitic magmatism of the eastern Arabian Nubian Shield. Geosci Front 10 (6):1981–1992 Cox G, Lewis CJ, Collins AS, Halverson GP, Foden J, Jourdan F, Nettle D, Kattan F (2012) Ediacaran terrane accretion in the Arabian Nubian Shield. Gond Res 21:341–352 Delfour J (1979) Explanatory notes to the geologic map of the Halaban quadrangle, Sheet 23G Kingdom of Saudi Arabia. Bur Rech Geol Min, Orleans Doebrich JL, Al-Jehani AM, Siddiqui AA, Hayes TS, Wooden JL, Johnson PR (2007) Geology and metallogeny of the Ar Rayn terrane, eastern Arabian shield: evolution of a Neoproterozoic continental-margin arc during assembly of Gondwana within the East African orogeny. Prec Res 158:17–50 El-Bialy MZ (2004) Petrologic, geochemical and petrogenetic characterization of the old granites of Sinai, Egypt. PhD thesis, Suez Canal, University, Ismailia El-Bialy MZ (2010) On the Pan-African transition of the Arabian-Nubian Shield from compression to extension: The post-collision Dokhan volcanic suite of Kid-Malhak region, Sinai, Egypt. Gond Res 17:26–43 El-Bialy MZ (2013) Geochemistry of the Neoproterozoicmetasediments of Malhaq and Um Zariq formations, Kid Metamorphic Complex, Sinai, Egypt: implications for source-area weathering, provenance, recycling and depositional tectonic setting. Lithos 175:68–85 El-Bialy MZ (2020) Precambrian basement complex of Egypt. In: Hamimi Z, El-Barkooky A, Martínez Frías J, Fritz H, Abd El-Rahman Y (eds) The Geology of Egypt. Regional Geology Reviews, Springer, Cham, pp 37–79 El-Bialy MZ, Omar MM (2015) Spatial association of Neoproterozoic continental arc I-type and post-collision A-type granitoids in the Arabian-Nubian Shield: the Wadi Al-Baroud Older and younger granites, North Eastern Desert, Egypt. J Afr Earth Sci 103:1–29 El-Bayoumi R, Abdel Wahed M, Tobschall HJ, Shallaly NA (2006). The Counter-Clockwise Pressure-Temperature-Time Path of metavolcanics and gneissose diorite from Central Wadi Kid Area, southeast Sinai, Egypt. In: Proceedings of the 8th International Conference on the Geology of the Arab World (GAW 8), Cairo University I, pp 233–240 El-Gaby S, El-Nady O, Khudeir A (1984) Tectonic evolution of the basement complex in the Central Eastern Desert of Egypt. Geol Rund 73:1019–1036 Eliwa HA, El-Bialy MZ, Murata M (2014) Ediacaran post-collisional volcanism in the Arabian-Nubian Shield: the high-K calc-alkaline

436 Dokhan Volcanics of GabalSamr El-Qaa (592 ± 5 Ma), North Eastern Desert, Egypt. Prec Res 246:180–20 El-Mahallawi MM, Ahmed AF (2012) Late Proterozoic older granitoids from the North Eastern Desert of Egypt: petrogenesis and geodynamic implications. Arab J Geosci 5:1527 El-Metwaly AA (1993) Microgranular enclaves in the Pan-African I-type granites from the Sinai Massif: petrology, mineralogy and geochemistry. J Afr Earth Sci 17:95110 El-Metwally AA, El Aasy IE, Ibrahim ME, Essawy MA, El-Mowafy AA (1999) Petrological, structural and geochemical studies on the basement rocks of Gabal Umm Zariq-Wadi Kid area, Southeastern Sinai, Egypt. J Geol 43:147–180 El Ramly MF (1972) A new geological map for the basement rocks in the Eastern and Southwestern Deserts of Egypt. Ann Geol Surv Egypt 2:1–18 El-Shazly AK, Khalil KI (2016) Metamorphic and geochronologic constraints on the tectonic evolution of the Central Eastern Desert of Egypt. Prec Res 283:144–168 Escuder-Viruete J, Valverde-Vaquero P, Rojas-Agramonte Y, Jabites J, Pérez-Estaún A (2013) From intra-oceanic subduction to arc accretion and arc-continent collision: Insights from the structural evolution of the Río San Juan metamorphic complex, northern Hispaniola. J Struc Geol 46:34–56 Eyal M, Be’erin-Shlevin Y, Eyal Y, Whitehouse M, Litvinovsky B (2014) Three successive Proterozoic island arcs in the Northern Arabian-Nubian Shield: Evidence from SIMS U-Pb dating of zircon. Gond Res 25(1):338–357 Farahat ES, Mohamed HA, Ahmed AF, El Mahallawi MM (2007) Origin of I- and Atype granitoids from the Eastern Desert of Egypt: implications for crustal growth in the northern Arabian-Nubian Shield. J Afr Earth Sci 49:43–58 Farahat ES, Zaki R, Hauzenberger C, Sami M (2011) Neoproterozoic calc-alkaline peraluminous granitoids of the Deleihimmi pluton, Central Eastern Desert, Egypt: implication for transition from late to post-collisional tectonomagmatic evolution of the northern Arabian-Nubian Shield. Geol J 46:544–560 Ferreira ACD, Dantas EL, Fuck RA, Nedel IM (2020) Arc accretion and crustal reworking from late Archean to Neoproterozoic in Northeast Brazil. Sci Rep 10:7855 Fritz H, Abdelsalam M, Ali KA, Bingen B, Collins AS, Fowler AR, Ghebreab W, Hauzenberger CA, Johnson PR, Kusky TM, Mace P, Muhongo S, Stern RG, Viola G (2013) Orogen styles in the East African Orogen: a review of the Neoproterozoic to Cambrian tectonic evolution. J Afr Earth Sci 86:65–106 Fritz H, Dalmeyer DR, Wallbrecher E, Loizenbauer J, Hoinkes G, Neumayr P, Khudeir AA (2002) Neoproterozoic tectonothermal evolution of the Central Eastern Desert, Egypt: a slow velocity tectonic process of core complex exhumation. J Afr Earth Sci 34:543–576 Furnes H, Shimron AE, Roberts D (1985) Geochemistry of Pan-African volcanic arc sequences in southeastern Sinai Peninsula and plate tectonic implications. Prec Res 29:359–382 Fowler A, Osman AF (2009) The Sha’it-Nugrus shear zone separating central and south eastern deserts, Egypt: a post-arc collision low-angle normal ductile shear zone. J Afr Earth Sci 53:16–32 Fowler A, Osman AF (2013) Sedimentation and inversion history of three molasse basins of the western Central Eastern Desert of Egypt: implications for the tectonic significance of Hammamat basins. Gond Res 23:1511–1534 Fowler AR, Hamimi Z (2020) Structural and tectonic framework of Neoproterozoic basement of Egypt: from gneiss domes to transpression belts. In: Hamimi Z, El-Barkooky A, Martínez Frías J, Fritz H, Abd El-Rahman Y (eds) The Geology of Egypt. Regional Geology Reviews, Springer, Cham, pp 81–129

A. A. Surour Fowler A, Hassan IH, Osman AF (2010) Neoproterozoic structural evolution of SE Sinai, Egypt: II. Convergent tectonic history of the continental arc Kid Group. J Afr Earth Sci 58:526546 Gad S, Kusky T (2007) ASTER spectral ratioing for lithological mapping in the Arabian-Nubian shield, the NeoproterozoicWadi Kid area, Sinai, Egypt. Gond Res 11:326–335 Genna A, Nehlig P, Le Goff E, Guerrot C, Shanti M (2002) Proterozoic tectonism of the Arabian Shield. Prec Res 117:21–40 Greiling RO, de Wall H, Sadek MF, Dietl C (2014) Late Pan-African granite emplacement during regional deformation, evidence from magnetic fabric and structural studies in the Hammamat-Atalla area, Central Eastern Desert of Egypt. J Afr Earth Sci 99:109–121 Hamimi Z (1996) Tectonic evolution of the shield rocks of Gabal El Sibai area, Central Eastern Desert, Egypt. Egypt J Geol 40:423–453 Hamimi Z, Abdel-Wahed MA (2020) Suture(s) and major shear zones in the Neoproterozoic basement of Egypt. In: Hamimi Z, El-Barkooky A, Martínez Frías J, Fritz H, Abd El-Rahman Y (eds) The Geology of Egypt. Regional Geology Reviews, Springer, Cham, pp 153–189 Hamimi Z, Abd El-Wahed MA, Gahlan HA, Kamh SZ (2019) Tectonics of the Eastern Desert of Egypt: Key to Understanding the Neoproterozoic Evolution of the Arabian-Nubian Shield (East African Orogen). In: Bendaoud A, Hamimi Z, Hamoudi M, Djemai S, Zoheir B (eds) The Geology of the Arab World - An Overview. Regional Geology Reviews, Springer, Cham, pp 1–81 Hamimi Z, El Amawy MA, Wetait M (1994) Geology and structural evolution of El Shalul Dome and environs, central Eastern Desert, Egypt. Egypt J Geol 38(2):575–595 Hassan MA, Hashad AH (1990) Precambrian of Egypt. In: Said R (ed) The geology of Egypt. Balkema, Rotterdam, pp 201–248 Hawkesworth C, Cawood PA, Dhuime B (2019) Rates of generation and growth of the continental crust. Geosci Front 10:165–173 Hsu KJ (2003) Role of ophiolites in archipelago model of orogenesis. Geological Society of America 373:159–171 Hussein AA, Ali MM, El Ramly FM (1982) A proposed new classification of the granites of Egypt. J Vol Geotherm Res 14 (1-2):187–198 Jacobs J, Thomas RJ (2004) Himalayan-type indenter-escape tectonics model for the southern part of the late Neoproterozoic-early Paleozoic East African Antarctic orogen. Geology 32:721–724 Johnson PR, Andresen A, Collins AS, Fowler AR, Fritz H, Ghebreab W, Kusky T, Stern RJ (2011) Late Cryogenian-Ediacaran history of the Arabian-Nubian Shield: a review of depositional, plutonic, structural, and tectonic events in the closing stages of the northern East African Orogen. J Afr Earth Sci 61(3):167–232 Johnson PR, Kattan F (2008) Lithostratigraphic revision in the Arabian Shield: the impacts of geochronology and tectonic analysis. Arab J Sci Engin 33(1C):3–16 Johnson PR, Stewart ICF (1995) Magnetically inferred basement structure in central Saudi Arabia. Tectonophysics 245(1–2):37–52 Johnson PR, Woldehaimanot B (2003) Development of the Arabian-Nubian Shield: perspectives on accretion and deformation in the northern East African Orogen and the assembly of Gondwana. In: Yoshida M, Windley BF, Dasgupta S (Eds), Proterozoic East Gondwana: Supercontinent Assembly and Breakup, Geol Soc London Sp Publ 206, pp 289–326 Khudeir AA, Abu El-Rus MA, El-Gaby S, El-Nady O, Bishara WW (2008) Sr-Nd isotopes and geochemistry of the infrastructural rocks in the Meatiq and Hafafit core complexes, Eastern Desert, Egypt: evidence for involvement of Pre-Neoproterozoic crust in the growth of Arabian-Nubian shield. Island Arc 17:90–108 Kröner A, Todt W, Hussein IM, Mansour M, Rashwan AA (1992) Dating of late Proterozoic ophiolites in Egypt and the Sudan using the single grain zircon evaporation technique. Prec Res 59:15–32

17

Arc Accretion and Calc-Alkaline Plutonism …

Lee C-TA, Morton DM, Kistler RW, Baird AK (2007) Petrology and tectonics of Phanerozoic continent formation: From island arcs to accretion and continental arc magmatism. Earth Planet Sci Lett 263 (3–4):370–387 Loizenbauer J, Wallbrecher E, Fritz H, Neumayr P, Khudeir AA, Kloetzli U (2001) Structural geology, simple zircon ages and fluid inclusion studies of the Meatiq metamorphic core complex: implications for Neoproterozoic tectonics in the Eastern Desert of Egypt. Prec Res 110:357–383 Lundmark AM, Andersen A, Hassan MA, Augland LE, Boghdady GY (2012) Repeated magmatic pulses in the East African Orogen in the Eastern Desert, Egypt: An old idea supported by new evidence. Gond Res 22(1):227–237 Makroum F (2017) Structural interpretation of the Wadi Hafafit culmination: a Pan African gneissic dome in the central Eastern Desert, Egypt. Lithosphere 9(5):759–773 Moore JM (1979) Tectonics of the Najd transcurrent fault system, Saudi Arabia. J Geol Soc London 136:441–454 Moussa EMM, Stern RJ, Manton WI, Ali KA (2008) SHRIMP zircon dating and Sm/Nd isotopic investigations of Neoproterozoic granitoids, Eastern Desert, Egypt. Prec Res 160:341–356 Nawab RJ (1979) Geology of the Al-Amar-Idsas region of the Arabian Shield. Inst Appl Geol Bull Jeddah 3:29–40 Nettle D, Halverson GP, Cox GM, Collins AS, Schmitz M, Gehling J, Johnson PR, Kadi K (2014) A middle-late Ediacaran volcano-sedimentary record from the eastern Arabian-Nubian shield. Terra Nova 26(2):120–129 Pallister JS, Stacy JS, Fisher LB, Premo WR (1988) Precambrian ophiolites of Arabia: geological settings, U-Pb geochronology, Pb-isotope characteristics, and implications for crustal accretion. Prec Res 38:1–54 Rollinson H (2010) Coupled evolution of Archean continental crust and subcontinental lithospheric mantle. Geology 38:1083–1086 Sattler HR (1995) Our patchwork planet: the story of plate tectonics. Maestro, Giulio 1st Ed, Lothrop, Lee & Shepard Books, New York Sengor AMC, Natal’in BA (1996) Turkic-type orogeny and its role in the making of the continental crust. Ann Rev Earth Planet Sci 24:263–337 Shallaly NAM (2006) Geology of central Wadi Kid area, southern Sinai, Egypt. PhD Thesis, Cairo University, Giza Shimron AE (1980) Proterozoic island arc volcanism and sedimentation in Sinai. Prec Res 12:437–458 Shimron AE (1987) Pan-African metamorphism and deformation in the Wadi Kid region, SE Sinai Peninsula: evidence from porphyroblasts in the Um Zariq Formation. Israel Earth Sci 36:173–193 Stacey JS, Stoeser DB, Greenwood WR, Fischer LB (1984) U-Pb zircon geochronology and geological evolution of the Halaban–Al Amar region of the Eastern Arabian Shield, Kingdom of Saudi Arabia. J Geol Soc London 14:1043–1055 Stein M, Goldstein S (1996) From plume head to continental lithosphere in the Arabian-Nubian Shield. Nature 382:773–778 Stern RJ (1979) Late Precambrian crustal environments as reconstructed from relict igneous minerals, Central Eastern Desert of Egypt. Ann Geol Surv Egypt 9:9–13

437 Stern RJ (1981) Petrogenesis and tectonic setting of late Precambrian ensimatic volcanic rocks, Central Eastern Desert of Egypt. Prec Res 16:195–230 Stern RJ (1985) The Najd fault system, Saudi Arabia and Egypt: a late Precambrian rift-related transform system. Tectonics 4:497–511 Stern RJ (1994) Arc assembly and continental collision in the Neoproterozoic East African orogen: implications for the consolidation of Gondwanaland. Ann Rev Earth Planet Sci 22:319–351 Stern RJ (2008) Neoproterozoic crustal growth: the solid Earth system during a critical episode of Earth history. Gond Res 14:33–50 Stern RJ (2018) The evolution of plate tectonics. Philosophical Trans Royal Soc A 376:20170406 Stern RJ, Hedge CE (1985) Geochronologic constraints on late Precambrian crustal evolution in the Eastern Desert of Egypt. Amer J Sci 285:97–127 Stern RJ, Johnson PR (2010) Continental lithosphere of the Arabian Plate: a geologic, petrologic, and geophysical synthesis. Earth-Sci Rev 101:29–67 Stoeser DB, Camp VE (1985) Pan-African microplate accretion of the Arabian Shield. Geol Soc Amer Bull 96:817–826 Stoeser DB, Frost CD (2006) Nd, Pb, Sr, and O isotopic characterization of Saudi Arabian Shield terranes. Chem Geol 226:163–188 Stoeser DB, Whitehouse MJ, Stacey JS (2001) The Khida Terranegeology of Paleoproterozoic rocks in the Muhayil Area, eastern Arabian Shield, Saudi Arabia. Gond Res 4:192–194 Surour AA, Kabesh ML (1998) Calc-alkaline magmatism and associated mafic microgranular enclaves of Wadi Risasa area, Southeastern Sinai, Egypt. Ann Geol Surv Egypt 21:35–54 Vail JR (1985) Pan-African (late Precambrian) tectonic terrains and the reconstruction of the Arabian-Nubian shield. Geology 13:839–842 van der Pluijm BA (2004). Earth structure: an introduction to structural geology and tectonics. Marshak, Stephen, 1955 (2nd ed), Norton, New York Whitehouse MJ, Pease V, Al-Khirbash S (2016) Neoproterozoic crustal growth at the margin of the East Gondwana continent- age and isotopic constraints from the easternmost inliers of Oman. Inter Geol Rev 58(16):2046–2064 Whitehouse MJ, Stoeser DB, Stacey JS (2001a) The Khida terrane geochronological and isotopic evidence for Paleoproterozoic and Archean crust in the eastern Arabian Shield of Saudi Arabia. Gond Res 4:200–202 Whitehouse MJ, Windley BF, Stoeser DB, Al-Khirbash S, Ba-Bttat MAO, Haider A (2001b) Precambrian basement character of Yemen and correlations with Saudi Arabia and Somalia. Prec Res 105:357– 369 Whitehouse MJ, Windley BF, Ba-Bttat MAO, Fanning CM, Rex DC (1998) Crustal evolution and terrane correlation in the eastern Arabian Shield, Yemen: geochronological constraints. J Geol Soc London 155:281–295 Windley BF, Whitehouse MJ, Ba-Bttat MAO (1996) Early Precambrian gneiss terranes and Pan-African island arcs in Yemen: crustal accretion of the eastern Arabian Shield. Geology 24:131–134

18

Origin of the Volcanic-Arc Signature in Late-Orogenic Granitoids from the Arabian–Nubian Shield Basem Zoheir and Aliaa Diab

Abstract

Although formed after subduction tectonics have already waned and terrane accretion has concluded, numerous late-orogenic granitoid phases in the Neoproterozoic Arabian–Nubian Shield (ANS) hold geochemical characteristics typical of the volcanic-arc granites (VAG). The latter are accentuated by the high abundances of LILE and LREE relative to the HFSE and HREE, and thus high Ba/Nb, Ba/Zr, Sr/Y, La/Yb ratios along with fractionated REE patterns. Synthesis of bulk-rock geochemical and zircon Hf isotope data of the ANS late-orogenic granitoids, exclusively dated as *640–600 Ma, is employed herein to reveal the origin of the VAG signature in many of these rocks. Data of widespread granitoid intrusions reveal that these rocks span typically metaluminous to peraluminous, calc-alkalic I-type to alkali-calcic ferroan A-type granite compositions. Mild-to-strong negative Eu anomalies generally correspond to increasing magmatic differentiation, and transition from hybrid to evolved granites is envisaged for the younger phases, irrespective of their distribution. The available bulk-rock geochemical data and zircon Hf isotope compositions weigh with parental magmas produced by high-temperature melting of subduction-related/subduction-metasomatised lower crustal rocks (e.g., TTGs, high-K mafic rocks, and metasediments) and mixing with mantle-derived sanukitoid-like melts. These magmas then experienced plagioclase-rich/poor fractionation (revealed by variable negative Eu anomalies) prior to crystallization and fractionation toward A‐type/ferroan granite geochemistry. The heat source and melting trigger were most likely B. Zoheir (&) Department of Geology, Faculty of Science, Benha University, 13518 Benha, Egypt e-mail: [email protected] A. Diab Department of Geology, Faculty of Science, Ain Shams University, Cairo, Egypt

asthenospheric mantle upwelling during the onset of extensional collapse of the ANS. The widespread distribution of the post-collisional granitic intrusions in most of the shield area can be explained by lithospheric delamination generally results in voluminous rather than restricted linear magmatism. The commonly described VAG signature in some ANS late-orogenic granitoids ( Kareim > Um Tawat > El Qash > Igla > Abu Sheqeili > Abu Gheryan (2.24). These strains were obtained by the Rf/u method for pebble sectional ellipses, and the Fry method using pebble centre-to-centre distances, deformed rain prints and mudcracks (Fig. 19.10d). Fowler and Abdeen (2014) explored the bases of these calculations and compared them to strain data from sand particles in the same outcrops, and to the shape parameters of pebbles from undeformed outcrops. They found that the strain ellipsoid from the sandstones was barely detectable above the primary fabric ellipsoid of the sandstones, even in rocks showing a crude cleavage, for which values of RXZ 1.20–1.30 were representative (Fig. 19.10e). Preferred orientation of pebbles with larger Ri led to tendencies for Rf values to be significantly overstated. The average Wadell sphericity of the undeformed pebbles showed a narrow range of 0.66–0.72 for both volcanic and granitic pebbles, regardless of pebble size, and could constitute an index for recognizing strained pebble shapes in these basins (Fowler and Hamimi 2021). The average sphericity of deformed pebbles in the Um Seleimat, Zeidun and Arak basins was 0.50, 0.49, 0.42, respectively. Similarly, the % spheroid pebbles in undeformed conglomerates ranged from 26 to 33%, dropping very quickly to 0–11% in the foliated metaconglomerates. Based on cross-sectional restoration, Abdeen and Greiling (2005) estimated the amount of NW–SE shortening in the Queih basin to be e = −25%. Using the cross-sections of the Zeidun and Arak basins provided by Fowler and Osman (2013), the N–S shortening of the Zeidun basin is about e = −18%, while the NW–SE shortening of the Arak basin is about e = −21%. The NW–SE shortening in the El Qash basin to the N (based on the presence of NE–SW trending folds with limb dips of 50–70°), gives e = −30%. However,

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A.-R. Fowler and Z. Hamimi

Fig. 19.10 Deformation features of the Hammamat associated with the Eastern Desert Shear Zone (EDSZ). a EDSZ-related foliation inclined to bedding (represented by the sandy layer above lens cap) (Wadi Hammamat). b EDSZ-related foliation inclined to bedding (Arak). c Foliated metaconglomerate with stretching lineation parallel to pencil (Um Seleimat). d Deformed mudcracks in siltstone (Kareim). e Cleavage developed in coarse-grained sandstone (Kareim). f Deformed cobble in metaconglomerate (Um Seleimat). g Undeformed and h foliated lithic arenite, demonstrating the changes occurring during cleavage formation in Hammamat sandstones (Um Seleimat). Sericitization of lithic particles and growth of metamorphic white mica and biotite accompanies cleavage formation

19

Post-amalgamation Depositional Basins in the Arabian-Nubian …

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Fig. 19.11 Map of the Central Eastern Desert showing the approximate extent of the Eastern Desert Shear Zone (EDSZ) (bold red dashed line) and how this area coincides with the limits of the distribution of the larger Hammamat basins. Areas of extension-related foliations and displacement sense of extension-related shears are also presented. The Shait-Nugrus Shear Zone (SNSZ) is shown; however, the relations between the SNSZ and EDSZ are not presently known

the Wadi Hammamat basin farther north lacks any obvious signs of NW–SE shortening, and has NW–SE trending approximately cylindrical folds (Fowler and Osman 2001). An estimate of NW–SE shortening for Kareim basin (our unpublished structural data) yields e = −21%. Abdeen and Greiling (2005) gave a rough estimate of the NE-SW shortening of the Queih basin at e = −15 to −17%. For Wadi Hammamat basin, the ENE shortening is a minimum of e = −26%, as thrusts displacements are unknown in the section. For the section from Meatiq westwards to the Nubia sandstone contact, Fowler and Hamimi (2020) estimated shortening at e = −20%. Folding at Um Seleimat is tighter, in accord with NE–SW shortening of about −30 to −35%. The NE–SW shortening for El Qash is approximately e = −17%. Our estimates of ENE–WSW shortening for Zeidun and Arak, based on cross-sections, are low at −9% and −15%, respectively. A similarly low value of NE–SW shortening at Kareim is estimated to be −9%. From these figures, it appears that the NW–SE shortening is typically

−20 to −25%, and the NE–SW shortening is more variable from −10 to −35%; however, exclusive of the Wadi Hammamat and Um Seleimat basins, it is typically −10 to −15%.

19.11

Basal Unconformity, Boundary Faults, Basin Shape

For most of the Hammamat basins, the present-day contacts with basement rocks are either reactivated original marginal faults or faults that formed during basin inversion. Preserved basal unconformities, however, have been identified in many of the basins. Basal unconformity has been reported for the Igla basin (Akaad 1957; Abdel Khalek and Hafez 1983; Rice et al. 1993b), Abu Gheryan basin (Noweir et al. 1983), Meesar basin (Osman 1996), Arak basin (Fowler and Osman 2013), Um Tawat basin (Osman et al. 2001); Kareim basin (El-Shazly, 1977; Noweir et al. 2005; Fowler et al. 2020, Kiyokawa et al. 2020), Queih basin (Abdeen and Warr 1998;

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Abdeen and Greiling 2005; Khalaf 2012), Abu Sheqeili basin (Abd El-Wahed 2007), Ranga, El Miyah, Wadi Bali and Wadi Dib basins (Akaad and El Ramly 1958). In all, except Wadi Bali, the unconformity was marked by basal conglomerates containing clasts of the underlying rocks. The rocks immediately below the unconformity are mainly metavolcanics and diorites, but Dokhan volcanics are also found, e.g. at Ranga and Queih basins. No examples have been found of Hammamat resting unconformably on gneissic rocks. The Hammamat beds at the unconformity dip

A.-R. Fowler and Z. Hamimi

gently to steeply (25°–60°) towards the E, N and S. Where bedding is present in the basement rocks (e.g. bedded Dokhan volcanics), the unconformity shows angular discordance, elsewhere it is a non-conformity. The shapes of the Hammamat basins are influenced by the orientations of the boundary faults. For most of the Hammamat basins (large and small), faults account for >80% of their boundaries (Fig. 19.12a), the remainder being unconformities, intrusive contacts or later sedimentary cover. Many of the faults are not in the ideal orientation with respect to the regional stresses, as they are reactivated fractures with inherited orientations. This results in common oblique-slip sense on the faults. The NW-elongate Wadi Hammamat basin is outlined by NW-trending thrusts and sinistral strike-slip faults (Fritz et al. 1996; Fowler and Osman 2001). Zeidun basin has a square shape confined by E–W reverse sinistral, and N–S (reverse) dextral faults, while Arak basin shares the E–W thrusts and N–S dextral faults, and has also NNW-trending thrusts (Fowler and Osman 2013) (Fig. 19.12a). The rhombic shape of the El Qash basin is defined by NW to WNW normal faults and NNW to N–S sinistral and dextral strike-slip faults (Abd El-Wahed 2010). The more polygon-shaped Kareim has well-developed NW, WNW and ENE faults that include thrusts, sinistral strike-slip and tear faults (Fig. 19.12b). There are also N–S (probably reverse) sinistral faults, and NE normal faults (Fritz and Messner 1999; Noweir et al. 2005; Akawy and Zaky 2008; Fowler et al. 2020). The Queih basin boundaries are mostly NE- and NW-trending thrusts, ENE-trending normal faults and NW sinistral strike-slip faults. The smaller Hammamat basins usually have a single dominant marginal fault set. These are mostly normal faults (N–S, NE or E–W trending), e.g. Sheqeili, Igla, Abu Gheryan, El Miyah and Um Tawat basins, though thrusts with these orientations have also been identified as boundary faults, e.g. Igla basin (Fritz et al. 2002).

19.12

Fig. 19.12 Within basin structures of some examples of the large Hammamat basins. a The Zeidun and Arak basins (yellow areas) showing E–W to NE–SW trending early folds (represented by yellow arrows) cut by faults with E–W, N–S, NE and NW trends that also define the present-day basin margins. Red arrows give the trend of smaller macroscopic later folds that refold the earlier folds and may also be related to fault compartmentalization of the deforming basin. Grey areas are undifferentiated basement rocks. b Kareim basin showing the segmented nature of the NE to E–W trending syncline deforming the sequence. The segmentation of the western half of the syncline is fault-related. The segmentation of the eastern half is due to folding of mutually inclined surfaces separated by angular unconformities. Different colours represent different stratigraphic units, according to Fowler et al. (2020)

Internal Deformation Structure— Folds and Faults

Some macrofolds in the larger Hammamat basins appear to be among the earliest deformation structures in the basins, as they are cut by the basin faults. Three examples of these folds will be briefly described. The first example is in the Arak basin where the sediments are deformed by a WSW-trending syncline that can be traced for 4.5 km along its S-dipping axial plane (Fig. 19.12a). The hinge plunges 33° SW (Makroum et al. 2001; Fowler and Osman 2013). The southern limb has overturned beds. The interlimb angle is estimated to be about 35–40°. The syncline is refolded about later SSE-trending folds. The syncline may have formed as a passive structure draped over syn-sedimentary

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basement faults, or it may be a forced fold associated with thrusting at the southern margin. The second example is in the southern part of the Zeidun basin, where an ENE-trending syncline was mapped by Fowler and Osman (2013) (Fig. 19.12a). The fold extends for 9.5 km along its steeply N-dipping axial plane. The hinge plunges gently to the west. The interlimb angle is estimated to be about 85°. A large monoclinal fold parallels the syncline to its north, interpreted by Fowler and Osman to be due to the inversion of a basin floor normal fault as a reverse fault. The third example is in the Kareim basin, where another large syncline exists. This syncline has been illustrated as having a continuous 16 km long axial plane curving from a NE trend at the western end of the fold to an E–W trend in the east, but is in fact a set of en échelon axial plane segments, due in part to faulting, but mainly due to folding of obliquely inclined surfaces (Ramsay 1967, pp 509–510) (Fig. 19.12b). The obliquity of layers is a result of angular unconformities within this basin (Fowler et al. 2020). The Kareim syncline plunges about 12° to the west and has an approximately 90° interlimb angle. There is an axial plane cleavage in the hinge dipping 60° NW, so this fold appears to be tectonic. The beds on the limbs are refolded about much smaller folds with N–S axial planes. There are other folds, with quite different style in the larger Hammamat basins. These folds have a mainly elliptical basinal geometry and may be described as periclines. They have large interlimb angles (*120°) and can be traced along their axial planes for 2–3 km. The locations and trends of these folds are controlled by within-basin tear faults, which have acted as compartmental faults allowing segments of the sequence between them to be folded individually, so that the folds cannot be followed across the faults into the next compartment. Examples of these periclines are found in the northern part of the Zeidun basin and the southern part of the El Qash basin (Fowler and Osman 2001, 2013) (Fig. 19.12a). The faults within the larger Hammamat basins include those extending from the margins, and fault sets parallel to those defining the margins. Thus, N–S to NNW dextral strike-slip faults and E–W thrusts are common, along with NW- to NNW-trending sinistral strike-slip faults and thrusts, and NE- to ENE-trending dextral strike-slip faults. In the smaller Hammamat basins, there is usually only one trend of folding, and is commonly a NE-trending subhorizontal fold set, especially where this is also the long axis direction of the basin. In narrow basins such as Um Tawat, Gebel El Urf, eastern El Miyah and Abu Sheqeili basin, it is a single upright to inclined syncline. Wider basins such as Igla and Queih have sets of folds with this and NW orientations. The interlimb angles of the synclines are variable but in the range 70–120°. Other basins may show steep

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to vertical or overturned bedding rather than folding, e.g. western El Miyah basin. The dominant faults within the smaller basins also trend parallel to the basin elongation and are represented by mainly NE-trending dextral or sinistral strike-slip faults. The basins are also cut by NW-trending dextral or sinistral strike-slip faults. The fault and fold deformations of the basins very likely resulted in termination of sedimentation and basin inversion (i.e. the proportion of syn-sedimentary normal fault displacement that has been completely reversed during later compression). This issue is further considered in the discussion.

19.13

Metamorphism

The status of metamorphism in the Hammamat molasse was a topic of debate in the 1960s—1980s, with a special focus on the Um Esh—Um Seleimat area, just east of the Hammamat Group type section in the CED. At Wadi Um Esh, Andrew (1939) reported low-grade metaconglomerates with green biotite in the foliation. These lay at the northwestern end of a narrow belt of deformed metaconglomerates that, when traced farther south, were found to be in direct contact with Hammamat conglomerates of the El Qash basin at the latitudes of Wadi Um Seleimat and Wadi Muweilih. From the early stages of development of a stratigraphic system for the EED (which culminated in Akaad and Noweir’s influential scheme in 1969) the Hammamat was deemed to be essentially unmetamorphosed and to lie above a well-defined regional unconformity that post-dated regional metamorphism. Therefore, Akaad and Noweir (1964) and Akaad (1972) argued against the Um Esh—Um Seleimat metaconglomerates being included in Hammamat. Instead, they identified the metaconglomerates as Atud Formation metasediments of the older Abu Ziran Group. Other workers, however, accepted these metaconglomerates as Hammamat (Stern 1979; Ries et al. 1983; Abu El Ela and El Bahariya 1998, Fowler and El Kalioubi 2004). Ries et al. (1983) and Abu El Ela and El Bahariya (1998) found the grade of metamorphism to increase from greenschist facies to amphibolite facies in the direction of the Meatiq Complex. Most other Hammamat basins have also exhibited low temperature effects, mainly deep diagenetic (zeolite) to epizonal (Chlorite Zone), revealed by illite and chlorite crystallinity studies. Grades between diagenesis and anchizone (zeolite facies) were found for the Igla basin (Samuel et al. 1978), and Zeidun basin (Rice et al. 1993a); and between incipient and very weak metamorphism (upper anchizone to lower greenschist facies) for the Queih basin (Ahmed et al. 1989; Rice et al. 1993a) and Wadi Hammamat basin (Soliman 1983; Messner 1996). Higher grades of upper

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anchizone (transition from prehnite-pumpelleyite to lower greenschist facies) were found for the Arak basin (Osman et al. 1993; Rice et al. 1993a); while epizonal conditions (greenschist Chlorite to Biotite Zones) were found for the Meesar basin (Rice et al. 1993a; Osman 1996), Kareim basin (Messner 1996; Fritz and Messner 1999), Um Tawat basin (Willis et al. 1988), and El Miyah basin (Kamal El-Din and Asran 1995; Shalaby et al. 2006) (Fig. 19.9). It was also discovered that grade increased towards the lower stratigraphic sections, where the molasse sediments became foliated and sheared (Messner 1996; Fowler and El Kalioubi 2004). Fritz and Messner’s (1999) illite-chlorite crystallinity data show higher temperature epizonal conditions in stratigraphically lower parts of the Kareim basin (Units 1–3 of Fowler et al. 2020), compared to the higher parts (Units 5–8). Using estimates of 200 °C and 300 °C to mark the diagenesis-anchizone and anchizone-epizone boundaries, respectively, and assuming the highest exposed units in the Hammamat basins, that show no signs of post-burial change, equilibrated at 100 °C km−1 and highly variable, even for adjacent basins (e.g. Zeidun 35–40 °C km−1 and nearby Arak 70–80 °C km−1). The calculated gradients are therefore very unlikely to represent the background geotherm for crust at that time. The high values and their variability are probably best explained by variable contributions of magmatic heat beneath these basins (Fritz and Messner 1999). The acceptance of metamorphism in the Hammamat, albeit constrained to the foliated lowermost strata of the basins, was also partly a result of El-Gaby et al.’s (1988) EED tectonic model that compared immature silty, sandy and pebbly metasediments and interbedded silicic metavolcanics occupying the Allaqi “swell” in the SED, with the Hammamat molasse and interbedded Dokhan volcanics. These immature metasediments were thought by El-Gaby to have filled a giant foreland basin in a similar setting to the Hammamat. The metasediments experienced low pressure— high temperature metamorphism (Abukuma or “Buchan” series), apparently due to the thermal effects of large gneissic granitoid intrusions. El-Gaby proceeded to correlate or even reassign several other low pressure immature metasediment sequences in the EED to the Hammamat Group, including high-grade metasediments enclosing the Um Had granite, and medium to high-grade metasediments east of the El Sibai complex (El-Gaby et al. 1988b; El-Gaby 1994). He further extended Hammamat status to metasediment sequences in the Sinai, including the Zaghra Conglomerates (El-Gaby et al. 2002), and the Kid Conglomerates (El-Gaby et al. 1991).

A.-R. Fowler and Z. Hamimi

19.14

Geochronology

Isotopic dating methods have been applied to the Hammamat sediments and associated Dokhan volcanics and Younger Granites to best define the beginning and end of molasse deposition, and to put constraints on the nature of the provenance for the sediments of these basins. Each basin must be individually dated as they appear to have different age ranges, at least relative to other events. For example, some Hammamat sequences are syn-Dokhan (Kharaza, Hamid, El Urf, Wadi Bali), others suggested to be post-Dokhan (Wadi Hammamat); some postdate the Younger Granites (Igla), and some predate them (W. Hammamat, Arak); some are older than the “post-Hammamat” felsites (Zeidun), whereas some are syn-felsite (Queih) (Rice et al. 1993a). In addition, there are reportedly separate age ranges for CED Hammamat basins (639–602 Ma) and NED basins (595–575 Ma) (El Kalioubi et al. 2020). Early geochronology relied on Rb–Sr whole rock and K– Ar data. In the CED, Ries et al. (1983) limited Wadi Hammamat basin sedimentation to lie between the Dokhan volcanics and the Um Had granite (Rb–Sr whole rock 616 ± 9 Ma, and 590 ± 11 Ma, respectively). Stern (1979) also used Rb–Sr method to date Dokhan in the CED to 618– 602 Ma, with errors of *10 Ma. In the NED, Massey (1984) (quoted by Stern and Voegeli 1987) used Rb–Sr to estimate Hammamat deposition in the Um Tawat area at *600 Ma. The same methods yielded age ranges for NED Dokhan (600–575 Ma), epizonal Younger Granites (600– 570 Ma) and crustal extension (600–575 Ma) (Stern et al. 1984; Stern and Hedge 1985). Moghazi (2003) gave slightly different Rb–Sr-based age ranges for NED Dokhan (610– 560 Ma); Younger Granites (610–550 Ma) and Hammamat (600–585 Ma). A narrower range for NED Dokhan (594– 583 Ma) was given by Stern et al. (1988). Willis et al.’s (1988) whole rock Rb–Sr dating of Um Tawat sandstones and mudstones gave 585 ± 15 Ma, and their coarser clay fraction gave K–Ar age of 588–567 Ma (the clays were cements and diagenetic replacements), yielding an overall age of *590 Ma for the NED Hammamat. SHRIMP, SIMS and TIMS U–Pb zircon studies have provided new age data defining the Hammamat depositional interval and the characteristics of the provenance rocks. In the CED, Andresen et al. (2009) confirmed the previous age estimates of the Um Had granite at 598–595 Ma, and further constrained Hammamat deposition to be before regional folding at 605–600 Ma. Bezenjani et al. (2014) discovered that the top of the Igla basin sequence cannot be older than 628 ± 6 Ma, and that rhyolitic clasts (dated 700 ± 6 Ma) in the formation were older than Dokhan. They found that 98% of the zircons collected were older than Dokhan (peaks at 690 and 652 Ma). Their data indicated island arc volcanic

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provenance, not Dokhan. Fielding et al. (2017) estimated the maximum age of deposition of the Wadi Hammamat molasse as 635 Ma. They found that the zircon population of an arenite sample was dominated by ages 630–780 Ma, suggesting oceanic-arc rocks as provenance. Abd El-Rahman et al. (2019) estimated the maximum age for the base of the Wadi Hammamat basin sediments at 625 Ma. The zircons in the Hammamat conglomerate showed peak ages of 840, 780, 725, 680 and 625 Ma. They noted that the shortage of Dokhan age zircons in some CED basins matches the rarity of the Dokhan in their surrounding areas. This not what would be expected if the rarity of Dokhan was simply due to its erosional removal and incorporation into Hammamat. Recently, Kiyokawa et al. (2020) dated the El Dabbah granitoid that intrudes the base of the Kareim sequence to 638 ± 3 Ma. The youngest zircons in this granitoid were 620–625 Ma. In a sandstone from the lower parts of the basin the youngest zircons dated to 640 Ma, with zircon age peaks at 790, 680 and 650 Ma. Again this does not favour Dokhan as a provenance for the early sedimentation of the basin. In the NED, Wilde and Youssef (2000) dated Gebel Dokhan volcanics bordering Um Tawaf basin to 602– 593 Ma (±10 Ma). In the same basin Wilde and Youssef (2002) found zircons to be younger at the base of the Hammamat (585 ± 13 Ma) than at the top (45 My older) consistent with gradual erosion of a nearby Dokhan volcanic pile, and onset of Hammamat deposition at 600 Ma or younger. They also found a large population of older zircons (peaks at 680 and 640 Ma). Eliwa et al. (2010) obtained U– Pb zircon ages on Dokhan above Hammamat at Gebel El Urf dated at 616 Ma. The radiometric data basically confirms the different age range of CED and NED basins and has implications for the provenance of the Hammamat. These points will be discussed ahead in the context of the evolution of the basins.

19.15

Discussion

19.15.1 Relations of the Hammamat Basins to Each Other The above sections illustrate diversity among the Hammamat basins in terms of size, shape, structure, metamorphism, stratigraphy, relations to other major rock units (Dokhan, Younger Granites, felsites), zircon population geochronology and geochemistry. Features in common include limited duration (all formed and were closed within a 55 My time range, 630–575 Ma), sedimentology, facies, terrestrial alluvial fan, braided stream and lake sedimentary environments, and petrography. Much of the information on the Hammamat has been collected from the larger Hammamat basins (Wadi Hammamat, El Qash, Zeidun, Arak and Kareim)

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which are concentrated into a limited area of the CED. With individual exceptions, there is a shortage of data on the remaining smaller basins of the NED to SED. Considering all of the basin properties together, a grouping of these larger basins appears justified on the basis of the following shared properties: (1) Their location and shape are generally equant (we will discuss the elongation of Wadi Hammamat basin ahead) (2) Their geochemistry (as much as we know of it) defines a group separate from most other smaller basins in terms of weathering indices, and major and minor chemical trends (3) These trends appear to strictly limit the contribution of Dokhan volcanics in the provenance of the basins (4) In accord with (3), there are few isolated Dokhan outcrops existing near the margins of these basins (5) The geochronology of zircon populations is more consistent with older metavolcanics, ophiolites and associated plutons in the provenance of the basins (6) They lie close to each other and all are located within an area affected by the EDSZ, while most other smaller basins lie beyond the EDSZ. All show some evidence for the lower stratigraphic levels of the basins being affected by EDSZ-related shearing, and the sheared zones show greenschist facies metamorphism accompanying significant degrees of elongation of basal conglomerate pebbles. The smaller Hammamat basins include some examples with similarities to the large basins in all but size and shape, e.g. the Um Seleimat and Meesar basins, and possibly the Atawi, Endiya and El Miyah basins, which lie within or near the EDSZ. Other small basins have a close association with Dokhan Volcanics and Younger Granites, which dominated the provenance of the basins, e.g. Um Tawat, Kharaga, Hamid, El Urf, Wadi Dib, Wadi Bali, Fatira, Queih, Sheqeili, Semna, Wadi Shait and Wadi Ranga basins.

19.15.2 Information Based on Basin Size, Width and Thickness of Sedimentary Fill A critical aspect of the Hammamat basins that stands in the way of understanding their tectonic setting is the question of how the basins formed and evolved in size and depth. In the absence of seismic geophysical data on these basins to reconstruct their history by back-stripping methods, we are left with less direct options for constraining models of basin histories. Plotting basin parameters one against the other may provide some clues to basin forming mechanisms. Plots comparing basin fill thickness (stratigraphic thickness), basin area and basin width are shown in Fig. 19.13a–d for 28

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Fig. 19.13 Plot of basin parameters (width, length, “depth”, i.e. stratigraphic fill thickness) for 28 NED and CED Hammamat basins (listed below). a Plot of basin area/width versus basin width (effectively a basin width versus length plot), gives some idea of basin shape. Lines of constant shape aspect ratio are given. The larger basins are more equant in shape while the smaller basins are more elongate. b Plot of basin area versus basin width also allows comparison of basin shapes. The blue curve represents the equation area = width2, showing the approximately equant shape of the larger Hammamat basins. The smaller basins are better approximated by the red line representing the equation area = 8  width emphasizing the elongate shape of these basins. c and d Plots of basin fill (i.e. “depth” or stratigraphic thickness) versus basin area and width (the latter two parameters effectively tracking basin growth). These plots show different growth—depth trends between the small and large Hammamat basins, represented by the curved red correlation lines. Explanation for these trends is given in the text. The orange lines in (d) represent growth—depth trends for grabens with differently dipping boundary normal faults, according to the graben model shown in Fig. 19.14. The small number of basins plotting along the low-dipping blue correlation line in (c) and (d) are also discussed in the text. The 28 basins plotted include: Abu Gheryan Abu Sheqeili (NE), Abu Sheqeili (SW), Andiya, Arak, Atawi, W. Bali, W. Dib, El Hida, El Miyah, El Qash, El Urf, Esh El Mellaha, Fatira, Hamid, W. Hammamat, Igla, Kareim, Kharaza, Meesar, Queih, Ranga, Riseis, Semna, W. Shait, Um Seleimat, Um Tawat, Zeidun

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Fig. 19.14 a and b explain the geometric basis of formulae relating basin parameters (initial width w, final width W, basin depth T, boundary fault dip a and magnitude of slip s on boundary faults) for the simple graben model explained in the text. c, d and e show the progressive footwall migration away from the basin centre, representing one model to explain the slow rate of basin subsidence relative to growth in basin width for the larger Hammamat basins. Faults are numbered in order of age 1 = oldest, 3 = youngest active

Hammamat basins from the NED, CED and SED (listed in Fig. 19.13 caption). The errors in the data are modest; however, the assumptions in the approach cannot be tested at present. These include the assumption that the basin areas and widths are representative of the original dimensions, and that the preserved stratigraphic thickness is also not too different to the original basin fill. Figure 19.13a plots width of the basins against area/width (effectively basin “length”). Lines of constant aspect ratio (values of basin length divided by basin width) radiate from the origin, with larger basins showing equant geometry (plotting close to an aspect ratio of 1). The smaller narrower basins have much higher aspect ratios of 4 or more. This demonstrates a trend for the smaller Hammamat basins to be markedly elongate in shape. The plot of basin width versus basin area is presented in Fig. 19.13b. For the larger basins the data fit a simple parabolic curve with equation: area = width2, but for the smaller basins the data is better approximated by a linear function: area = 8  width. As with Fig. 19.13a, b shows that smaller basins have larger aspect

ratios. The basin stratigraphic thickness is plotted against basin area in Fig. 19.13c. The data can be fitted to a smooth curve that rises steeply from the origin in order to fit the data for the small basins, then the curve gradient declines for the large basin data. This change in correlation line gradient demonstrates that the small basins subsided rapidly as they modestly grew in size, whereas the large basins experienced relatively slower subsidence as they grew. A similar pattern is seen in Fig. 19.13d, which shows stratigraphic thickness plotted against basin width. The property of rapid subsidence compared to basin width or basin area growth is typical of a graben with steep boundary faults, as shown in Fig. 19.14a, b. This is the subsidence mechanism we will work with for the smaller basins and will quantify in the following section. For the larger Hammamat basins, the rapid growth in area of the basin relative to its depth may be explained by a number of mechanisms including sag basin subsidence, and fault-bound basin with footwall migration away from or towards the basin centre (Fig. 19.14c–e).

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There is a third data trend seen on the basin fill versus basin area and width plots (Fig. 19.13c, d) where several basins (Fatira, Um Tawat, El Urf, Abu Gheryan and Igla) show basin widths 625–605 Ma) than the smaller (NED) basins (595–575 Ma). Plotting one basin dimension against another (basin area or width against basin fill thickness) shows different trends for the smaller and larger basins. The smaller basins show plots expected for grabens with 50° dipping boundary faults and subsidence histories yielding relatively faster basin deepening than widening. By contrast the larger basins show significantly slower rate of deepening compared to areal growth rates that may indicate migration of fault boundaries towards or away from the basin centre or sag basin formation. Activity of these faults may be the origin of intraformational unconformities, known to exist in these basins. The latter style of basin growth is similar to that of the supra-detachment basins (basins lying above a detachment fault, such as the EDSZ). Future studies of Hammamat basins may benefit from application of remote sensing, zircon population geochronology and chemostratigraphy. One of the most urgent properties of these basins to explore is their three-dimensional shape, for which geophysical data will be needed.

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References Abd El-Rahman Y, Polat A, Fryer BJ, Dilek Y, El-Sharkawy M, Sakran S (2010) The provenance and tectonic setting of the Neoproterozoic Um Hassa Greywacke Member, Wadi Hammamat area, Egypt: evidence from petrography and geochemistry. J Afr Earth Sci 58:185–196 Abd El-Rahman Y, Abu Anbar M, Li X-H, Li J, Ling X-X, Wu L-G, Masoud AE (2019) The evolution of the Arabian-Nubian Shield and survival of its zircon U–Pb–Hf–O isotopic signature: a tale from the Um Had Conglomerate, central Eastern Desert, Egypt. Precam Res 320:46–62 Abd El-Wahed MA (2004) Structural analysis of Wadi El Qash area, Central Eastern Desert, Egypt. Ann Geol Surv Egypt 27:79–108 Abd El-Wahed MA (2007) Late Pan-African tectonic evolution and strain determinations in the late Neoproterozoic molasse sediments, Eastern Desert, Egypt: evidence for post-Hammamat compression and transpression. Egypt J Geol 51:1–39 Abd El-Wahed MA (2010) The role of the Najd fault system in the tectonic evolution of the Hammamat molasse sediments, Eastern Desert, Egypt. Arab J Geosci 3:1–26 Abdeen MM, Abdelghaffar AA (2011) Syn- and post-accretionary structures in the Neoproterozoic Central Allaqi-Heiani suture zone, Southeastern Egypt. Precam Res 185:95–108 Abdeen MM, Greiling RO (2005) A quantitative structural study of Late Pan-African compressional deformation in the Central Eastern Desert (Egypt) during Gondwana assembly. Gond Res 8:457–471 Abdeen MM, Warr LN (1998) Late orogenic basin evolution, deformation and metamorphism in the Pan-African basement, Wadi Queih, Eastern Desert of Egypt. Forschungszentrum Jülich, Sci Ser Int Bureau 42:272 Abdeen MM, Dardir AA, Greiling RO (1997) Rain drop prints mud crack polygons and pebbles as strain markers in a Pan-African molasse basin, Wadi Queih, Arabian Desert, Egypt. Annals Geol Surv Egypt 20:185–191 Abdel Khalek ML, Hafez AMA (1983) Geology of Wadi Igla area. Int Basement Tect Assoc Publ 5:125–130 Abdelsalam MG, Stern RJ (1996) Sutures and shear zones of the Arabian-Nubian Shield. J Afr Earth Sci 23:289–310 Abu El Ela AM, El Bahariya GA (1998) The Hammamat molasse sediments between Wadi Um Esh and Wadi Muweilih, Qift-Quseir road region, Central Eastern Desert, Egypt. Egypt J Geol 41:5–35 Abu Sharib ASAA, Maurice AE, Abd El-Rahman YM, Sanislav IV, Schulz B, Bakhit BR (2019) Neoproterozoic arc sedimentation, metamorphism and collision: evidence from the northern tip of the Arabian-Nubian Shield and implication for the terminal collision between East and West Gondwana. Gond Res 66:13–42 Ahmed EA, Khudeir AA, El-Taky M (1989) Sedimentological and geochemical studies on the Hammamat sediments at Wadi Queh area, Central Eastern Desert, Egypt. Bull Fac Sci Assiut Uni 18:1– 20 Akaad MK (1957) Petrography, structure and time relations of the Igla formation, Eastern Desert. Egypt J Geol 1:93–105 Akaad MK (1972) A contemplation and assessment of the 1960–1961 classification of the rocks of the Central Eastern Desert. Ann Geol Surv Egypt 2:19–44 Akaad MK (1996) Rock succession of the basement. An autobiography and assessment. Geol Surv Egypt Paper No. 71, Cairo Akaad MK, El Ramly MF (1958) Seven new occurrences of the Igla formation in the Eastern Desert of Egypt. Geol Surv Cairo, Paper No. 2. 38 pp Akaad MK, Noweir AM (1964) A new occurrence of the Atud formation at Wadi El-Muweilih, Eastern Desert of Egypt. Bull Sci Tech Assiut 7:31–48

479 Akaad MK, Noweir AM (1969) Lithostratigraphy of the Hammamat-Um Seleimat District, Eastern Desert, Egypt. Nature 223:284–285 Akaad MK, Noweir A (1980) Geology and lithostratigraphy of the Arabian Desert Orogenic Belt of Egypt between latitude 25° 35′ and 26° 30′. Bull Inst Appl Geol King Abdulaziz Univ 3:127–135 Akaad MK, Noweir AM, Shazly AG, Sief El-Din SA (1979) Geology, petrography and condition of sedimentation of the Igla formation at Wadi El-Hammamat, Eastern Desert, Egypt. Delta J Sci 3:52–90 Akawy A, Zaky K (2008) Structural analysis of the tectonically inverted Kareim Basin, Eastern Desert, Egypt. N Jb Geol Palaont Abh 250:9–30 Ali KA, Stern RJ, Manton WI, Kimura J-I, Khamees HA (2009) Geochemistry, Nd isotopes and U–Pb SHRIMP dating of Neoproterozoic volcanic rocks from the Central Eastern Desert of Egypt: new insights into the  750 Ma crust-forming event. Precam Res 171:1–22 Ali-Bik MW, Hassan SA, Abou El Maaty MA, Abd El Rahim SH, Abayazeed SD, Abdel Wahab W (2018) The late Neoproterozoic Pan-African low-grade metamorphic ophiolitic and island-arc assemblages at Gebel Zabara area, Central Eastern Desert, Egypt: Petrogenesis and remote sensing—based geologic mapping. J Afr Earth Sci 144:17–40 Alsubhi MB (2012) Geological and structural studies on Jabal Daf-Jabal Abu Bakr area, Wadi Fatima, Kingdom of Saudi Arabia. M.Sc. Thesis, King Abdulaziz University, 125 pp Amin BEM (2002) Geological studies with application of remote sensing on the basement rocks between latitudes 25° 05′ and 25° 25′ N and Longitudes 33° 55′ and 34° 10′ E, Eastern Desert, Egypt. PhD Thesis, Al Azhar University, Egypt Andresen A, Abu El-Rus MA, Myhre PI, Boghdady GY, Corfu F (2009) U–Pb TIMS age constraints on the evolution of the Neoproterozoic Meatiq Gneiss Dome, Eastern Desert, Egypt. Int J Earth Sci 98:481–497 Andrew G (1939) The greywackes of the Eastern Desert of Egypt. Bull Inst, Egypte 21:153–190 Asran AMH, Azer MK, Aboazom AS (2005) petrological and geochemical investigations on Dokhan Volcanics and felsites in Wadi Sodmein, Central Eastern Desert, Egypt. Egypt J Geol 49:1–20 Assran ASM, El-Shayeb HM, Abouelnaga HSO, Rabie SI, El Qassas RA (2014) Airborne radiometric and magnetic investigations for Wadi El Tarfawy—Wadi El Dabbah area, Central Eastern Desert, Egypt. Egypt J Appl Geophys 13:55–70 Atalla RF (1978) Geology and petrography of the Hammamat conglomerate of Wadi Kareim, eastern Desert, Egypt. M.Sc. Thesis, Assiut University, 105 pp Azzaz SA, Kharbish S, El-Alfy HM (2015) Geological and geochemical investigations on Hammamat molasse sediments, G. Kharaza, Eastern Desert, Egypt. Acta Universitatis Matthiae Belii séria Environmentálne manažérstvo 17:24–34 Barthoux J (1922) Chronologie et description des roches ignées du Desert Arabique. Institut d’Egypte, Mémoires 5:262 Basahel AN, Groen C, Alwash MM (1984) Structural history of the Precambrian of western Wadi Fatima, in Pan-African crustal evolution of the Arabian-Nubian Shield, I.G.C.P. Project 164: King Abdulaziz University, Jeddah, Saudi Arabia. Fac Earth Sci Bull 6:422–434 Bezenjani RN, Pease V, Whitehouse MJ, Shalaby MH, Kadi KA, Kozdroj W (2014) Detrital zircon geochronology and provenance of the Neoproterozoic Hammamat group (Igla Basin), Egypt and the Thalbah group, NW Saudi Arabia: implications for regional collision tectonics. Precam Res 245:225–243 Bhatia MR, Crook KAW (1986) Trace element characteristics of graywackes and tectonic setting discrimination of sedimentary basins. Contrib Mineral Petrol 92:181–193

480 Breitkreuz C, Eliwa H, Khalaf I, El Gameel K, Bühler B, Sergeev S, Larionov A, Murata M (2010) Neoproterozoic SHRIMP U–Pb zircon ages of silica-rich Dokhan Volcanics in the North Eastern Desert, Egypt. Precam Res 182:163–174 Collettini C (2011) The mechanical paradox of low-angle normal faults: current understanding and open questions. Tectonophys 510:253– 268 Davies FB (1985) Geological map of the Al Wajh Quadrangle, Sheet 26B, Kingdom of Saudi Arabia. Saudi Arabian Deputy Ministry for Mineral Resources Geoscience Map GM 83, Scale 1:250,000, 27 pp de Rozière FM (1813) Description minéralogique de la Vallée de Qoceyr. In: Jomard E-F (ed) Description de l’Égypte, Histoire Naturelle (vol 2). Paris, pp 86–98 Delfour J (1970) Le groupe de J’balah une nouvelle unite du bouchlier Arabe. Bureau de Recherches Géologiques et Minières Bulletin 2:19–32 Dessouky OK, Dardier AM, Abdel Ghani IM (2018) Egyptian Hammamat molasse basins and their relations to arc collision stages: implications for radioactive elements mineralization potential. Geol J 54:1205–1222 Dickinson WR, Suczek CA (1979) Plate tectonics and sandstone compositions. Am Assoc Petrol Geol Bull 63:2164–2182 Dickinson WR, Beard LS, Brakenridge GR, Erjavec JL, Ferguson RC, Inman KF, Knepp RA, Lindberg FA, Ryberg PT (1983) Provenance of North American Phanerozoic sandstones in relation to tectonic setting. Geol Soc Am Bull 94:222–235 El Kalioubi B (1996) Provenance, tectonic setting and geochemical characteristics of the Hammamat molasse sediments around Um Had pluton, Eastern Desert, Egypt. M.E.R.C Ain Shams Univ Earth Sci Ser 10:75–88 El Kalioubi B, Fowler A, Abdelmalik K (2020) The metamorphism and deformation of the basement complex in Egypt. In: Hamimi Z, El-Barkooky A, Frias JM, Fritz H, Abd El-Rahman Y (eds) The geology of Egypt. Springer, pp 191–251 El Ramly MF (1972) A new geological map the basement rocks in the Eastern and South-Western Deserts of Egypt (Scale 1:1,000,000). Ann Geol Surv Egypt 2:1–18 El Ramly MF, Akaad MK (1960) Geological history and classification of the basement rocks of the Central Eastern Desert of Egypt. Geol Sur Cairo, Paper No. 9, 24 pp El Shazly EM (1977) The geology of the Egyptian region. In: Nairn AEM, Kanes WH, Stehli FG (eds) The Ocean basins and margins. Plenum Press, pp 379–444 El-Desoky HM (2013) Geochemistry and petrogenesis of Late Neoproterozoic Dokhan Volcanics at Wadi Abu Hamra area, Central Eastern Desert, Egypt. J Am Sci 9(5):212–235 El-Desoky HM, El Mezayen AM, Serag AE (2014) Petrology, geochemistry and petrogenetic aspects of Ediacaran Dokhan Volcanics at Wadi Zareib, Central Eastern Desert, Egypt. Stand Glob J Geol Explor Res 1:053–073 El-Gaby S (1983) Architecture of the Egyptian basement complex. Int Basement Tect Assoc Publ 5:1–8 El-Gaby S (1994) Geologic and tectonic framework of the Pan-African orogenic belt in Egypt. In: 2nd International conference geology of the Arab world, Cairo University, pp 3–17 El-Gaby S (2007) Integrated classification and evolution of the Neoproterozoic Pan-African belt in Egypt. In: 5th International conference on the geology of Africa, vol 1, pp V143–V154 El-Gaby S, El-Nady O, Khudeir A (1984) Tectonic evolution of the basement complex in the Central Eastern Desert of Egypt. Geol Rundsch 73:1019–1036 El-Gaby S, List FK, Tehrani R (1988) Geology, evolution and metallogenesis of the Pan-African belt in Egypt. In: El-Gaby S, Greiling RO (eds) The Pan-African belt of NE Africa and adjacent

A.-R. Fowler and Z. Hamimi areas (Earth evolution sciences). Vieweg & Sohn, Wiesbaden, pp 17–68 El-Gaby S, List FK, Tehrani R (1990) The basement complex of the Eastern Desert and Sinai. In: Rushdi S (ed) The geology of Egypt. Balkema, Rotterdam, pp 175–184 El-Gaby S, Khudeir AA, Abd El-Tawab M, Atalla RF (1991) The metamorphosed volcanosedimentary succession of Wadi Kid, southeastern Sinai, Egypt. Ann Geol Surv Egypt 17:19–35 El-Gaby S, Khalaf IM, Eliwa HA, El-Miligy A, Gomaa RM (2002) Volcanosedimentary successions of the Wadi Sa’al-Wadi Zaghra area, southeastern Sinai, Egypt. In: 6th International conference on the geology of the Arab world, Cairo University, pp 25–44 Eliwa H, Breitkreuz C, Khalaf I, El Gameel K (2010) Depositional styles of early Ediacaran terrestrial volcano-sedimentary succession in Gebel El Urf area, North Eastern Desert, Egypt. J Afr Earth Sci 57:328–344 Eliwa HA, Breitkreuz C, Murata M, Khalaf IM, Bühler B, Itaya T, Takahashi T, Hirahara Y, Miyazaki T, Kimura J-I, Shibata T, Koshi Y, Kato Y, Ozawa H, Daas MA, El Gameel K (2014a) SIMS zircon U–Pb and mica K–Ar geochronology, and Sr–Nd isotope geochemistry of Neoproterozoic granitoids and their bearing on the evolution of the north Eastern Desert, Egypt. Gond Res 25:1570– 1598 Eliwa HA, El-Bialy MZ, Murata M (2014b) Edicaran post-collisional volcanism in the Arabian-Nubian Shield: the high-K calc-alkaline Dokhan Volcanics of Gabal Samr El-Qaa (592 ± 5 Ma), North Eastern Desert, Egypt. Precam Res 246:180–207 El-Kammar AM, Salman AE, Shalaby MH, Mahdy AI (2001) Geochemical and genetical constraints on rare metals mineralization at the central Eastern Desert of Egypt. Geochem J 35:117–135 El-Shazly SM, Hegazy HA (2000) Geochemistry and petrogenesis of a Late Proterozoic volcanic sequencein the Magal Gebriel area, South Eastern Desert, Egypt. Qatar Univ Sci J 20:181–195 El-Tokhi M, Alaabed S, Amin BEM (2010) Precambrian metavolcanics of Um Anab North Eastern Desert, Egypt, geochemistry and tectonic environment studies. Eur J Sci Res 3:507–524 Fedo CM, Nesbitt HW, Young GM (1995) Unraveling the effects of potassium metasomatism in sedimentary rocks and paleosols, with implications for paleoweathering conditions and provenance. Geology 23:921–924 Fielding L, Najman Y, Millar I, Butterworth P, Ando S, Padoan M, Barfod D, Kneller B (2017) A detrital zircon record of the Nile River and its catchment. J Geol Soc 174:301–317 Fowler A, Abdeen MM (2014) Reappraisal of strain estimations and measurement methods in the Hammamat Group sediments: comparison of primary and secondary grain fabrics with new data from Wadi Zeidun and Wadi Arak basins, CED, Egypt. Precam Res 247:1–32 Fowler A, El Kalioubi B (2004) Gravitational collapse origin of shear zones, foliations and linear structures in the Neoproterozoic cover nappes, Eastern Desert, Egypt. J Afr Earth Sci 38:23–40 Fowler A, Hamimi Z (2020) Structural and tectonic framework of Neoproterozoic basement of Egypt: from gneiss domes to transpression belts. In: Hamimi Z, El-Barkooky A, Frias JM, Fritz H, Abd El-Rahman Y (eds) The geology of Egypt. Springer, pp 81–130 Fowler TJ, Osman AF (2001) Gneiss-cored interference dome associated with two phases of late Pan-African thrusting in the Central Eastern Desert, Egypt. Precam Res 108:17–43 Fowler A, Osman AF (2013) Sedimentation and inversion history of three molasse basins of the western Central Eastern Desert of Egypt: Implications for the tectonic significance of Hammamat basins. Gond Res 23:1511–1534 Fowler A, Baghdady A, Abdelmalik K, Gad A (2020) Remote sensing-guided stratigraphic dissection of an Ediacaran terrestrial

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Post-amalgamation Depositional Basins in the Arabian-Nubian …

molasse basin (Kareim basin, Egypt), with implications for sedimentary evolution. Precam Res 338(2020):105589 Fowler A, Hamimi Z (2021) A review and synthesis of strain studies in the Egyptian Nubian Shield: materials, methods, problems and prospects. Precam Res 355 https://doi.org/10.1016/j.precamres. 2020.106072 Francis MH (1972) Geology of the basement complex in the North Eastern Desert between latitudes 27o 30′ and 28o 00′ N. Ann Geol Survey Egypt 2:161–180 Friedmann SJ, Burbank DW (1995) Rift basins and supradetachment basins: intracontinental extensional end-members. Basin Res 7:109– 127 Fritz H, Messner M (1999) Intramontane basin formation during oblique convergence in the Eastern Desert of Egypt: magmatically versus tectonically induced subsidence. Tectonophys 315:145–162 Fritz H, Wallbrecher E, Khudier AA, Abu El Ela F, Dallmeyer RD (1996) Formation of Neoproterozoic metamorphic core complexes during oblique convergence, Eastern Desert, Egypt. J Afr Earth Sci 23:311–329 Fritz H, Dalmeyer DR, Wallbrecher E, Loizenbauer J, Hoinkes G, Neumayr P, Khudeir AA (2002) Neoproterozoic tectonothermal evolution of the Central Eastern Desert, Egypt: a slow velocity tectonic process of core complex exhumation. J Afr Earth Sci 34:543–576 Gaschnig RM, Rudnick RL, McDonough WF, Kaufman AJ, Valley JW, Hu Z, Gao S, Beck ML (2016) Compositional evolution of the upper continental crust through time, as constrained by ancient glacial diamictites. Geochim Cosmochim Acta 186:316–343 Gheith BM, Nigm AA (2009) The structural framework of Gabal El Hassnawia area, Central Eastern Desert, Egypt, as deduced from airborne magnetic data. NRIAG J Geophys Spec Iss 2008:123–142 Ghoneim MF, Noweir MA, Abu-Alam TS (2015) Magmatic evolution of the area around Wadi Kariem, Central Eastern Desert, Egypt. Arab J Geosci 8:9221–9236 Gomaa MM (2013) Forward and inverse modelling of the electrical properties of magnetite intruded by magma, Egypt. Geophys J Int 194:1527–1540 Grothaus B, Eppler D, Ehrlich R (1979) Depositional environment and structural implication of the Hammamat formation, Egypt. Ann Geol Surv Egypt 9:564–590 Hamimi Z (2000) Quantitative strain analyses of some late-tectonic Hammamat sediments, Eastern Desert, Egypt. Egypt J Geol 44:85– 107 Hamimi Z, El Kazzaz Y (2000) Fold geometry and folding mechanism in Um Krush klippe, South Eastern Desert, Egypt. 5th Internat Conf Geology of the Arab World, Cairo Uni. pp 605–618 Hamimi Z, Matsah M, El-Shafei M, El-Fakharani A, Shujoon A, Al-Gabali M (2012a) Wadi fatima thin-skinned foreland FAT belt: a post amalgamation marine basin in the Arabian shield. Open J Geol 2:271–293 Hamimi Z, El-Shafei M, Kattu G (2012b) Transpressional regime in southern Arabian shield: insights from Wadi Yiba area, Saudi Arabia. Miner Petrol 107:849–860 Harnois L (1988) The CIW index: a new chemical index of weathering. Sediment Geol 55:319–322 Harrell JA, Brown VM (1992) The world’s oldest surviving geological map: the 1150 BC. Turin Papyrus from Egypt. J Geol 100:3–18 Hassan MA, Hashad AH (1990) Precambrian of Egypt. In: Said R (ed) The geology of Egypt. Balkema, Rotterdam, pp 201–245 Holail HM, Moghazi AM (1998) Provenance, tectonic setting and geochemistry of greywackes and siltstones of the Late Precambrian Hammamat Group, Egypt. Sed Geol 116:227–250 Hume WF (1934) Geology of Egypt. Volume II: the fundamental pre-cambrian rocks of egypt and the sudan: their distribution, age

481 and character, Part I: metamorphic rocks. In: Geological survey of egypt, pp 1–300 Johnson PR (2003) Post-amalgamation basins of the NE Arabian Shield and implications for Neoproterozoic tectonism in the northern East African Orogen. Precam Res 123:321–337 Johnson PR, Woldehaimanot B (2003) Development of the Arabian– Nubian shield: perspectives on accretion and deformation in the northern East African Orogen and the assembly of Gondwana. In: Yoshida M, Windley BF, Dasgupta S (eds) Proterozoic East Gondwana: supercontinent assembly and breakup; Geol Soc Lond Spec Publ 206:290–325 Johnson PR, Andresen A, Collins AS, Fowler AR, Fritz H, Ghebreab W, Kusky T, Stern RJ (2011) Late Cryogenian-Ediacaran history of the Arabian-Nubian shield: a review of depositional, plutonic, structural, and tectonic events in the closing stages of the northern East African Orogen. J Afr Earth Sci 61:167–232 Johnson PR, Halverson GP, Kusky TM, Stern RJ, Pease V (2013) Volcanosedimentary basins in the Arabian-Nubian shield: markers of repeated exhumation and denudation in a Neoproterozoic Accretionary Orogen. Geosciences 3:389–445 Kamal El-Din GM, Asran AM (1995) Structural elements and metamorphism controlling the late Pan-African Hammamat sediments at the up-stream of Wadi El-Miyah, Central Eastern Desert of Egypt. In: Proceedings of international conference 30 years cooperation on the geology of egypt and related sciences (April 1993), pp 213–228 Khalaf EA (2010) Stratigraphy, facies architecture, and palaeoenvironment of Neoproterozoic volcanics and volcaniclastic deposits in Fatira area, Central Eastern Desert, Egypt. J Afr Earth Sci 58:405– 426 Khalaf EA (2012) Mechanisms of volcaniclastic aggradation in fluvial systems influenced by explosive volcanism: an example from Neoproterozoic Hammamat Group, Wadi Queih area, Central Eastern Desert, Egypt. J Afr Earth Sci 68:44–66 Khalaf EA (2013) Syn-eruptive/inter-eruptive relationships in Late Neoproterozoic volcano-sedimentary deposits of the Hamid area, North Eastern Desert, Egypt. Bull Volcanol 75(693) Khalil AE, El-Desoky HM, Shahin TM, Abdelwahab W (2018) Late Cryogenian arc-related volcaniclastic metasediment successions at Wadi Hammuda, Central Eastern Desert, Egypt: geology and geochemistry. Arab J Geosci 11:74. https://doi.org/10.1007/s12517017-3370-9 Khameis AA, Nigm AA (2010) Magnetic interpretation of north Gebel El Shallul area, central Eastern Desert, Egypt. Egypt J Rem Sens Space Sci 13:53–62 Khudeir AA, Ahmed EAS (1996) Tectono-sedimentary framework of the molasse sediments, Wadi Abu Shiqieli, Egypt. Egypt J Geol 40:197–226 Kiyokawa S, Suzuki T, Horie K, Takehara M, El-Dokouny HA, Dawoud M, Abuelhasan MM (2020) Tectonic and sedimentary history of the Neoproterozoic metavolcanic–volcaniclastic rocks of the El-Dabbah group, Central Eastern Desert, Egypt. J Afr Earth Sci 165: Makovicky E, Frei R, Bailey JC, Karup-Moller S (2016) Imperial porphyry from Gebel Abu Dokhan, the Red Sea Mountains, Egypt: Part II geochemistry. N Jb Miner Abh 193:29–44 Makroum FM, Assaf HS, Saleh GM (2001) Tectonic evolution of the basement complex in the Wadi Arak area, central Eastern Desert, Egypt. Egypt J Geol 45:51–64 Marsaglia KM, Ingersoll RV (1992) Compositional trend in arc-related, deep-marine sand and sandstone: a reassessment of magmatic-arc provenance. Geol Soc Am Bull 104:1637–1649 Massey KW (1984) Rubidium-strontium geochronology and petrography of the Hammamat formation in the Northeastern Desert of Egypt. M.Sc. Thesis, Univ. Texas, Dallas, 75 pp

482 Matsah MI, Kusky TM (1999) Sedimentary facies of the Neoproterozoic Al-Jifn basin, NE Arabian shield: relationships to the Halaban-Zarghat (Najd) faults system and the closure of the Mozambique Ocean. In: Greiling R (ed), Pan-African of northern Africa–Arabia. Forschungszentrum Jülich. Geologisch Palaeontoloigsches Institut, Ruprecht-Karls-Universitaet, Heidelberg, Proceedings of a Workshop, 22–23 Oct, 1998 Matsah MI, Kusky TM (2001) Analysis of Landsat ratio imagery of the Halaban Zarghat fault and related Jifn basin, NE Arabian shield: implications for the kinematic history of the Najd fault. Gond Res 4:182 Maurice AE, Basta FF, Khiamy AA (2012) Neoproterozoic nascent island arc volcanism from the Nubian Shield of Egypt: Magma genesis and generation of continental crust in intra-oceanic arcs. Lithos 132–133:1–20 Messner M (1996) The Hammamat formation: preliminary comparison of the Wadi Hammamat—Wadi El Qash and Wadi Kareim basins, Pan-African molasse basins, Eastern Desert, Egypt. In: Proceedings of the Egyptian geological survey centennial conference, Cairo, pp 529–545 Moghazi AM (2003) Geochemistry and petrogenesis of a high-K calc-alkaline Dokhan Volcanic suite, South Safaga area, Egypt: the role of late Neoproterozoic crustal extension. Precam Res 125:161– 178 Mohamed FH, Moghazi AM, Hassanen MA (2000) Geochemistry, petrogenesis and tectonic setting of late Neoproterozoic Dokhan-type volcanic rocks in the Fatira area, eastern Egypt. Int J Earth Sci 88:764–777 Moufti MRH (2001) Age, geochemistry, and origin of peraluminous A-type granitoids of the Ablah-Shuwas pluton, Ablah Graben Arabian shield. Acta Mineralogica-Petrographica, Szeged 42:5–20 Nebert K, Alshaibi AA, Awlia M, Bounny I, Nawab ZA, Sharief OH, Sherbini OA, Yeslam AH (1974) Geology of the area north of Wadi Fatima, Kingdom of Saudi Arabia. Center Appl Geol Bull Jeddah 1:31 Nesbitt HW, Young GM (1982) Early Proterozoic climates and plate motions inferred from major element chemistry of lutites. Nature 299:715–717 Noweir AM (1968) Geology of the Hammamat-Umm Seleimat district, Eastern Desert. PhD Thesis, Faculty of Science, Assiut University (unpubl) Noweir AM, Sief El-Din SA (1979) Petrochemistry and geochemistry of the Dokhan volcanic—Igla formation association of the Hammamat district, Eastern Desert, Egypt. Delta J Sci 3:91–106 Noweir AHM, Ghoneim MF, El Anwar MA (1983) Geology and lithostratigraphy of the Precambrian around Wadi Umm Gheig, Eastern Desert, Egypt. Int Basement Tect Assoc Publ 5:77–83 Noweir MA, Ghoneim MF, Abu Alam TS (2005) Deformation phases and tectonic evolution of the Neoproterozoic rocks of Wadi Kareim area, Central Eastern Desert, Egypt. Ann Geol Surv Egypt 28:81– 109 Osman AF (1996) Structural geological and geochemical studies on the Pan-African basement rocks, Wadi Zeidun, Central Eastern Desert, Egypt (Scientific Series of the International Bureau), vol 39. Forschungszentrum Julich GmbH, 262 pp Osman AF, Greiling RO, Warr LN, Hilmy ME, Ragab AI, El Ramly MF (1993) Distinction of different syn- and late orogenic Pan-African sedimentary sequences by metamorphic grade and illite crystallinity (W. Zeidun, Central Eastern desert, Egypt). In: Thorweihe U, Schandelmeier H (eds) Geoscientific research in Northeast Africa. Balkema, Rotterdam, pp 21–25 Osman AF, El Kalioubi B, Gaafar AS, El Ramly MF (2001) Provenance, geochemistry and tectonic setting of Neoproterozoic molasse-type sediments, North Eastern Desert, Egypt. Ann Geol Surv Egypt 24:93–114

A.-R. Fowler and Z. Hamimi Ragab AI (1987) The Pan-African basement of the northern segment of the Eastern Desert of Egypt: a crustal evolution model and its implications on tectono-stratigraphy and granite types. M.E.R.C Ain Shams Univ Earth Sci Ser 1:1–18 Ragab AI, El Alfy Z (1996) Arc-arc collision model and its implications on a proposed classification of the Pan-African rocks of the Eastern Desert of Egypt. M.E.R.C Ain Shams Univ Earth Sci Ser 10:89–101 Ragab AI, El-Gharbawi RI, El-Alfy Z (1993) Pan-African tectonostratigraphic assemblages of Gabal Meatiq—Wadi Atalla area, Eastern Desert, Egypt: evidence for arc-arc suturing. M.E.R.C. Ain Shams Univ Earth Sci Ser 7:131–145 Ramsay JG (1967) Folding and fracturing of rocks. McGraw-Hill, New York, p 568 Rice AHN, Osman AF, Abdeen MM, Sadek MF, Ragab AI (1993a) Preliminary comparison of six late- to post-Pan-African molasse basins, E. Desert, Egypt. In: Thorweihe U, Schandelmeier H (eds) Geoscientific research in Northeast Africa. Balkema, Rotterdam, pp 41–45 Rice AHN, Sadek MF, Rashwan AA (1993b) Igneous and structural relations in the Pan-African Hammamat Group, Igla Basin, Egypt. In: Thorweihe U, Schandelmeier H (eds) Geoscientific research in Northeast Africa. Balkema, Rotterdam, pp 35–39 Ries AC, Shackleton RM, Graham RH, Fitches WR (1983) Pan-African structures, ophiolites and melanges in the east Desert of Egypt: a traverse at 26°N. J Geol Soc Lond 140:75–95 Roser BP, Korsch RJ (1988) Provenance signatures of sandstonemudstone suites determined using discriminant function analysis of major-element data. Chem Geol 67:119–139 Sabet AH (1972) On the stratigraphy of the basement rocks of Egypt. Ann Geol Surv Egypt II:79–101 Samuel MD, El-Sokkary A, Ali AM (1978) Contribution to the petrography, mineralogy and geochemistry of the Igla Red Beds, Eastern Desert, Egypt. Bull. N.R.C Cairo, Egypt 3:95–108 Schürmann HME (1966) The Precambrian along the Gulf of Suez and the Northern part of the Red Sea. EJ Brill, Leiden, p 404 Shalaby A, Stüwe K, Fritz H, Makroum F (2006) The El Mayah molasse basin in the Eastern Desert of Egypt. J Afr Earth Sci 45:1– 15 Soliman MA (1983) Degree of metamorphism of the Hammamat Group, Qift-Quseir Road, Central Eastern Desert, Egypt. Bull Fac Sci Assiut Univ 12:117–128 Stern RJ (1979) Late Precambrian ensimatic volcanism in the central eastern desert of Egypt. Ph.D. Thesis, University of California, San Diego, CA, 210 pp Stern RJ (1981) Petrogenesis and tectonic setting of late Precambrian ensimatic volcanic rocks, C.E.D., of Egypt. Precam Res 16:195– 230 Stern RJ (1994) Arc assembly and continental collision in the Neoproterozoic East African Orogen implications for the consolidation of Gondwanaland. Annual Rev Earth Planet Sci 22:319–351 Stern RJ (2018) Neoproterozoic formation and evolution of Eastern Desert continental crust—the importance of the infrastructuresuperstructure transition. J Afr Earth Sci 146:15–27 Stern RJ, Hedge CE (1985) Geochronologic and isotopic constraints on late Precambrian crustal evolution in the Eastern Desert of Egypt. Am J Sci 285:97–127 Stern RJ, Voegeli DA (1987) Geochemistry, geochronology, and petrogenesis of a Late Precambrian (590 Ma) composite dike from the North Eastern Desert of Egypt. Geol Rundsch 76:325–341 Stern RJ, Gottfried D, Hedge CE (1984) Late Precambrian rifting and crustal evolution in the Northeastern Desert of Egypt. Geology 12:168–172 Stern RJ, Sellers G, Gottfried D (1988) Bimodal dyke swarms in the North Eastern Desert of Egypt: significance for the origin of late

19

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Precambrian “A-type” granites in northern Afro-Arabia. In: El-Gaby S, Greiling RO (eds) The Pan-African belt of NE Africa and adjacent areas. Vieweg and Sohn, pp 147–179 Taylor SR, McLennan SM (1985) The continental crust: its composition and evolution. Blackwell, Oxford, p 312 Van Houten FB (1973) Origin of Red Beds—a review 1961–1972. Annu Rev Earth Planet Sci 1:39–61 Vetti VV, Fossen H (2012) Origin of contrasting Devonian supradetachment basin types in the Scandinavian Caledonides. Geology 40:571–574 Wallbrecher E, Fritz H, Khudeir AA, Farahad F (1993) Kinematics of Pan-African thrusting and extension in Egypt. In: Thorweihe U, Schandelmeier H (eds) Geoscientific research in northeast Africa. Balkema, Rotterdam, pp 27–30

483 Wilde SA, Youssef K (2000) Significance of SHRIMP U-Pb dating of the Imperial Porphyry and associated Dokhan Volcanics, Gebel Dokhan, North Eastern Desert. Egypt J Afr Earth Sci 31:403–413 Wilde SA, Youssef K (2002) A re-evaluation of the origin and setting of the Late Precambrian Hammamat Group based on SHRIMP U– Pb dating of detrital zircons from Gebel Umm Tawat, North Eastern Desert, Egypt. J Geol Soc Lond 159:595–604 Williams GD, Powell CM, Cooper MA (1989) Geometry and kinematics of inversion tectonics. Geol Soc Lond Spec Publ 44:3–15 Willis KM, Stern RJ, Clauer N (1988) Age and geochemistry of late precambrian sediments of the Hammamat series from the Northeastern desert of Egypt. Precam Res 42:173–187

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Volcanism During the Post-accretionary Stage of the Arabian–Nubian Shield Mokhles K. Azer, Paul D. Asimow, and Simon A. Wilde

Abstract

The Arabian–Nubian Shield (ANS), formed by collision between East and West Gondwana during the Pan-African orogeny, serves as an excellent example of the product of major crustal accretion processes. One of the most striking features of the ANS is the widespread abundance of post-collisional plutons and associated volcano-sedimentary sequences. The magmatic history of the post-accretionary stage of the ANS includes two major episodes, both Ediacaran in age: the earlier calc-alkaline phase (635–610 Ma) and the later alkaline/peralkaline phase (610–580 Ma). Each of these magmatic episodes includes plutonic rocks and their volcanic equivalents. Early calc-alkaline volcanism was emplaced during two cycles of eruption of medium- to high-K calc-alkaline volcanic rocks, each accompanied by deposition of immature clastic sediments that remain undeformed and unmetamorphosed. The volcanic rocks of the early episode include intermediate to felsic subaerial lava flows, tuffs and ignimbrites, as well as subvolcanic bodies associated with minor basalt. The mostly high-K calc-alkaline character and other traits previously interpreted to indicate arc magmatism may simply reflect remelting of earlier arc-related material from the pre-accretionary stage (850–740 Ma). Some of the more evolved early episode calc-alkaline volcanic sequences are transitional to alkaline A-type, but this is interpreted to reflect extensive fractionation of I-type parental magmas rather than a switch in the tectonic

regime. In addition, some of the early calc-alkaline volcanic rocks are adakitic in character, perhaps recording early manifestations of lithospheric delamination, including melting of the mafic lower crust. The late alkaline/peralkaline volcanic units were emplaced after the termination of the Pan-African Orogeny. They are represented by alkali rhyolite, comendite and pantellerite flows, with abundant ignimbrites and pyroclastic deposits. They were emplaced during a non-orogenic period associated with tensile stresses, block faulting and differential uplift. During the closing stages of this volcanic phase, large-scale caldera subsidence occurred and ring dykes were injected into the bordering fractures. There is some overlap in the emplacement of calc-alkaline and alkaline rocks between 610 and 590 Ma, and peralkaline magmatism also extended into the early Cambrian (*530 Ma). The overlap period of coeval calc-alkaline and alkaline volcanic series implies that, on a regional scale, two magma sources coexisted during a complex transitional phase in the evolution of the ANS. Keywords

20.1 M. K. Azer (&) Geological Sciences Department, National Research Centre, Cairo, Egypt P. D. Asimow Division of Geological and Planetary Sciences, California Institute of Technology, Pasadena, CA 91125, USA S. A. Wilde School of Earth and Planetary Sciences, Curtin University, Perth, WA 6845, Australia











Volcanic rocks Calk-alkaline sequences Alkaline/peralkaline sequences Dokhan Volcanics Rutig Volcanics Katharina Volcanics Volcano-sedimentary successions







Introduction

The Arabian–Nubian Shield (ANS) represents the largest tract of juvenile continental crust of Neoproterozoic age on Earth (Stern 2002; Patchett and Chase 2002; Hargrove et al. 2006). It extends from Jordan and southern Israel in the north to Eritrea and Ethiopia in the south and from Egypt in the west to Saudi Arabia in the east (Fig. 20.1). Prior to the opening of the Red Sea–Gulf of Aden rift, the Arabian and

© The Author(s), under exclusive license to Springer Nature Switzerland AG 2021 Z. Hamimi et al. (eds.), The Geology of the Arabian-Nubian Shield, Regional Geology Reviews, https://doi.org/10.1007/978-3-030-72995-0_20

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Fig. 20.1 Simplified map of the Arabian–Nubian Shield (after Johnson and Woldehaimanot 2003). Basement outcrops are white; Phanerozoic cover is shown in light yellow; structural trends are highlighted; ophiolitic rocks are shown in black, and gneissic rocks are stippled. The locations of some examples of the post-collisional volcanics in this work are indicated: (1) Dokhan Volcanics in the Eastern Desert of Egypt, (2) Rutig Volcanics in Sinai, (3) Ferani Volcanics in Sinai, (4) Murdama volcano-sedimentary succession in Saudi Arabia, (5) Bani Ghayy Group in Saudi Arabia, (6) Attala felsite in the Eastern Desert of Egypt and (7) Katharina Volcanics in Sinai

Nubian shields were conjoined as a continuous tract of juvenile Neoproterozoic crust. On a regional scale, the ANS is the northern continuation of the Mozambique belt, and together they have been referred to as the East African Orogen (EAO) (Stern 1994). This is one of the largest known orogenic systems on Earth and produced both a large volume and a large area of continental crust (Rino et al. 2008). The EAO marks one of Earth’s greatest collision zones, a global feature in space (about 6000 km long, where it is preserved in Africa and Antarctica) and in time (350 million years of evolution) (Stern 1994). The term Pan-African was first introduced by Kennedy in 1964 on the basis of an assessment of available Rb–Sr and K–Ar ages in Africa. It was defined as an “important and widespread tectonic and thermal event” which affected the

African continent and led to its “structural differentiation into cratons and orogenic areas” at 500 ± 100 Ma ago, during which a number of mobile belts formed, surrounding older cratons. The term Pan-African was then extended to the rest of Gondwana, although regional names were also proposed, such as Brasiliano for South America, Adelaidean for Australia and Beardmore for Antarctica (Kröner and Stern 2005). Subsequent studies on several Pan-African regions and the Arabian Shield have expanded this age range to 1200–500 Ma (Hashad 1980). Kröner (1984) favored a time span from 950 to 450 Ma, whereas Kröner and Stern (2005) narrowed the time span to 870–550 Ma. Globally, the Pan-African orogeny culminated in the formation of the supercontinent Gondwana during the late Neoproterozoic to Paleozoic.

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The ANS preserves the processes of major crustal accretion that accompanied collision between East and West Gondwana during the Pan-African orogeny (Stern 1994; Dilek and Ahmed 2003; Meert 2003). The ANS is a collage of Neoproterozoic juvenile arcs, younger sedimentary and volcanic basins, voluminous granitoid intrusions and enclaves of pre-Neoproterozoic crust that crop out in the western Arabian Plate and the northeastern African Plate (Kröner et al. 1991; Stern 1994; Kusky et al. 2003; Johnson and Woldehaimanot 2003; Reischmann and Kröner 1994; Abdallah et al. 2019). The stabilization of the ANS occurred before the development of an extensive peneplain in mid-Cambrian times (*520 Ma) and was exhumed in the Neogene as a consequence of Red Sea rifting and uplift. The latest tectonic event in the ANS (starting *30 Ma ago) was the formation of the Red Sea and related structures, including the Gulf of Suez and the Gulf of Aqaba. The northernmost segment of the ANS is characterized by an abundance of post-collisional plutons and associated volcano-sedimentary sequences (Bentor 1985; Stern and Hedge 1985; Bentor and Eyal 1987; Jarrar et al. 2003; Stein 2003; Khalil et al. 2018). Post-collisional associations in the ANS include calc-alkaline (ca. 635–610 Ma) and alkaline (ca. 610–580 Ma) plutonic and volcanic suites (Beyth et al. 1994; Jarrar et al. 2003; Ali et al. 2009a; Azer et al. 2010; Be’eri-Shlevin et al. 2009a, 2011). The post-accretionary volcanic units are a distinctive part of the ANS and represent a key component of the shield, providing important clues about its origin and mineralization. They include different phases of volcano-sedimentary successions that can be separated into early calc-alkaline volcanic rocks and late alkaline/peralkaline volcanic rocks, both of which are interbedded with immature sediments, including conglomerates, sandstones and siltstones. In contrast to continuing studies of post-accretionary volcanism in the Nubian Shield, those of the Arabian Shield have yet to be comprehensively surveyed. This work is a review of what is known about the post-accretionary volcanic sequences in the ANS. It focuses on the distribution, stratigraphy, field characteristics, geochemistry, ages, petrogenetic models and tectonic setting of these volcanic units, with the goal of tracking the main trends in chemical and tectonic change over time.

20.2

Geotectonic Evolution of the Arabian– Nubian Shield

The ANS is the best preserved and most widely exposed juvenile continental crust of Neoproterozoic age on Earth (Reymer and Schubert 1984; Stern 2002; Hargrove et al. 2006). The evolution of the ANS commenced with breakup of the Rodinia supercontinent and formation of the

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Mozambique Ocean and island arcs (Li et al. 2008; Johnson et al. 2011). During the closure of the Mozambique Ocean, the east and west parts of Gondwana began to collide and the island arcs amalgamated with the African cratons (Meert 2003; Johnson et al. 2011; Cox et al. 2012; Eyal et al. 2014a). The collisional stage was followed by calc-alkaline and alkaline magmatism during the post-collisional stage (Eyal et al. 2010, 2019; Farahat and Azer 2011; Azer et al. 2014; Khalil et al. 2018). Geochronological and isotopic data can be utilized to infer the evolution of the ANS during the Neoproterozoic East African Orogeny (870–550 Ma ago; Kröner and Stern 2005). These data have traditionally been interpreted to show a lack of involvement of pre-Neoproterozoic crust in its formation (Stein and Goldstein 1996; Stern 1994, 2002). Nevertheless, a large number of pre-870 Ma inherited and detrital zircons (0.9–3.0 Ga) have been reported within the ANS magmas (Calvez et al. 1985; Sultan et al. 1990; Kröner et al. 1992; Loizenbauer et al. 2001; Kennedy et al. 2004; Hargrove et al. 2006; Ali et al. 2009a; Be’eri-Shlevin et al. 2009a, 2011, and many others) and sediments (Dixon 1981; Wilde and Youssef 2002; Avigad et al. 2007). Also, Archean and Paleoproterozoic rocks of the East Sahara Craton, west of the Nile, have been recognized at Oweinat in SW Egypt (Harris et al. 1984), in scattered basement exposures (Sultan et al. 1994) and even close to the Eastern Desert at Wadi Halfa on the Nile, just north of the border with Sudan (Stern 1994). Thus, some authors (e.g., El-Gaby et al. 1988; Sultan et al. 1990; Khudeir et al. 2008) consider that extensive tracts of older pre-Neoproterozoic continental crust exist beneath the Pan-African rocks, at least in the Eastern Desert of Egypt. Other authors (e.g., Stern 2002; Liégeois and Stern 2010) stressed that the presence of abundant pre-Neoproterozoic zircons only indicates the presence of older zircons, not the presence of extensive tracts of older crust. Nevertheless, Stern et al. (2010a) stated that “it seems clear that *1 Ga crust can be found near the Sa’al–Zaghra region of the Sinai Peninsula.” Several attempts have been made to classify the Proterozoic rocks of the ANS. There is general consensus that the high-grade gneisses and migmatites occupy the lowermost position, which represents the substratum of all other rock units (El-Gaby et al. 1988; Abdel-Khalek et al. 1992; El-Gaby 2007; Li et al. 2018). Some of the high-grade gneisses in the ANS are late Neoproterozoic in age (e.g., Andresen et al. 2009; Liégeois and Stern 2010; Augland et al. 2012; Ali et al. 2012, 2015). Therefore, some authors have considered them as a part of the juvenile crust of the ANS and interpreted U–Pb zircon dates of pre-Pan-African age (Be’eri-Shlevin et al. 2012; Abu El-Enen and Whitehouse 2013) as inherited zircons, as evident in numerous late Neoproterozoic juvenile crustal exposures in the ANS (e.g., Ali et al. 2009a; Be’eri-Shlevin et al. 2009b; Breitkreuz et al.

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2010; Stern et al. 2010a; Be’eri-Shlevin et al. 2009c, 2011, 2012; Samuel et al. 2011; Ali et al. 2015). However, Azer and Asimow (2021) considered that Pan-African dates obtained for some gneisses and sheared granites, especially in the northwestern part of the ANS, may represent episodes of metamorphism rather than protolith ages. Other authors considered that the high-grade gneisses of the northwestern part of the ANS may represent an easterly extension of the East Sahara Craton that was remobilized, and then exposed in tectonic windows during the Neoproterozoic East African Orogeny (e.g., Sturchio et al. 1983; Habib et al. 1985; El-Gaby et al. 1990; Hassan and Hashad 1990; El-Gaby 1994; Khudeir et al. 2008). Many models have been proposed for the evolution of the ANS (e.g., Greenwood et al. 1976; El-Gaby 1983; Kröner 1985; Kröner et al. 1987; El-Gaby et al. 1988; Stern 1994; Abdelsalam and Stern 1996; Blasband et al. 2000; Johnson et al. 2011). Bentor (1985) proposed four phases for the evolution of the Precambrian basement rocks of the ANS. Phase I (oceanic ophiolites phase; *1000 Ma) includes peridotites, gabbros, sheeted dykes, pillow lavas and pelagic sediments formed in a deep-sea environment. Phase II (island arc phase; *950–650 Ma) includes a thick sequence of basic to intermediate volcanic rocks and their intrusive equivalents, as well as volcanogenic clastic sediments. Phase III (batholithic phase; *640–590 Ma) includes subaerial, medium- to high-K calc-alkaline volcanic units of andesite to rhyolite composition and their plutonic equivalents, as well as volcanogenic clastics. During phase III, the

Fig. 20.2 Schematic illustration of stages in the development of ANS, showing its setting in the supercontinental cycle bracketed by the breakup of Rodinia and the assembly of Gondwana (after Stern and Johnson 2010)

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ANS was cratonized. Unlike those of phases I and II, all rock units of Phase III are unmetamorphosed. Finally, phase IV (alkaline phase; *590–580 Ma) includes intraplate alkaline rocks, high-level granites and syenites, and their volcanic equivalents (Katharina Province). The rocks of this phase were emplaced during the anorogenic stage of ANS evolution. According to Johnson et al. (2011), the ANS formed during four major tectonomagmatic episodes between about 870 and 550 Ma (Fig. 20.2). Based on this model, the initial rifting of the supercontinent Rodinia was followed by seafloor spreading that opened the Mozambique Ocean, accompanied by formation and accretion of fringing arc and back-arc basins to form the juvenile material of the ANS. Terrane accretion resulted in the formation of arc–arc sutures. Closure of the Mozambique Ocean led to collision that formed the ANS from rifted block fragments of East and West Gondwana along north-trending arc-continent sutures. The continuation of convergence of East and West Gondwana resulted in crustal shortening in the ANS, localized along north-trending linear belts. The shortening deformation lasted from *600 to 550 Ma with the formation of northwest-trending sinistral and northeast-trending dextral strike-slip fault systems. Also, it was suggested that in the last phase, escape tectonics and orogenic collapse occurred. Several workers have recognized a major uplift and erosional phase between the end of accretion in the ANS (ca. 635–625 Ma; e.g., Stern 1994; Katz et al. 2004; Be’eri-Shlevin et al. 2009a) and the formation of a widespread

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Cambrian peneplain in the northernmost ANS (Bentor 1985; Garfunkel 1999; Avigad and Gvirtzman 2009). In their model of post-collisional lithospheric delamination, Avigad and Gvirtzman (2009) postulated that the ANS was uplifted to elevations of more than 3 km at 630 Ma, thus triggering exceptionally rapid erosional unroofing of a ca. 10-km-thick rock carapace by 600 Ma.

20.3

General Classification of the ANS Volcanic Rocks

Three major Neoproterozoic volcanic episodes have been recognized in the ANS, namely: ophiolitic metavolcanic units, island arc metavolcanic units and post-collisional volcanic units. Also, Phanerozoic volcanic rocks are recorded in the ANS, extruded onto the Neoproterozoic rocks and their sedimentary cover. In the following paragraphs, we will give a brief outline of all these events. Following this, the post-collisional (post-accretionary) volcanic rocks, the target of the present chapter, will be discussed in greater detail. The ophiolitic metavolcanic units represent the oldest volcanic phase in the shield, but are not common in the ANS. In the Eastern Dessert of Egypt, they have been informally named the “Older Metavolcanics” (OMV; Stern 1981). These units are composed of pillow lavas and sheeted dykes. Low-K tholeiitic basalts with well-preserved pillow structures are the most common rock type; rare andesites are associated locally with the pillow basalt (Zimmer et al. 1995; Stern et al. 2004; Gahlan et al. 2015). The sheeted dykes include metabasalt with rare metadolerite and meta-andesite (Zimmer et al. 1995). The ophiolitic metabasalts are considered to have been generated in either a mid-ocean ridge (Zimmer et al. 1995) or supra-subduction zone setting (Basta et al. 2011; Gahlan et al. 2015; Obeid et al. 2016). The island arc metavolcanic sequences (750–650 Ma) represent an important rock unit in the Pan-African of the ANS. They have been informally named the “Younger Metavolcanics” (YMV; Stern 1981) and have geochemical characteristics of island arc volcanic rocks (e.g., Stern 1981; Maurice et al. 2012; Abdel-Karim et al. 2019). Importantly, Ali et al. (2009b) and Andresen et al. (2009) concluded that the OMV and the YMV overlap in time. The island arc metavolcanic units in the south Eastern Desert of Egypt are also known as the “Shadli Metavolcanics,” which refers to their type locality near the Tomb of Sheikh Shadli, whereas in Sinai they are known as the Kid and Sa’al Metavolcanics (Shimron 1984, 1987; Furnes et al. 1985; Shimron et al. 1993; Be’eri-Shlevin et al. 2012; El-Bialy 2013; Eyal et al. 2014a; Ali-Bik et al. 2017). In Saudi Arabia, different names have been assigned to the island arc metavolcanic units in different localities, including the Mahd Group Metavolcanics, Murshid Metavolcanics, Urayfi Formation, Fashghah

489

Formation, Siqam Formation, Siham Group and Ablah Group (e.g., Genna et al. 1999; Kennedy et al. 2004, 2005; Doebrich et al. 2007). The island arc volcanic rocks were deposited contemporaneously with volcanogenic sedimentary sequences that are commonly layered and graded, indicating subaqueous deposition. They have been metamorphosed at low- to medium-grade, i.e., greenschist to amphibolite facies conditions (Farahat 2006; Maurice et al. 2012). The volcanic units cover the entire spectrum from low-K metabasalt and meta-andesite to medium-K metadacite and metarhyolite, of calc-alkaline affinity (Alaabed and Amin 2010; Bühler et al. 2014; Abdel-Karim et al. 2019). The associated metasediments comprise thick sequences of regionally metamorphosed layered volcaniclastic turbidite, including interbedded metagreywackes, metamudstones, metasiltstones, metaconglomerates and schists. Geochemically, the island arc metavolcanic rocks have a transitional character from tholeiite to calc-alkaline affinity (e.g., Khudeir et al. 1988; Obeid 2006; Maurice et al. 2012). The post-collisional volcanic sequences (post-accretionary volcanic rocks) in the ANS are represented by two suites: calc-alkaline and alkaline/peralkaline volcanic suites. These volcanic rocks are unmetamorphosed, except for a few cases where the early phase of the calc-alkaline volcanic units is weakly metamorphosed and deformed due to their proximity to later granitic intrusions (El-Gaby et al. 1991; Obeid and Azer 2015; Be’eri-Shlevin et al. 2011; Abdelfadil et al. 2018). The calc-alkaline volcanic rocks mainly comprise intermediate to felsic subaerial lava flows, tuffs and ignimbrites, as well as subvolcanic felsites. The Neoproterozoic alkaline volcanic rocks in the ANS comprise alkali rhyolite, comendite and pantellerite flows, with abundant ignimbrites and pyroclastic deposits. The type locality of these rocks is exposed at Gebel Saint Katharina in south-central Sinai, where they are known as the Katharina Volcanics. Phanerozoic volcanic rocks in the ANS occur in association with alkaline ring complexes, as well as forming local lava flows, sills, dykes and small stocks (Saleeb-Roufaiel et al. 1989; Baldridge et al. 1991; Hassan et al. 1997, 2001; Samuel et al. 1999, 2000, 2002). The Natash Volcanics are the most common Phanerozoic volcanic rocks in the Nubian Shield (Hubbard 1981; Mohamed 2001; Darwish and El-Tohamy 2014; Khalaf et al. 2018; Abu El-Rus et al. 2019), whereas the Harrat volcanic flows are the most common in the Arabian Shield (e.g., Moufti et al. 2012; Moufti and Németh 2013, 2016; Murcia et al. 2017; Downs et al. 2018). The Natash Volcanics are intercalated with Nubian sandstones and were dated as Late Cretaceous, with a Rb/Sr age of 104 ± 7 Ma (Hashad and El Reedy 1979), which is consistent with the 100 Ma K/Ar age reported by Abul Gadayel (1972). The Natash Volcanics are mainly of alkaline (sodium-dominated) composition and include alkali

490

M. K. Azer et al.

olivine basalt, hawaiite, mugearite and benmoreite, associated with felsic volcanic flows of trachytic to rhyolitic composition (Mohamed 2001). The Harrat volcanic flows represent the largest intraplate volcanic fields in the Arabian Shield (Shaw et al. 2003). Together they form one of the most voluminous alkali volcanic provinces on Earth, covering an area of  180,000 km2 (Coleman et al. 1983). They occur widely in the western Arabian plate, where continental intraplate volcanic centers extend for >3000 km south to north from Yemen through Saudi Arabia to Jordan, Syria and Turkey (Downs et al. 2018). More than 21 eruptions have occurred over the past  1500 years, the last of which occurred at Dhamar in northern Yemen in 1937 (Shaw et al. 2003; Downs et al. 2018).

20.4

Post-accretionary Volcanic Sequences in the ANS

One of the most striking features of the ANS is the abundance of post-collisional plutons and associated volcano-sedimentary sequences, whereas older rocks, now comprising parts of metamorphic complexes, ophiolites and island arc assemblages, are less well represented. Rocks of the post-collisional phase have U–Pb zircon ages ranging from ca. 635 to 580 Ma (Moussa et al. 2008; Breitkreuz et al. 2008; Ali et al. 2009a; Azer et al. 2010; Be’eri-Shlevin 2009a, 2011; Abu El-Enen et al. 2018). The magmatic history of the post-accretionary stage of the ANS includes two major magmatic episodes: the calc-alkaline phase (630– 610 Ma) and the alkaline phase (610–580 Ma) (Beyth et al. 1994; Bentor and Eyal 1987; Jarrar et al. 2003; Ali et al. 2009a; Be’eri-Shlevin et al. 2009a, 2011; Eyal et al. 2010). Both of these two magmatic episodes include plutonic rocks and their volcanic equivalents. Post-collisional volcano-sedimentary rock associations are widespread in the ANS, but their mode of occurrence, relationship with surrounding rocks, and the clast component and associated provenance patterns of the sedimentary rocks are variable. Two Ediacaran post-accretionary volcanosedimentary successions are recognized within the ANS after *630 Ma (Eyal et al. 2010; Be’eri-Shlevin et al. 2011), which indicates a changing geodynamic setting during the post-collisional collapse of the ANS. The earlier succession is represented by a high-K calc-alkaline volcanic sequence, whereas the later episode is characterized by an alkaline/ peralkaline volcanic sequence. These volcanic rocks postdate the volcanic arc and collisional stages and mark the latest Neoproterozoic magmatic events in the ANS. The associated sediments, where preserved, are immature and include conglomerates, sandstones and siltstones (El-Gaby et al. 2002; Moussa 2003a; Azer 2007; El-Bialy 2010; Eliwa et al. 2010;

Stern et al. 2010a; Samuel et al. 2011; Bezenjani et al. 2014). The stratigraphic position of these successions in relation to post-collisional batholiths is not fully resolved, but it is clear that most of the post-accretionary volcano-sedimentary successions were intruded by ca. 595–580 Ma A-type granites (e.g., El-Gaby et al. 2002; Samuel et al. 2001, 2011; El-Bialy 2010; Stern et al. 2010a). However, this provides only a minimum estimate for their formation age and some of these successions may be much older.

20.4.1 Early Calc-Alkaline Volcanic Sequences The early post-accretionary volcano-sedimentary sequences dominate in the northern part of the ANS. This stage begins with the extrusion of medium- to high-K calc-alkaline volcanic rocks and deposition of immature clastic sedimentary units that are undeformed and unmetamorphosed. The volcanic rocks include intermediate to felsic subaerial lava flows, tuffs and ignimbrites, as well as subvolcanic bodies and minor basaltic andesite. The end of the early post-accretionary stage at *610–600 Ma is marked by injection of numerous dyke swarms of different compositions (Bentor 1983; Eyal et al. 2019). In many cases, the time interval between the syn-collisional and early post-collisional volcanic stages was very short. Also, the timing of post-accretionary magmatism was not contemporaneous across the whole of the ANS, since Eyal et al. (2014a, b, 2019) observed that early post-collisional calc-alkaline magmatism in Israel predates early post-collisional calc-alkaline magmatism in central and southern Sinai by 20–30 My. In the Nubian Shield, early calc-alkaline volcanic sequences are widespread in the Eastern Desert and Sinai, where they are associated with Hammamat clastic sedimentary deposits. In the Eastern Desert of Egypt, they are known as the Dokhan Volcanics and dated at 630–590 Ma (Willis et al. 1988; Wilde and Youssef 2000, 2001, 2002; Breitkreuz et al. 2010; Bezenjani et al. 2014). In the Sinai Peninsula, there are several localities where high-K calc-alkaline volcanic units occur (e.g., Bentor 1985; Bentor and Eyal 1987; Basta 1997; El-Bialy 2010; Samuel et al. 2011; Be’eri-Shlevin et al. 2011; Abdelfadil et al. 2018; Samuel et al. 2019). Different names have been assigned to the Dokhan Volcanics in different localities in Sinai, including the Ferani, Rutig, Khashabi, Iqna Shar'a and Meknas Volcanics. The early calc-alkaline volcanic sequences in the Eastern Desert and Sinai are mainly high-K volcanics ranging from basaltic andesite to rhyolite in composition (Fig. 20.3a). They plot in the post-collisional field on the discrimination diagram of Pearce et al. (1984) (Fig. 20.3b). On the SiO2 versus FeOT/(FeOT + MgO) diagram, the published data for the post-accretion calc-alkaline

20

Volcanism During the Post-accretionary Stage …

491

G. El-Kharaza (Abdel-Rahman 1996)

b

1000

I

Um Sidra-Um Asmer area (Eliwa et al. (2006)

Syn-collisional Granites

Iqna Shar'a (Samuel et al. 2001)

Within-plate Granites

W. Rufaiyil (Azer and Farahat 2011)

R

Rutig (Be'eri-Shlevin et al. 2011)

100

b

G. Samr El-Qaa (Eliwa et al. 2014)

Post-collisional Granites

Ferani (Moussa 2003; Be'eri-Shlevin et al. 2011) G. Abu Had area (Obeid and Azer 2015) Khashabi (Abdelfadil. (2018)

10

W. Hamed (Maurice et al. 2018)

Volcanic Arc Granites

W. Zaghra ( Samuel et al. 2019) G. Um Guruf (Abuamarah et al. 2021)

5

I

Oceanic Ridge Granites

1

a

1

100

10

1000

Y+Nb 1.0

Shoshonitic

4

c

I

3

FeOT/(FeOT+MgO)

K2O (wt.%)

Ferroan rocks

High K

Medium K

2

1

Basalt

Basaltic andesite

Andesite

0.6

Rhyolite

Dacite

0.0

45

50

55

60

65

Magnesian rocks

0.4

0.2

Low K

0

0.8

70

SiO2 (wt.%)

75

45

50

55

60

65

70

75

SiO2(wt.%)

Fig. 20.3 a K2O versus SiO2 diagram (after Rickwood 1989). b Rb versus Y + Nb tectonic discrimination diagram (after Pearce et al. 1984) and c FeOT/(FeOT + MgO versus SiO2 diagram for some post-collisional calc-alkaline volcanics in the Eastern Desert and Sinai

volcanics in the Eastern Desert and Sinai are magnesian (Fig. 20.3c), which is characteristic of post-collisional rocks (Frost et al. 2001; Frost and Frost 2008). In NE Sudan, the Amaki volcano-sedimentary series consists of late Neoproterozoic molasse-type sediments associated with post-orogenic high-K volcanic rocks that are similar to the Hammamat Group and Dokhan Volcanics, respectively, in Egypt (Almond et al. 1984). In Saudi Arabia, the post-accretionary high-K calc-alkaline volcanic sequences are well represented in the Jibalah Group (*630 Ma). They consist of coarse-grained continental clastic sedimentary rocks interbedded with volcanic units (Hadley and Schmidt 1980; Johnson and Woldehaimanot 2003). Also, similar volcanic sequences are known from several localities in Saudi Arabia with different names, including the Murdama, Bani Ghayy, Jurdhawiyah, Ablah, Hadn, Siqam and Murshid Volcanics (Johnson et al. 2011; Kennedy et al. 2004, 2005). In Israel, the early post-accretionary volcanic stage is represented by the Gishron Formation, exposed in an E-W (3.5  0.5 km) outcrop between the El-Hura and Haneikiya plutons (Eyal et al. 2016, 2019). It is composed of conglomerate, shale and volcanic rocks, unconformably overlies the El-Hura quartz dioritic gneiss, and is intruded by the calc-alkaline Haneikiya granite. Eyal et al. (1980) correlate the Gishron Formation with the Ferani Group in south Sinai because both are intruded by calc-alkaline plutons,

indicating erosion and deposition before the calc-alkaline magmatism. The volcanic rocks of the Gishron Formation consist of tuffs (lapilli tuffs and crystal tuffs) of andesitic to dacitic composition.

20.4.2 Late Alkaline/Peralkaline Volcanic Sequences At the end of the post-collisional phase, the alkaline/peralkaline volcanic sequence erupted during a non-orogenic period extending from the late Ediacaran (610–590 Ma) into the early Cambrian (*540 Ma). Their geochemical characteristics indicate a within-plate tectonic setting (Fig. 20.4). These volcanic units represent a well-defined phase of magmatism in the northern part of the ANS, as they occur in Sinai, southern Negev (Israel) and southwest Jordan. The late alkaline volcanic units in the ANS are small in volume, but nonetheless can be found widely scattered over the entire shield (Bentor 1985; Mushkin et al. 1999). They comprise alkali rhyolite, comendite and pantellerite flows, with abundant ignimbrites and pyroclastic deposits. The type locality is at Gebel Saint Katharina in south-central Sinai, where they are known as the Katharina Volcanics (605–580 Ma; Katzir et al. 2007; Moreno et al. 2012, 2014; Eyal et al. 2014b; Azer et al.

492

Fig. 20.4 Rb versus Y + Nb tectonic discrimination diagram (after Pearce et al. 1984) for some post-collisional alkaline volcanics in the Eastern Desert and Sinai

2014). The Katharina Volcanics are associated with the Katharina ring complex, a typical ring complex of late Ediacaran age (605–580 Ma; Katzir et al. 2007; Moreno et al. 2012, 2014; Eyal et al. 2014b). They were emplaced under tensile stress conditions, manifested by block faulting and differential uplift (Bentor 1985; Samuel et al. 2007). During the closing stages of this volcanic phase, large-scale caldera subsidence occurred and ring dykes were injected into the bordering fractures. Correlated alkaline/peralkaline volcanic rocks, with different local names, are exposed in a number of localities in Sinai and Israel: in south-central Sinai (Gebel Ma’ain), around Taba–Nuweiba (Wadi El-Mahash, Wadi Khileifiya and Gebel El-Homra), and at Biq’at Hayareah (Sinai–Negev border). These alkaline volcanic units are unconformably overlain by early Paleozoic sandstones, including the Cambrian Araba Formation (El-Araby and Abdel-Motelib 1999), that mark the onset of stable platform conditions in the Taba–Nuweiba district of Sinai (Azer 2004). In the Eastern Desert, exact correlatives of the Katharina Volcanics have not been described, but alkali rhyolites, felsites and alkali microgranites have been recognized (Essawy 1972; El-Mahallawi 1999a). El-Mahallawi and Ahmed (1993) and El-Mahallawi (1999a) concluded that the Igla felsites are calc-alkaline (I-type felsites), whereas the Atalla felsites are alkaline (A-type felsites). Asran et al. (2005) indicated that the felsites of Wadi Sodmein are similar to the Atalla felsites. Abu El-Ela (2001) equated the Atalla felsites with the Katharina Volcanics and showed that they are composed mainly of comendite, pantellerite and rhyolitic tuffs, and thus represent the extrusive equivalents of alkaline to peralkaline granites. Their Rb/Sr age (588 Ma; Hassan 1998a) and geochemistry are both comparable to the Katharina volcanic series in south Sinai. In the Midyan terrane of northeastern Saudi Arabia, the equivalent Minaweh Formation contains cobbles of *600 Ma monzogranite and is intruded by *575 Ma

M. K. Azer et al.

alkaline granites of the Midyan suite (Clark 1985). Some A-type volcanic rocks in the Arabian Shield were described as alkaline–peralkaline bimodal volcanic deposits that occurred in pull-apart grabens of the Najd strike-slip shear zone (Roobel et al. 1983; Agar 1986). In southern Jordan, the post-accretionary alkaline volcanic series is represented by the Haiyala volcaniclastic formation and the Saramuj conglomerate, which are intruded by a *595 Ma monzogabbroic body (Jarrar et al. 1993). The late post-accretionary alkaline volcanic series in Israel is characterized by basaltic and andesitic lava flows and pyroclastic rocks of the Mapalim stratovolcano edifice that was followed by alkali rhyolite volcanism of the Ramat Yotam Caldera and the Amram–Neshef volcanic field (Eyal et al. 2019). The latter is characterized by abundant subvolcanic intrusions of porphyritic rhyolite, alkali rhyolite, quartz syenite, alkali quartz syenite and monzonite. This stage also includes the Timna and Yehoshafat plutons and many dyke suites.

20.5

Ages of Post-accretionary Volcanic Sequences in the ANS

The most acceptable available radiometric ages for the post-accretionary volcanic sequences in the ANS are listed in Table 20.1. Old geochronological methods applied to the post-accretionary volcanic rocks in the ANS, such as Rb–Sr whole rock, gave inaccurate and scattered ages (e.g., El-Ramly 1962; El-Shazly et al. 1973); therefore, we exclude most of these data from Table 20.1. On the other hand, modern geochronology techniques using zircon U–Pb have been applied by many authors and give more precise ages (Wilde and Youssef 2000, 2002; Johnson et al. 2011; Kennedy et al. 2004, 2005; Be’eri-Shlevin et al. 2009b, 2011, Samuel et al. 2011; Moghazi et al. 2012; Andresen et al. 2014; Abu El-Enen et al. 2018; Abd El-Rahman et al. 2019). The modern techniques applied to the post-accretionary volcanic rocks in the ANS are based mainly on SIMS (SHRIMP and CAMECA) secondary ion mass spectrometry and LA-ICP-MS (laser ablation inductively coupled plasma mass spectrometry). The available ages for both the post-accretionary calc-alkaline and alkaline volcanic units from the ANS have been mostly obtained from the Eastern Desert of Egypt (Stern 1979; Dixon 1979; Ries et al. 1983; Stern and Hedge 1985; Abdel-Rahman and Doig 1987; Wilde and Youssef 2000; Breitkreuz et al. 2008), Sinai (Halpren and Tristan 1981; Bielski 1982; Basta 1997; Be’eri-Shlevin et al. 2011), Israel (Bielski 1982; Be’eri-Shlevin 2008), Jordan (Jarrar et al. 1992) and Saudi Arabia (Johnson et al. 2011; Kennedy et al. 2004, 2005). The first systematic geochronological study of the post-accretionary calc-alkaline volcanic units in the Nubian Shield was that of Stern and Hedge (1985), who dated ten

20

Volcanism During the Post-accretionary Stage …

493

Table 20.1 Age determinations of post-accretionary volcanic units in the Arabian-Nubian Shield Locality

Material, rock type

Method

Age (Ma)

References

Saudi Arabia Ablah group

Zircon, rhyolite

U–Pb

613 ± 7

Johnson et al. (2011)

Hadn Formation

Zircon, felsic volcanic

U–Pb

614 ± 6 Ma

Kennedy et al. (2004, 2005)

Jurdhawiyah

Zircon, andesite–dacite

U–Pb

594 ± 16

Kennedy et al. (2004, 2005)

Jurdhawiyah

Zircon, andesite–dacite

U–Pb

612 ± 4

Kennedy et al. (2004, 2005)

Bani Ghayy group Murdama

Zircon, rhyolite

U–Pb

620 ± 5

Stacey and Agar (1985)

Whole rock, andesite and basalt

Rb–Sr

620 ± 95

Fleck et al. (1980)

Zircon, andesite

U–Pb

623 ± 6

Kennedy et al. (2004, 2005)

Zircon, rhyolite

620 ± 4

Zircon, rhyolite

607 ± 3

Eastern Desert Qena–Safaga

Whole rock, rhyolite and andesite

Rb–Sr

602 ± 13

Stern (1979)

W. Sodmein

Whole rock, volcanic rocks

Rb–Sr

616 ± 9

Ries et al. (1983)

W. Fatira

Whole rock, volcanic rocks

Rb–Sr

620 ± 12

Ressetar and Monrad (1983)

W. El Mahdaf

Whole rock, volcanic rocks

Rb–Sr

622 ± 6

Stern and Hedge (1985)

G. Nuqrah

Whole rock, felsic rocks

Rb–Sr

581 ± 7

G. El Kharaza

Whole rock, volcanic rocks

Rb–Sr

620 ± 16

G. Dokhan

Abdel-Rahman and Doig (1987)

Whole rock, andesites

Rb–Sr

592 ± 15

Stern and Hedge (1985)

Zircon, imperial porphyry

U–Pb

602 ± 9

Wilde and Youssef (2000)

Zircon, quartz andesite

U–Pb

593 ± 13

Wilde and Youssef (2000)

Zircon, ignimbrites

U–Pb

615 ± 4

Breitkreuz et al. (2008)

Zircon, imperial porphyry

U–Th–Pb

609–600

Abu El-Enen et al. (2018)

616 ± 5.4 Wadi Umm Sidri Wadi Abu Ma’amel Sinai W. Kid

W. Meknas Rutig Volcanics

Feirani Volcanics

Whole rock, rhyolite

Rb–Sr

595

Halpren and Tristan (1981)

Whole rock, rhyolite to dacite

Rb–Sr

609 ± 12

Bielski (1982)

Whole rock, andesite to rhyolite

Rb–Sr

587 ± 9

Whole rock, volcanic rocks

Rb–Sr

574 ± 12

Basta (1997)

Whole rock, volcanic rocks

Rb–Sr

587 ± 9

Bielski (1982)

Zircon, Upper Rutig rhyolite

U–Pb

607 ± 3

Be’eri-Shlevin et al. (2011)

Zircon, Upper Rutig dacite

U–Pb

611 ± 4

Zircon, Lower Rutig rhyolite

U–Pb

619 ± 4

Zircon, Lower Rutig dacite

U–Pb

623 ± 4

Zircon, rhyodacite

U–Pb

622 ± 3

Moreno et al. (2012) Be’eri-Shlevin et al. (2011)

Zircon, Upper Feirani rhyolite

U–Pb

597 ± 4

Zircon, Lower Feirani dacite

U–Pb

599 ± 3

Zircon, Lower Feirani dacite

U–Pb

603 ± 4

Mahash Volcanics

Zircon, granophyre

U–Pb

587.7 ± 5.6

Eyal et al. (2019)

G. Musa

Zircon, subvolcanic

U–Pb

593 ± 2

Moreno et al. (2012)

Jordan and Israel Wadi Araba, Jordan Ghuweir Mafic Suite

Whole rock, rhyolite

Rb–Sr

553 ± 11

Jarrar et al. (1992)

Rhyolite

Single zircon SIMS

598 ± 5

Jarrar et al. (2012)

Mafic Volcanics

Rb–Sr

572 ± 48

Jarrar et al. (2008) (continued)

494

M. K. Azer et al.

Table 20.1 (continued) Locality

Material, rock type

Method

Age (Ma)

References

Aheimir Volcanic, Jordan

Composite dyke–rhyolite core

U–Pb

607 ± 6

Ghanem et al. (2020)

Biotite in rhyolite dyke

40

600 ± 4

Amphibole in andesite dyke

40

Whole rock total gas age in dolerite dyke

40

Whole rock, rhyolite

Rb–Sr

Biq’at Hayareah, Israel

Ar/39Ar Ar/39Ar 39

Ar/ Ar

594 ± 3 *579 548 ± 5

Bielski (1982) Be’eri-Shlevin (2008)

Dyke suite, Israel

Zircon, dacite, rhyolite

U–Pb

591 ± 13

Dyke suite, Israel

Microdiorite dyke

U–Pb

614 ± 6

Amram, Israel

Alkali rhyolite

U–Pb

606 ± 3

samples of andesite and dacite from the Gabal Dokhan area and defined a Rb–Sr whole-rock isochron age of 592 ± 26 Ma with an initial 87Sr/86Sr ratio of 0.7028 ± 2 (MSWD = 1.8). Six ignimbrite samples and one rhyolite porphyry from the Gabal Nuqrah area defined a Rb–Sr age of 581 ± 14 Ma, with an initial 87Sr/86Sr ratio of 0.7033 ± 6 (MSWD = 2.32). Abdel-Rahman and Doig (1987) obtained a seven-point whole-rock Rb–Sr isochron age of 620 ± 16 Ma for the Dokhan Volcanics in the Ras Gharib area (G. El Kharaza). However, their data have an MSWD of 18, which indicates a significant geochronological error. The Dokhan Volcanics at the type locality have SHRIMP U–Pb zircon ages ranging from 602 ± 9 Ma for the lower part of the sequence and 593 ± 13 for the upper part (Wilde and Youssef 2000). Breitkreuz et al. (2008) reported SHRIMP U–Pb zircon ages of 615 ± 4 Ma and 616 ± 5 Ma for two lithic- and crystal-poor ignimbrites of the Dokhan Volcanics in the Northeastern Desert. Uranium– Th–Pb zircon geochronology on two samples of imperial porphyry and one sample of the common porphyry of the Dokhan Volcanics yielded ages ranging from 609 to 600 Ma (Abu El-Enen et al. 2018). Four imperial porphyry samples from Gabal Dokhan yielded an errorchron with an age of 560 ± 42 Ma and an initial 87Sr/86Sr ratio of 0.70283 ± 0.00011 (Makovicky et al. 2016), but this age is not acceptable for the Dokhan Volcanics, as it is both imprecise and much younger than the more precise zircon U–Pb ages. In south Sinai, Rb–Sr ages of 609 ± 12 Ma and 587 ± 9 Ma were obtained by Bielski (1982) for the volcanic series at Wadi Madsus (northern Kid area) and Wadi Rutig, respectively. The Meknas Volcanics of Wadi Watir yielded a Rb–Sr age of 574 ± 12 Ma (Basta 1997). More recently, Be’eri-Shlevin et al. (2011) published SIMS zircon U–Pb ages for the Dokhan Volcanics at Wadi Rutig and Gebel Ferani. The Rutig Volcanics record two groups of eruption ages of 620–618 Ma and 611–595 Ma, while the Ferani Volcanics define a restricted age range of 603–

595 Ma. Moreno et al. (2012) reported SHRIMP U–Pb zircon ages of 622 ± 3 for the Wadi Rutig Volcanics and 593 ± 2 Ma for the Gebel Musa neck of the Katharina Volcanics. In Saudi Arabia, many ages were obtained for the volcanic rocks in the Arabian Shield including both volcanic arc and post-accretionary phases (Kennedy and de Laeter 1994; Kennedy et al. 2004, 2005, 2010, 2011a, b). The more precise ages of post-accretionary volcanism are listed in Table 20.1. Few ages are available for the post-accretionary volcanic rocks in Jordan, but the rhyolite at Wadi Araba is reported to be much younger, with an age of 553 ± 11 Ma (Jarrar et al. 1992). In Israel, the rhyolite at Biq’at Hayareah yielded a whole-rock (Rb–Sr) age of 548 ± 5 Ma (Bielski 1982). Be’eri-Shlevin (2008) obtained many SIMS U–Pb zircon ages for alkali rhyolite (606 ± 3 Ma) and dyke suites (591 ± 13 to 614 ± 6 Ma) at Amram. According to Eyal et al. (2019), the age of the Amram rhyolite flows is 606 ± 3 Ma and may be broadly compared to the Gebel Musa neck in the Katharina Volcanics at 593 ± 2 Ma (Moreno et al. 2012).

20.6

Collision of East and West Gondwana

The ANS is composed of Neoproterozoic crust that formed in and adjacent to the Mozambique Ocean (Stern and Johnson 2010; Fritz et al. 2013) over a protracted period of *300 Ma, passing through a series of distinct stages (Stern 1994; Stern and Johnson 2010; Johnson et al. 2011; Fritz et al. 2013). The rifting and breakup of Rodinia into East and West Gondwana fragments occurred sometime in the Tonian (Stern and Johnson 2010). This was followed by a prolonged episode of formation of island arc terranes and their accretion, lasting from *800 to 670 Ma (Stern and Johnson 2010). East and West Gondwana and intervening fragments were then re-assembled by a continent–continent collision at

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*630 Ma. Stern and Johnson (2010) proposed that several intra-oceanic subduction zones were active from 870 to 630 Ma and that these arcs collided to form the ANS. In Sinai, Eyal et al. (2014a) concluded that the last island arc (Kid) in the northern ANS existed from *640 Ma to *620 Ma, implying that the Mozambique Ocean persisted (at least locally) until *620 Ma. The collision between East and West Gondwana was a complex and lengthy process, and it is challenging to interpret the age of any particular sample as the time of collision or even to argue that there was a single collision time (Andresen et al. 2009; Stern and Johnson 2010; Fritz et al. 2013). At the largest scale, the assembly of Gondwana was defined by the collision of two main fragments, East and West Gondwana (McWilliams 1981), but along the whole length of the suture this event was clearly diachronous and, given the irregular shapes of the continental margins, the age of “final collision” in one area may not preclude the continuing existence of open ocean elsewhere along the front (Fritz et al. 2013). There was also a lengthy series of small and large terrane collisions throughout the whole period from *850 to 620 Ma (Fritz et al. 2013). However, many published studies have attempted to define the general chronology of the collision on the basis of local or regional studies. This kind of reasoning has led several authors to assign the high-K calc-alkaline Dokhan Volcanics series in the Eastern Desert of Egypt to the pre-collisional phase of ANS development (e.g., Maurice et al. 2018). This interpretation is generally based on geochronological correlation between the Dokhan volcanic rocks and various deformed granite plutons, rather than on structural, stratigraphic or metamorphic evidence that might indicate that the Dokhan lavas themselves were caught up in collisional deformation after their eruption. Andresen et al. (2009), for example, studied the Meatiq area in the Eastern Desert of Egypt and found that deformation of I-type granite occurred at 605–600 Ma. From this evidence, they reasoned that collision between East and West Gondwana occurred at this late time, younger than many exposures of the Dokhan Volcanics, in support of their assignment of all Dokhan Volcanics to a pre-collisional phase. Without questioning the accuracy of the age of deformation of this granite, we must still ask whether it supports a broad regional interpretation. Deformation of various rock types may occur after collision and post-collisional plutons may inherit geochemical characteristics from earlier island arc stages. Hence, the measured age may only date the collision locally and a deformed pluton may not be an island arc stage pluton. We find the absence of deformation or metamorphism of the Dokhan sequence itself to be the most compelling evidence that the

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Dokhan Volcanics are post-collisional, but locally they may be weakly metamorphosed due to later granitic intrusions.

20.7

Existence of Old Continental Crust Beneath the ANS

Based on geochronological and isotopic data, the ANS is a product of the Neoproterozoic East African Orogeny (870– 550 Ma; Stern 1994). As described above, there is nonetheless disagreement in the literature concerning whether the ANS was entirely built of juvenile material, with no pre-Neoproterozoic crust involved in its formation, or whether instead extensive tracts of older pre-Neoproterozoic continental crust exist beneath the Pan-African rocks of the ANS, at least in the Eastern Desert of Egypt. Much of this debate has focused on the pre-870 Ma inherited zircons found in both ANS magmas and sediments. Some authors (e.g., Stern 2002; Liégeois and Stern 2010) have attempted to reconcile these views by stressing the longevity, durability and inheritability of zircons, such that the presence of pre-Neoproterozoic zircons need not require the existence of extensive tracts of older crust. Other work has argued that preserved crustal oxygen isotope signatures rule out recycling of zircons into the mantle and later incorporation in juvenile crust; this led Stern et al. (2010) to state that “it seems clear that *1 Ga crust can be found near the Sa’al–Zaghra region of the Sinai Peninsula.” Other work has focused on the controversial high-grade gneisses and migmatites in the Eastern Desert of Egypt, and whether their exposures are windows into remobilized Sahara metacratonic material as suggested by pre-Pan-African U–Pb zircon dates, despite their late Neoproterozoic whole-rock ages. Azer and Asimow (2021) preferred the classical interpretation of these gneisses and sheared granites as cratonic, arguing that the Pan-African ages, especially in the northwestern part of the ANS, date episodes of metamorphism rather than protolith ages. Maurice et al. (2018) claimed that the Dokhan volcanic rocks of the Eastern Desert of Egypt developed in a subduction zone environment and preferred an intra-oceanic arc setting over an active continental margin setting due to the absence of isotopic evidence for involvement of old continental crust. However, there is in fact extensive evidence for the involvement of old continental crust in the generation of the calc-alkaline Dokhan Volcanics (Breitkreuz et al. 2010; Stern et al. 2010a; Be’eri-Shlevin et al. 2011, 2012). For example, Li et al. (2018) studied a *755 Ma rhyolite cobble from the Atud Formation in the Eastern Desert. This cobble is characterized by exclusively negative zircon eHf(t) values, ranging from –1.9 to –16.3, indicating a highly

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heterogeneous source enriched in Hf and not a juvenile oceanic arc source. The rhyolite cobble from the Atud Formation was generated by melting of a Paleoproterozoic– Archean protolith at ca. 755 Ma (Li et al. 2018). It represents the first igneous rock identified from the subduction stage of the ANS that was unquestionably derived from old continental crust and confirms the presence of an old continental crust substrate beneath the ANS.

20.8

Examples of Post-accretionary Volcanic Sequences in the ANS

In the following section, we will provide several examples of Ediacaran post-accretionary volcanic sequences in the ANS, which include both early calc-alkaline volcanic units and late alkaline volcanic rocks. The calc-alkaline phase is represented by the following examples: (1) the Dokhan Volcanics in the Eastern Desert of Egypt; (2) the Rutig and Ferani Volcanics in Sinai; (3) the Murdama and Bani Ghayy groups in Saudi Arabia; (4) the Saramuj–Haiyala volcaniclastic units in Jordan; and (5) the Elat conglomerate in Israel. Examples of the late alkaline phase include: (1) the Attala felsite in the Eastern Desert; (2) the Katharina Volcanics in Sinai; (3) the Biq’at Hayareah Volcanics on the border between Sinai (Egypt) and the Negev (Israel); (4) the Elat Volcanics in Israel; and (5) the Aheimir Volcanic Suite in Jordan.

20.8.1 Early Calc-Alkaline Volcanic Sequences 20.8.1.1 Dokhan Volcanics in the Eastern Desert of Egypt Post-accretionary high-K calc-alkaline volcanic rocks are the most common type among the basement rocks of Egypt and are represented by the intercalation of Dokhan-type volcanic rocks and Hammamat-type sedimentary sequences (Fig. 20.5). Ali et al. (2010) proposed a simplified evolutionary diagram to summarize the tectonic events in the Eastern Desert of Egypt during Neoproterozoic time (Fig. 20.6) which can be applied to the northernmost ANS. This diagram is based on published geochronology, geochemical and Nd isotope data (Kröner et al. 1992; Ali et al. 2009b; Moussa et al. 2008; Wilde and Youssef 2000; Ali et al. 2010; Abu El-Enen et al. 2018). The Dokhan Volcanics are best exposed north of latitude 26° N in the Eastern Desert of Egypt, especially at their type locality at Gabal Dokhan. The south Eastern Desert lacks significant Ediacaran volcano-sedimentary successions such as the Dokhan Volcanics and Hammamat sedimentary succession that outcrop in the north and central parts of the

Eastern Desert (Stern and Ali 2020). Some volcanic successions in the south Eastern Desert have been mistakenly assigned to the Dokhan. El-Nisr (1997) mapped some mafic to felsic metavolcanic units along Wadi Allaqi as Dokhan Volcanics. Also, some volcanic exposures at Wadi Ranga (Gharib and Ahmed 2012) and again at Wadi Allaqi (e.g., El-Sayed et al. 2004) were assigned to the Dokhan Volcanics, although they have low- to medium-K characteristics similar to immature island arc volcanic rocks (Maurice et al. 2012). The Dokhan Volcanics differ from the preceding arc metavolcanic sequences in the abundance of felsic varieties, high potassium contents, the presence of ignimbrites and the absence of any metamorphism. The Dokhan Volcanics constitute a thick sequence of stratified lava flows of intermediate (andesite) to felsic (dacite, rhyodacite and rhyolite) compositions with minor basaltic andesite and basalt. The pyroclastic deposits and ignimbrites have the same composition and are commonly interbedded with the lavas. The Dokhan Volcanics have been separated in many areas into a lower and upper sequence (Basta et al. 1980; El-Gaby et al. 1989; El-Mahallawi 1999b; Khalaf 2012; Azzaz et al. 2015; Obeid and Azer 2015; Abdelfadil et al. 2018). The Hammamat sedimentary rocks associated with the Dokhan Volcanics are composed of continental molasse facies sediments and crop out sporadically throughout the Eastern Desert and Sinai, attaining a thickness of about 4000 m at the type locality in Wadi Hammamat in the central Eastern Desert. They include terrigenous clastic deposits, with abundant conglomerates, and are commonly intercalated with the Dokhan Volcanics (e.g., El-Gaby et al. 1989; El Kalioubi 1996; El-Gaby 2007; Eliwa et al. 2010; Bezenjani et al. 2014). The deposition of the Hammamat sedimentary rocks took place in several environments including alluvial fans, braided rivers and lakes (Samuel 1977; Grothaus et al. 1979; Ries et al. 1983; El-Gaby 1983, 1994; Stern et al. 1984; Ahmed et al. 1989). Deposition of the Hammamat sediments was in many cases synchronous with the eruption of the Dokhan Volcanics. There are many discrepancies in the literature regarding the nomenclature, age and setting of the Dokhan–Hammamat successions. The discrepancies focus around: (1) the stratigraphic relation between the Dokhan Volcanics and the Hammamat Group; (2) the tectonic setting of their emplacement; (3) the age range of the volcanic phase; (4) the presence or absence of compositional gaps; (5) the presence or absence of signs of metamorphism within the successions; (6) the geochemical characteristics of the Dokhan Volcanics; and (7) overall paleoenvironmental and facies models of their emplacement. In the following sections, we will discuss in detail all these sources of confusion.

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Fig. 20.5 Geologic map of the Eastern Desert and Sinai showing the distribution of the main geologic units and broad structural trends. Notice also the location of Dokhan Volcanics (after Johnson et al. 2011)

Stratigraphic Relation Between the Dokhan Volcanics and the Hammamat Sediments It has been commonly reported that the Dokhan Volcanics of the Eastern Desert are associated and intercalated with the Hammamat sediments and the same situation is recorded in Sinai. The Hammamat sediments have also been called the Hammamat Formation, the Hammamat Group or the Hammamat Clastics. Many workers consider them to be a continental molasse facies, largely formed of terrigenous clastic

rocks with abundant conglomerates (e.g., Grothaus et al. 1979; Reis et al. 1983; El-Gaby 1983; Kalioubi 1996; Samuel et al. 2011; Fowler and Osman 2013; Bezenjani et al. 2014). However, other workers consider them to be mainly fluvial deposits, with Ries et al. (1983) suggesting that the lower sequence was continuous across the Northeastern Desert rather than being deposited in restricted intermontane basins. This view was extended by Wilde and Youssef (2002) to the whole of the Hammamat sequence,

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Fig. 20.6 Simplified evolutionary diagram showing major tectonic events in the Neoproterozoic basement complex of Egypt (after Ali et al. 2010)

based on U–Pb zircon dating at Gabal Umm Tawat. Here, abundant Proterozoic and Archean zircons were identified that were considerably older than any known rocks in the Northeastern Desert, requiring transportation from a considerable distance. The stratigraphic position of the Hammamat sediments relative to the Dokhan Volcanics has long been a matter of debate. Some authors believe that the Hammamat sediments overlie the Dokhan Volcanics (e.g., Holail and Moghazi 1998; Akaad and Abu El-Ela 2002; Wilde and Youssef 2002), while others advocate that the sedimentary units exclusively underlie the volcanic deposits (e.g., Ghobrial and Lofti 1967; Stern and Hedge 1985; Willis et al. 1988). El-Gaby et al. (1984) indicated that the Hammamat sediments are penecontemporaneous with the Dokhan Volcanics and belong to a Cordilleran-type orogeny (i.e., tectogenetic uplift and erosion stages). The most recent studies indicate at least two pulses of Dokhan volcanic activity in the Eastern Desert and Sinai, both pulses including volcanic activity and deposition of Hammamat sediments. The two pulses are separated in time and can be distinguished by the sedimentological characteristics of the associated Hammamat rocks. Now, there is a consensus among geologists that the Hammamat sediments are penecontemporaneous with the Dokhan Volcanics and that they are intercalated with one another (e.g., Eliwa et al. 2010; Obeid and Azer 2015; Maurice et al. 2018). Tectonic Setting of the Dokhan Volcanics There is considerable discrepancy between the various interpretations that have been proposed concerning the tectonic setting of the Dokhan Volcanics in Egypt, and no consensus has been achieved among researchers. The dispute is centered around whether the Dokhan Volcanics were emplaced in a subduction environment, an extensional

setting that postdates crustal thickening or perhaps a post-collisional transition period from subduction to extension. Authors that have adopted a subduction-related tectonic setting for the Dokhan Volcanics reasoned mainly from chemical composition, which is similar to those of calc-alkaline volcanic rocks of typical Andean type (Basta et al. 1980; Ragab 1987; El-Gaby et al. 1988; Abdel-Rahman 1996; Samuel et al. 2001; Abdel Wahed et al. 2012; Alaabed and El-Tokhi 2014; Makovicky et al. 2016; Maurice et al. 2018). However, the high-K calc-alkaline character and other traits previously interpreted to indicate arc magmatism for the Dokhan Volcanics may simply reflect remelting of arc material or enriched lithospheric mantle from the earlier subduction period of ANS evolution at ca. 850–740 Ma. Similar inheritance is observed in some post-collisional plutonic rocks in the ANS (Azer et al. 2012; Gahlan et al. 2016) and elsewhere around the world (Coulon et al. 2002; Kay et al. 2013). Consequently, the persistence of a subduction signature in the magma source from the end of island arc activity until the early phase of Dokhan volcanism seems reasonable. This is supported by the lull in magmatic activity (740–640 Ma) between the end of the subduction-related island arc stage and the beginning of post-collisional volcanic activity in the northernmost ANS (Samuel et al. 2011), although ages spanning this gap have been reported from the southern portion of the ANS (Stern and Hedge 1985; Johnson and Kattan 2007; Ali et al. 2010). Potentially, island arc material from the 740–640 Ma lull period may be present in the sources of southern exposures of the Dokhan Volcanics. An extensional tectonic setting for the Dokhan Volcanics was introduced by Stern et al. (1984). This model was then adopted by many authors that described the Dokhan

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Volcanics as a syn-rift volcanic series (Stern and Gottfried 1986; Hegazy and Mawas 1992; Mohamed et al. 2000; Akaad and Abu El-Ela 2002; Abu El-Enen et al. 2018). The extensional model was based mainly on the bimodal nature of some Dokhan Volcanics (Stern et al. 1984; Stern and Gottfried 1986; Stern et al. 1988) and synchroneity with high-level A-type granite intrusions and dyke swarms. Subsequently, though, studies revealed that the Dokhan Volcanics were not synchronous with the A-type granites; rather, the dykes intruding the volcanic units differ from the lavas in geochemical characteristics (Abdel Wahed et al. 2012). The highly evolved rhyolites of the upper Dokhan Volcanics show some A-type characteristics due to the extensive degree of fractional crystallization. These apparently A-type rhyolites do not record a change in tectonic regime; they appear to belong to both the older (620– 618 Ma) and younger (610–595 Ma) magmatic cycles (Be’eri-Shlevin et al. 2011; Abdelfadil et al. 2018). Moreover, some island arc and collisional magmas are bimodal in composition (e.g., Smith et al. 2003; Aranda-Gomez et al. 2003), which indicates that bimodality by itself does not require a rift tectonic environment as previously proposed. A post-collisional transitional tectonic model for the Dokhan Volcanics was first proposed by Ressetar and Monrad (1983) and then applied by many authors (Mohamed et al. 2000; Moghazi 2003; El-Sayed et al. 2004; Eliwa et al. 2006; El-Bialy 2010; Azer and Farahat 2011; Khalaf 2012; El-Bialy and Ali 2013; Obeid and Azer 2015). Ressetar and Monrad (1983) determined that some Dokhan Volcanics are calc-alkaline in nature but contain slightly higher contents of alkalis and other incompatible elements than do typical continental margin volcanic suites; they interpreted these rocks to have formed in an environment transitional between compressive and extensional tectonic settings. However, slight enrichment in alkalis and in some incompatible elements in continental margin volcanic rocks is not uncommon (Pearce 1983; Ringwood 1990). High-precision U–Pb dating of subduction-related volcanic rocks in the Eastern Desert showed that their ages range from 770 to 720 Ma (Andresen et al. 2009; Ali et al. 2009b, 2010; Bühler et al. 2014; Abd El-Rahman et al. 2017). Recent U–Pb dating of the post-collisional phase in the ANS included both plutonic rocks and their volcanic equivalents and gave ages ranging between 630 and 590 Ma (Beyth et al. 1994; Jarrar et al. 2003; Be’eri-Shlevin et al. 2009a; Moussa et al. 2008; Be’eri-Shlevin et al. 2009a, 2011; El-Bialy and Ali 2013; Eliwa et al. 2014b). U–Pb zircon dating of the Dokhan Volcanics indicates that the early pulse began at 630 Ma, whereas the later pulse ended around 590 Ma (Wilde and Youssef 2000; Breitkreuz et al.

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2010; Be’eri-Shlevin et al. 2011; Abu El-Enen et al. 2018). Many authors have emphasized that there is no evidence for subduction or collisional-related processes in the ANS later than *635–620 Ma (e.g., Avigad and Gvirtzman 2009; Eyal et al. 2010), which suggests that the Dokhan Volcanics postdate the subduction period. Therefore, there is a general consensus that the Dokhan Volcanics formed in a post-accretionary tectonic setting (Khalaf 2012; Be’eri-Shlevin et al. 2011; Azer and Farahat 2011; and many others). Also, the two pulses of volcanic activity in the Dokhan Volcanics are comparable with the two magmatic phases of post-collisional calc-alkaline plutons: an early magmatic phase at ca. 635–620 Ma and a later one at 610– 590 Ma (Eyal et al. 2010). Maurice et al. (2018) considered that dating of the younger Dokhan felsic rocks (Breitkreuz et al. 2010; Abu El-Enen et al. 2018) does not support dividing the Dokhan volcanic rocks in the Northeastern Desert into two groups. Thus, they propose restricting use of the term Dokhan Volcanics to refer only to the calc-alkaline andesitic to dacitic volcanic rocks, excluding the later high-SiO2 rhyolites. This proposal would represent a substantial change in nomenclature, given the general consensus among scientists studying the Dokhan Volcanics since 1934 that they can be subdivided into two units, with the upper one consisting essentially of rhyolite (Hume 1934; El-Gaby et al. 1991; Moussa 2003b; Obeid and Azer 2015; Abdelfadil et al. 2018 and many others). Also, rhyolite is present in the lower Dokhan volcanic series (Be’eri-Shlevin et al. 2011). For example, the ca. 619 Ma rhyolites of the lower Dokhan volcanic series (Rutig Volcanics in south Sinai; Be’eri-Shlevin et al. 2011) clearly display A-type affinities that result from extensive fractionation of a calc-alkaline magma. Similar trends may characterize much of the Dokhan Volcanics in the Eastern Desert, but in order to assess this, a more direct correlation between age and chemistry throughout the different successions is needed. Compositional Gaps Geochemical studies carried out by several workers (e.g., Basta et al. 1980; Abdel-Rahman 1996; Samuel et al. 2001; Moussa 2003b; El-Sayed et al. 2004; Eliwa et al. 2006; Be’eri-Shlevin et al. 2011; Obeid and Azer 2015, Maurice et al. 2018) on the Dokhan Volcanics in different localities in Egypt indicate that these volcanic rocks do not exhibit a compositional gap. Rather, they populate continuous differentiation trends of medium- to high-K calc-alkaline affinity. Some early data suggested the presence of a compositional gap (bimodal suite) in some Dokhan Volcanics, but this has rarely been argued in recent years (Ressetar and Monrad

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1983; Stern et al. 1984; Stern and Gottfried 1986; El-Gaby et al. 1989, 1991; Mohamed et al. 2000; Asran et al. 2005). Abdel-Rahman (1996) suggested that the inference of bimodality in the Dokhan Volcanics is in part the result of equating unrelated stratigraphic units. Metamorphism In the central Eastern Desert, the island arc volcanosedimentary successions are metamorphosed (Akaad and El-Ramly 1960; El-Gaby et al. 1988; Hassan and Hashad 1990; El Habaak 2005; Ali et al. 2009b; Stern 2018; Abdel-Karim et al. 2019), whereas the Dokhan Volcanics and the Hammamat sediments are considered to be unmetamorphosed (Akaad and El-Ramly 1960; Akaad and Noweir 1969; Ghanem et al. 1973; Samuel et al. 2001; El-Bialy 2010; Maurice et al. 2018). However, many authors have described low-grade metamorphism of the Dokhan Volcanics (Ghobrial and Lotfi 1967; Ressetar and Monrad 1983; El Kalioubi 1988; El-Gaby et al. 1988, 1991; Eliwa et al. 2010; Makovicky et al. 2016). Some workers have indicated that only the earlier members of the Dokhan Volcanics and the lower Hammamat sediments in some areas of the north Eastern Desert and Sinai were subjected to low-pressure metamorphism (El-Gaby et al. 1991; Moussa 2003a; Samuel et al. 2011, Obeid and Azer 2015). This was confirmed by the presence of schistose texture and the development of chlorite, epidote and actinolite. However, the low-pressure metamorphism of some early Dokhan Volcanics and intercalated lower Hammamat sediments is most probably contact metamorphism due to the intrusion of the earlier phase of younger calc-alkaline granitoids (Ghanem et al. 1973; El Kalioubi 1988; Moussa 2003a; Be’eri-Shlevin et al. 2011; Azzaz et al. 2015; Abdelfadil et al. 2018). By contrast, Willis et al. (1988) reported that these sediments had undergone regional greenschist facies metamorphism. El-Gaby (1994, 2002) also considered that in some areas the lower part of the Dokhan Volcanics and associated Hammamat sediments underwent low-pressure regional metamorphism. Geochemical Characteristics The Dokhan Volcanics in the Eastern Desert of Egypt are mostly medium- to high-K calc-alkaline rocks (Abdel-Rahman 1996; Mohamed et al. 2000; Moghazi 2003; Eliwa et al. 2014a, Obeid and Azer 2015; Maurice et al. 2018) (Fig. 20.3a), except for those of the Southeastern Desert (Wadi Ranga and Wadi Allaqi), which belong to the low- to medium-K calc-alkaline volcanic series (El-Sayed et al. 2004; Gharib and Ahmed 2012). The low- to medium-K calc-alkaline members are comparable to immature island arc felsic metavolcanic rocks at Wadi Ranga and

Fig. 20.7 Sr/Y versus Y classification diagram (after Defant and Drummond 1990) for some post-accretionary calc-alkaline volcanics in Eastern Desert and Sinai showing an adakitic character

Wadi Shilman in the Eastern Desert (Ali et al. 2010; Maurice et al. 2012). Therefore, the so-called Dokhan Volcanics in the Southeastern Desert, especially those of Wadi Allaqi, require more detailed studies. The Dokhan Volcanics are characterized by their strong enrichment in the large ion lithophile elements (LILE) Rb, Ba, K and Th relative to the high field strength elements (HFSE) Nb, Zr, Y, P and Ti. They exhibit high LILE/HFSE ratios and depletion of Nb relative to N-MORB. Locally, adakitic lavas are recorded in some Dokhan volcanic successions in the Northeastern Desert and Sinai (Fig. 20.7), and are distinguished geochemically by very low Y (2.5 wt%), Cr (>50 lg/g), Ni (>25 lg/g) and Sr (>650 lg/g, with Sr/Y >40) contents (Eliwa et al. 2006; Be’eri-Shlevin et al. 2011; Obeid and Azer 2015; Abdelfadil et al. 2018; Maurice et al. 2018). They additionally have low Nb and Rb in contrast to the coexisting calc-alkaline lavas. The data from many upper Dokhan Volcanics show a range from calc-alkaline to alkaline. This feature can be attributed in part to a shift in the tectonic environment at 610–600 Ma as predicted by many workers (Stern and Gottfried 1986; Beyth et al. 1994; Be’eri-Shlevin et al. 2009a, 2011; Eyal et al. 2010) or to strong fractional crystallization. However, it is important to note that highly fractionated felsic I-type granites can have some major and trace element values that overlap those of typical A-type granites (Whalen et al. 1987). Various sources for the parent magmas of the Dokhan Volcanics have been proposed. The most likely source was partial melting of juvenile crustal rocks belonging to older accreted arc terrains of the ANS crust (Breitkreuz et al. 2010; El-Bialy 2010; Azer and Farahat 2011; Eliwa et al. 2014a, b), with minor contamination by older continental crust. This is supported by the existence of old inherited zircon crystals

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of Cryogenian age in some Dokhan Volcanics (Wilde and Youssef 2000; Breitkreuz et al. 2010). On the other hand, some authors consider the Dokhan Volcanics were generated through partial melting of a lithospheric mantle source with varying contributions from either older crustal rocks or slab-derived melts (e.g., Mohamed et al. 2000; Eliwa et al. 2006; Khalaf 2012; Alaabed and El-Tokhi 2014; Makovicky et al. 2016; Maurice et al. 2018). Adakitic magmas have been generated in different tectonic settings including subduction zones, collisional regions, active continental margins and intra-continental environments (e.g., Martin 1999; Xu et al. 2002; Guo et al. 2006, 2007; Dimalanta and Yumul 2008; Castillo 2012). The source of adakitic rocks has been debated extensively over the last few decades, and many petrogenetic models have been proposed to explain the mechanism of their formation (Defant and Drummond 1990; Defant et al. 1991; Fig. 20.8 Proposed schematic model for the evolution of Dokhan Volcanics and Hammamat sediments (after Eliwa et al. 2010); see text for discussion

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Castillo et al. 1999; Martin et al. 2005; Zhang et al. 2007; Wang et al. 2007; Li et al. 2009; Karsli et al. 2010). The model invoking partial melting of delaminated lower crust is the favored model for generation of the late Neoproterozoic adakitic rocks of the northernmost ANS (Obeid and Azer 2015), rather than melting of an oceanic ridge as proposed by Eliwa et al. (2006). In the northern ANS, lithospheric delamination resulted in upwelling of hot asthenosphere and crustal uplift, accompanied by melting of the delaminated lower continental crust (Avigad and Gvirtzman 2009; Farahat and Azer 2011). Evolution Models for the Dokhan Volcanics Eliwa et al. (2010) suggested a depositional model for the early Ediacaran volcano-sedimentary sequences in the Eastern Desert of Egypt (Fig. 20.8). It was based on work in the Gabal El Urf area, and Eliwa et al. (2010) applied this to

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the entire Eastern Desert. In this model, the geotectonic evolution of the Dokhan–Hammamat sequence displays four cycles. Phase I started with deposition of thick polymictic, well-rounded conglomerates resting on old rocks with an erosional unconformity. This scenario requires the presence of high and extended mountain ranges surrounding the depositional basin exposing metasedimentary, metavolcanic, plutonic and volcanic rocks of previous Pan-African orogenic stages in extended high-relief drainage systems. Rounding of boulders took place during fluvial transport in mountain rivers. Temporary storage of boulders in fluvial terraces inside the catchment area is assumed. At the beginning of Phase II, the depositional basin experienced a drastic change in environment. The conglomerates of Phase I were covered by a thick deepwater succession implying the presence of a lake. During this phase, two transitions from lacustrine to sandy braided river plain are documented, indicating a strong fluctuation of lake size and level. The shift from a high-energy environment in Phase I to the medium to low energy system of Phase II can theoretically be explained by strong denudation of the surrounding mountain ranges. However, more likely, this shift was related to tectonic damming and/or basin widening due to tectonic extension, leading to basin subsidence and a strong reduction of the river gradient. Tectonic or volcanogenic damming of the river created the lake. Tectonic extension probably resulted in down-faulting of the basin shoulders and basin widening. Phase III was characterized by perishing or shrinking of the lake and conglomerate sheets prograded into the center of the basin, suggesting a return to high-energy sedimentation. Various volcanic centers of silica-rich or silica-poor composition were active. In the central basin area, an andesitic phreatomagmatic vent situated in unconsolidated sediments led to the formation of

Fig. 20.9 Simplified model for the geodynamic setting of the post-collisional calc-alkaline Dokhan Volcanics in the Eastern Desert of Egypt

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volcanogenic mass flow deposits, hyaloclastite deposits and lavas. The last phase (Phase IV) represented the volcanic climax in the evolution of Dokhan–Hammamat sequence. The proposed model for the genesis of the Dokhan Volcanics is partial melting of lower continental crust and magma mixing at upper crustal levels. The extrusion of the Dokhan Volcanics was associated with extensive denudation of the pre-630 Ma orogenic edifice as well as the 630– 600 Ma post-collisional products. Such a scenario implies thinning of the previously thickened lithosphere, facilitating the upwelling of hot asthenospheric material and possibly associated with lithospheric delamination (e.g., Avigad and Gvirtzman, 2009). This may have been the trigger for the transition into dominantly alkaline A-type magmatism at 600 Ma. During the lithospheric thinning or delamination stage, upwelling of asthenospheric mantle would deliver heat sources that not only might trigger melting in the remaining lithospheric mantle and lower crust, but might also drive an extensional tectonic regime that would open conduits for small-volume alkaline magmas and create basins to provide space for volcano-sedimentary accumulation. This assertion is in good agreement with the record of peak extensional tectonics in the northernmost ANS at ca. 600 Ma (Stern et al. 1984; Garfunkel, 1999). The proposed geodynamic model for generation of the Dokhan volcanic sequences is summarized in Fig. 20.9, which shows the following steps: (1) lithospheric mantle delamination (though a similar series of events could follow from erosional thinning); (2) upwelling of asthenospheric mantle material; (3) low-degree partial melting of subcontinental lithospheric mantle to yield primary basaltic magmas; (4) partial melting of lower crust and generation of underplated basaltic magma that fractionated to produce intermediate magma; (5) magmatic heat input contributing to assimilation of continental crust,

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leading to formation and eruption of the basaltic andesite, andesite and dacite; and (6) after a hiatus, renewed eruption of an upper felsic volcanic cycle of rhyodacite and rhyolite generated by crustal melting.

20.8.1.2 Rutig and Ferani Volcanics in Sinai There are many high-K calc-alkaline volcanic successions in Sinai such as the Kid, Ferani, Rutig, Sa’al–Zaghra, Tarbush, Khashabi and Meknas Volcanics (Fig. 20.10). These

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sequences have been equated with the Dokhan–Hammamat successions of the Eastern Desert of Egypt (Schürmann 1966; El-Gaby et al. 1991; Samuel et al. 2001; Moussa 2003b; Azer 2007; El-Bialy 2010; Azer and Farahat 2011; Be’eri-Shlevin et al. 2011; Moghazi et al. 2012; Abdelfadil et al. 2018). Eyal et al. (1980) grouped the volcanosedimentary successions of Wadi Rutig and Gebel Ferani into the Ferani Group and argued that it was intruded by post-collisional granites (Eyal et al. 2010). Samuel et al.

Fig. 20.10 Geological map of the Sinai Peninsula (modified after Be’eri-Shlevin et al. 2011) with inset index map of the Arabian–Nubian Shield. Eight localities of post-accretionary high-K calc-alkaline volcanic series and six localities of alkaline volcanic series are indicated. High-K calc-alkaline volcanic units: (1) Iqna Shar'a, (2) Tarbush, (3) Rutig, (4) Sa’al–Zaghra, (5) Khashabi, (6) Kid–Malhak, (7) Ferani and (8) Meknas Volcanics. Alkaline volcanic units: (A) Gebel Katharina, (B) Gebel Ma'ain, (C) Gebel Abu Durba, (D) Wadi El-Mahash, (E) Wadi Khileifiya and (F) Gebel El-Homra

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a

b

Fig. 20.11 a Geological map of Saint Katharina area modified after Katzir et al. (2007), showing the distribution of post-collisional volcanics, including the Rutig and Katharina Volcanics. b Schematic diagram and age data showing the relations between different units within the Rutig volcano-sedimentary succession and field relations with country rocks (after Be’eri-Shlevin et al. 2011). AB–AFS = albite alkali feldspar, PA = peralkaline granite

(2011) have shown that ca. 610 calc-alkaline plutons intrude only the lower sections, whereas A-type granites intrude the entire section. There are some published Rb–Sr and SHRIMP U–Pb zircon ages for the post-accretionary high-K calc-alkaline volcanic rocks exposed in Sinai. Ages of 609 ± 12 Ma and 587 ± 9 Ma were given by Bielski (1982, in Bentor 1985) for volcanic units at Wadi Kid and Wadi Rutig (south Sinai). The Meknas Volcanics of Wadi Watir yield an age of 574 ± 12 Ma (Basta 1997). Recently, Be’eri-Shlevin et al. (2011) reported SIMS zircon U–Pb ages of 630– 590 Ma and 607–593 Ma for the Rutig and Ferani Volcanics, respectively, implying that these eruptions were coeval with those of the Dokhan Volcanics of the Eastern Desert. The most significant examples of post-accretionary volcanic series in south Sinai will be discussed in the following sections. Rutig Volcanics The Rutig volcano-sedimentary succession is exposed in the central part of southern Sinai ca. 5 km to the SE of Saint Katharina town (Fig. 20.11a). At the type locality in Wadi Rutig, the succession reaches a thickness of more than 2000 m, where it is overlain by Katharina Volcanics (*1000 m) with an angular unconformity (Eyal and Hezkiyahu 1980). Rutig Volcanics are also exposed to the west of the Katharina ring complex in the Wadi Rufaiyil–G. Tarbush area (Azer and Farahat 2011). Here, they overlie older rocks, including granitic gneiss, gabbro–diorite and

granodiorite and are in turn intruded or overlain by alkaline rocks (quartz syenite, alkali granite and peralkaline volcanic units). The Rutig Volcanics at Gebel Katharina and Gebel Tarbush are unconformably overlain by the ca. 590–585 Ma alkaline/peralkaline volcanic members of the Katharina Ring Complex and intruded by the ca. 583 Ma Katharina pluton (Bentor 1985; Azer 2007; Be’eri-Shlevin et al. 2009a; Moreno et al. 2012). The field relationships between the Rutig Volcanics and their country rocks are shown in a schematic diagram in Fig. 20.11b. The Rutig Volcanics comprise a series of intermediate to silicic lava flows and pyroclastic rocks with interbedded sediments, mainly conglomerates and arkoses (Eyal and Hezkiyaho 1980; Bentor 1985; Eyal et al. 1994; El-Masry and Hegazi 2005; Azer 2007; Be’eri-Shlevin et al. 2011; Samuel et al. 2011). Abdel Maksoud et al. (1993) and Abdel Khalek et al. (1994) divided the Rutig Volcanics into older and younger members, which are comparable to the ophiolitic Older Metavolcanics (OMV) and the island arc Younger Metavolcanics (YMV) of Stern (1981) in the Eastern Desert of Egypt. The presence of ophiolites in Sinai is doubtful (Bentor 1985; El-Gaby et al. 1990; Azer and Asimow 2021), although small bodies of serpentinite, a few tens of meters in length, are known in south Sinai at Kabr El-Bonaya in the Wadi Kid area (Beyth et al. 1978; Shimron 1981, 1984; Madbouly 1991; Abdel Khalek et al. 1994; Moussa 2002; Abu El-Enen and Makroum 2003; Mogahed 2020). Oweiss (1994) described the subduction-related volcanic sequences at Saint Katharina (including the Rutig

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Volcanics) as weakly metamorphosed Dokhan Volcanics, noting the presence of banded iron formations within the sequence. In recent years, the volcano-sedimentary succession at Rutig has been subdivided according to field relations and geochronology into lower and upper Rutig Volcanics, intercalated with clastic sediments (Azer 2007; Azer and Farahat 2011; Be’eri-Shlevin et al. 2011). The upper unit has a Rb–Sr age of 587 ± 9 Ma (Bielski, 1982). SIMS zircon U–Pb ages enable the recognition of two eruption phases within these successions, at ca. 620–618 Ma and 611– 595 Ma (Be’eri-Shlevin et al. 2011). The timing of the unconformity separating the lower and upper Rutig successions is constrained by the age of the Rahba pluton (610 ± 5 Ma; Be’eri-Shlevin et al. 2009a), which intrudes the lower Rutig succession, and the age of dacite (611 ± 4 Ma; Be’eri-Shlevin et al. 2011) from the upper Rutig succession, since the unconformity cuts not only the lower Rutig succession but also the Rahba pluton (Samuel et al. 2011). The lower Rutig succession comprises volcanic flows and pyroclastic deposits interbedded with conglomerates and sandstones. The volcanic flows and pyroclastic deposits (ca. 620–618 Ma; Be’eri-Shlevin et al. 2011) are dominantly dacite to rhyodacite, with lesser amounts of andesite and rhyolite. They are intercalated in the upper part by clastic sedimentary layers (Hammamat sediments) that consist of coarsely bedded sequences of conglomerate, greywacke and siltstone. The entire lower Rutig section is tilted, dipping ca. 50° to the west, and it is intruded by the Rahba granodiorite– monzogranite pluton (ca. 610 Ma; Be’eri-Shlevin et al. 2009a) and various dykes of unknown age. Both the pluton and the dykes abut against the upper unconformable contact with the overlying upper Rutig unit, indicating the unconformity is younger than ca. 610 Ma. The lower Rutig Volcanics are weakly thermally metamorphosed (Oweiss 1994; Azer 2007; Be’eri-Shlevin et al. 2011). The upper Rutig succession rests uncomfortably on top of the lower Rutig section and is itself unconformably overlain by peralkaline volcanic members of the Katharina sequence. It dips shallowly to the north and is intruded by only a few dykes (Be’eri-Shlevin et al. 2011). Similar to the lower Rutig succession, the upper Rutig section includes volcanic flows, pyroclastic deposits and sedimentary layers that include conglomerates. The upper Rutig Volcanics are unmetamorphosed, and the volcanic flows and pyroclastic deposits are mainly rhyolite and rhyodacite with minor dacite; andesite is completely absent.

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Geochemically, the Rutig volcanic rocks of both successions display large silica variations and are mostly high-K calc-alkaline rocks, corresponding to andesite through dacite–rhyodacite to rhyolite (Azer 2007; Azer and Farahat 2011; Be’eri-Shlevin et al. 2011). The high-K calc-alkaline character and other traits previously interpreted to indicate arc magmatism may simply reflect remelting of previously formed arc material of ca. 850–740 Ma age. The evolved rhyolites of the Rutig Volcanics display geochemical characteristics transitional to alkaline A-type due to the extensive degree of fractionation, rather than a switch in the tectonic regime, as this is manifested in both the older (620–618 Ma) and younger (610–595 Ma) magmatic cycles. Zircon in dacite of the lower Rutig sequence has d18O within the mantle range (5.3 ± 0.6, 2r error; Valley et al. 1998; Be’eri-Shlevin et al. 2011), indicating that the source of these magmas did not include significant supracrustal material, and also that magma ascent did not involve significant upper crustal assimilation. This feature is compatible with the general scarcity of xenocrysts in these dacites and in other dated volcanic rocks in both the upper and lower Rutig successions (Be’eri-Shlevin et al. 2011). The geochemical characteristics of the Rutig volcanic rocks, and their intimate association with molasse-type sediments, would appear to equate them with the Dokhan Volcanics and associated Hammamat sediments of the Eastern Desert of Egypt. In the Eastern Desert, these successions were developed within elongated N–S to NW–SE fault-bounded basins (Stern et al. 1984), indicating a partially extensional regime during this time (post ca. 620 Ma). In the case of the Rutig volcano-sedimentary successions, fault-bounded basins have not been demonstrated. Nevertheless, the preservation of the Rutig succession below a major (ca. 590–585 Ma) unconformity probably indicates that the succession was confined to a down-faulted block at this time. Furthermore, conglomerates in both the Rutig successions include detrital components as young as 620–600 Ma (Samuel et al. 2011). Some of these granitoid boulders are very similar to the currently exposed 635–600 Ma post-collisional plutons, implying that extensive erosion was ongoing during or very shortly after the emplacement of these plutonic rocks. Overall, this is taken to indicate generation of the post-accretionary calc-alkaline volcano-sedimentary successions in Sinai in an extensional regime. In the Rutig succession, basal andesites have characteristics of low-silica adakitic rocks (Sr = 1050–1420 lg/g; Y = 8–18 lg/g; Yb = 0.79–1.47 lg/g; Sr/Y = 57–178; Be’eri-Shlevin et al. 2011). It is noteworthy that adakitic

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b

a

Fig. 20.12 a Simplified geological map of the Gebel Ferani area (after Be’eri-Shlevin et al. 2011). b Schematic diagram and age data showing the relations between different units within the Ferani volcano-sedimentary succession and the relation with country rocks (after Be’eri-Shlevin et al. 2011)

rocks of relatively similar age have been reported in the Khashabi Volcanics of southern Sinai (Abdelfadil et al. 2018), showing that the occurrence of adakitic compositions is not a peculiarity of the Rutig succession and may reflect large-scale processes in the northern ANS. Ferani Volcanics The Ferani Volcanics are exposed at Gebel Ferani, which lies to the west of Dahab town on the Gulf of Aqaba, southeastern Sinai (locality 7 in Fig. 20.10). The Ferani Volcanics were first briefly described by Hume (1906) and were mapped by Eyal et al. (1980) as part of the Ferani Group. The succession of Ferani Volcanics (3–5 km thick) includes intermediate to felsic lava flows and pyroclastic deposits, alternating with immature sedimentary members. Bentor and Eyal (1987) noted that the Gebel Ferani succession uncomfortably overlies eroded upper amphibolite facies rocks of the Kid Group. The volcanics are folded and metamorphosed at the base to lower greenschist facies (Bentor and Eyal 1987; Moussa 2003a; Be’eri-Shlevin et al. 2011). The volcanic succession at Gebel Ferani has been studied by many workers, with most finding it to be similar to the Dokhan Volcanics of the Eastern Desert (Shimron 1980; Bentor 1985; Bentor and Eyal 1987; Khalaf et al. 1994; Moussa 2003a, b; Be’eri-Shlevin et al. 2011). However, Abu El-Leil et al. (1990) considered the Ferani Volcanics to be younger than the Dokhan Volcanics and to have calc-alkaline to alkaline affinities. On the other hand, Khalaf

(2002) considered the Ferani Volcanics to be an alkaline anorogenic volcanic series based on the high Nb concentrations (up to 127 lg/g), higher than the typical alkaline/peralkaline volcanic rocks of Katharina type (Azer et al. 2014; Eyal et al. 2014a, b). The Ferani Volcanics were extruded onto older calc-alkaline granite (G1) and then intruded by the Dahab alkali granite (591 ± 6 Ma; Be’eri-Shlevin et al. 2009a) (Fig. 20.12a). The relationships between the different rock units within the Ferani volcano-sedimentary succession and their country rocks are shown schematically in Fig. 20.12b. The younger calc-alkaline granite (G2) has sharp contacts against the G1 granites and the Ferani Volcanics and encloses xenoliths of Ferani andesite. The Ferani Volcanics and surrounding granites are dissected by several dykes with a variety of compositions (basalt, andesite and rhyolite). In detail, the Ferani Volcanics can be differentiated into two phases that are separated locally by coarse conglomerate beds (Moussa 2003b; Be’eri-Shlevin et al. 2011). The early (lower) phase is composed of basaltic andesite, andesite and dacite with subordinate tuffs, whereas the younger (upper) phase comprises dacite, rhyodacite, rhyolite, ignimbrite and pyroclastic deposits (agglomerates and tuffs). Both pulses of volcanism are intercalated with immature clastic sediments, including siltstone and conglomerate. The Ferani Volcanics have been dated using the SIMS technique by Be’eri-Shlevin et al. (2011). Dacites and rhyolites from the lower and upper parts of the Ferani Volcanics

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yield ages of 603–597 Ma, with zircon d18O values within the mantle range. Rare 635–620 Ma inherited zircons indicate minor contamination by earlier but still juvenile Neoproterozoic crust. The positive whole rock eNd(t) values (+2.7 and +3.2) of the Ferani Volcanics are consistent with the mostly juvenile character of the ANS magmas (Be’eri-Shlevin et al. 2011). Geochemically, the Ferani Volcanics are enriched in most LILE and LREE and depleted in most HFSE, as well as in HREE (Moussa 2003b; Be’eri-Shlevin et al. 2011). They evolved from a high-K calc-alkaline magma that was generated in a post-collisional regime. This magma was derived from the mafic lower crust, which likely melted due to lithospheric delamination (Be’eri-Shlevin et al. 2011). Contamination by continental crust, followed by fractional crystallization, was responsible for the variation observed within this suite, with plagioclase and amphibole being the dominant fractionating phases. There has been some disagreement among workers concerning the tectonic setting of the Ferani Volcanics. This is centered around whether the Ferani Volcanics were emplaced in an island arc or active continental margin subduction environment (Khalaf et al. 1994; Moussa 2003b), during the post-collisional stage (Be’eri-Shlevin et al. 2011), or in a within-plate setting (Khalaf 2002). Based on recent geochronological data and the geochemical characteristics of the Ferani Volcanics, there is now a general consensus emerging that they formed in a post-collisional tectonic setting (Be’eri-Shlevin et al. 2011). The geochemical characteristics and age of the Ferani Volcanics confirm that they correlate with the high-K calc-alkaline Dokhan Volcanics of the Eastern Desert. The volcanic rocks exposed in the Ferani succession also compare well in terms of age and geochemical patterns with the 610–590 Ma calc-alkaline plutonic suite, although some of the upper Ferani Volcanics display geochemical similarities with the later alkaline plutonic suite (608–580 Ma). This may in part reflect a change in the tectonic regime at ca. 610–600 Ma, with a transition to more alkaline compositions as reflected in the plutonic suites (Eyal et al. 2010), or it may be due to more extensive fractionation.

20.8.1.3 Murdama and Bani Ghayy Groups in Saudi Arabia The distribution of the Murdama and Bani Ghayy groups in the post-amalgamation basins in the northeastern Arabian Shield is shown in Fig. 20.13. Murdama Volcano-Sedimentary Succession The Murdama volcano-sedimentary succession, also known as the Murdama Group, represents the largest of the early

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Ediacaran post-amalgamation depositional assemblages in the northeastern part of the Arabian Shield (Johnson 2003). It crops out within an elongate area (600 km in length and 120 km in width) and continues a further 100 km southeast beneath Phanerozoic cover (Johnson and Stewart 1995). The volcano-sedimentary succession of the Murdama Group unconformably overlies metamorphosed volcanic and plutonic rocks belonging to the Afif composite terrane. It was deposited at the end of, and soon after, the Nabitah orogeny (680–640 Ma). According to Stoeser and Stacey (1988), the Nabitah orogeny was a major central shield deformational event that resulted from the convergence and suturing of the Afif composite terrane with the Asir, Jiddah, and Hijaz terranes. Greene (1993) estimated the Murdama Group was originally *20 km thick. The group contains a vast amount of detrital material, consistent with its deposition during and soon after uplift associated with the Nabitah orogeny. It has an age of 630–624 Ma in the southern part of the Maslum basin (Kennedy et al. 2010), but may be as old as 650 Ma farther north (Cole and Hedge 1986). It is intruded by granitoids belonging to the Idah suite (*620–615 Ma) and the Abanat suite (585–570 Ma) (Cole and Hedge 1986). The Murdama volcanic rocks are interbedded with sandstone and conglomerate, with subordinate siltstone and limestone, all regionally metamorphosed to lower greenschist facies. The sandstones are well-bedded lithic (volcanic) arenite and arkosic arenite derived from a volcanic arc source. Sedimentary structures include planar cross-bedding, ripple crosslamination, planar lamination, grading and scour-and-fill structures. Fine-grained sandstone grades locally into siltstone, and coarse-grained sandstone grades into pebbly sandstone and conglomerate; in several places, the rocks form upward-fining cycles of sandstone, siltstone and shale, less than one to several meters thick (Wallace 1986). Conglomerate, particularly abundant at the base of the group, is indistinctly bedded to massive. It is mainly clast-supported and is composed of subangular to subrounded pebbles, cobbles and rare boulders of meta-andesite, metadiorite, metadacite, marble, various felsic metavolcanic rocks, granodiorite, granite and mica schist. Lenses of limestone, tens of meters to 300 m thick, occur in the northern part of the outcrop, and a limestone unit more than 1000 m thick occurs toward the base of the group along the eastern margin of the Maslum basin. The limestone is commonly well bedded and includes tan-colored, fine-grained micrite, dark brown siliceous impure limestone, black carbonaceous limestone and tan stromatolitic limestone. Basalt, andesite and rhyolite of the Afif formation underlie the Murdama sandstone along the southwestern side of the Maslum basin. The volcanic rocks have a fairly mature calc-alkaline and high-K calc-alkaline geochemical

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signature. There is a compositional gap between the most evolved andesites and most primitive rhyolites in the suite; this bimodal character is possibly evidence for extension and rifting during initiation of Murdama Group deposition. The group is deformed by gently plunging, open, upright, north-trending folds with an associated subvertical axial plane cleavage. En echelon folds at the southwestern margin of the Maslum basin probably developed during a period of

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ductile shearing in the adjacent Ar Rika fault zone (Johnson 2003) (Fig. 20.13), one of the many northwest-trending sinistral transcurrent Najd faults that cut the northern Arabian Shield. The trend of the folds indicates deformation by bulk east–west shortening. Vertical bedding, indicative of strong rotation of the Murdama Group, may reflect adjustment to basement faulting (Johnson and Kattan 2012). The abundance of limestone in the Murdama Group is evidence

Fig. 20.13 Geologic map showing the distribution of the Murdama and Bani Ghayy groups in the post-amalgamation basins of the northeastern Arabian Shield (after Johnson 2003)

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that the ANS orogenic belt was connected with an ocean during the late Cryogenian–early Ediacaran, and the group may well have been deposited in an arm of what remained of the Mozambique Ocean following arc collisions during the preceding episodes of terrane amalgamation. Bani Ghayy Group The Bani Ghayy Group crops out in narrow belts (20–50 km wide) in the central part of the Arabian Shield (Fig. 20.13). The separate basins are up to 350 km long and add up to a combined strike length of more than 600 km. They are commonly inferred to be fault basins or grabens. The Jabal Bani Ghayy exposures are in the south-central part of the Mujayrib basin. The group is reported to be in excess of 6 km thick (Agar 1986) and has been dated by the multigrain U–Pb zircon technique on rhyolite at 620 ± 5 Ma (Stacey and Agar 1985). There is also a 3-point Rb–Sr whole-rock isochron age of 620 ± 95 Ma from andesite and basalt (Fleck et al. 1980). However, the group appears to be intruded by granite of the Haml batholith (650–600 Ma; Agar et al. 1992; Aleinikoff and Stoeser 1988) and by a quartz porphyry dyke with a SIMS U–Pb age of 646 ± 11 Ma (Doebrich et al. 2007), suggesting the group is older than indicated by the multigrain zircon U–Pb and Rb–Sr ages. The volcano-sedimentary succession of the Bani Ghayy Group consists of virtually unmetamorphosed volcanic flows (basalt and rhyolite) intercalated with sedimentary members (sandstone, conglomerate and limestone); this succession closely resembles the Murdama Group. In the north, the Bani Ghayy and Murdama groups are separated by the Halaban–Zarghat fault zone (Fig. 20.13). Because of this proximity and the similarity of lithology and age, some workers treat the groups as correlative and do not recognize the Murdama Group as a separate entity (e.g., Brown et al. 1989). At Jabal Bani Ghayy, the group includes pebble-to-boulder-sized basal conglomerate that transitions upward, through agglomerate, into tuffaceous wacke and siltstone. Basalt, andesite and rhyolite crop out northeast of Jabal Bani Ghayy. Agar (1988) mapped an unconformity at the base of the Bani Ghayy Group, with a basal conglomerate overlying red alkali feldspar granite that intruded into the underlying Siham Group. The eastern margin of Bani Ghayy exposures is a serpentinite-decorated west-vergent thrust, the Tawilah fault zone, separating the Bani Ghayy Group in the footwall from the Siham Group in the hanging wall. The fault zone continues to the south, swings to the SE and merges with the Ruwah fault zone. This structure is important as a site of gold and tungsten mineralization; it is being extensively explored by Ma’aden Mining Company.

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20.8.1.4 Saramuj Conglomerate in Jordan The Saramuj Conglomerate Formation (Fig. 20.14) and the overlying Haiyala Volcaniclastic Formation together make up the Safi Group (Ghanem and Jarrar 2013). They were deposited in a post-orogenic basin comprised of molasse-type sediments, i.e., poorly sorted conglomerate and sandstone. The age of the Saramuj conglomerate is constrained by the intrusion of a 595 Ma monzogabbro in Wadi Qunai (Jarrar et al. 1993). The intrusive contact has a distinct contact aureole (Jarrar et al. 1993; Ghanem 2009). Furthermore, the Saramuj conglomerate unconformably overlies the *610 Ma Aqaba Complex granitoids in Wadi Abu Barqa (Jarrar 1985). This implies an age for the Saramuj conglomerate between 610 and 595 Ma. The ten youngest 207 Pb-corrected zircon U–Pb analyses from matrix samples in the conglomerate define a maximum depositional age  615 Ma, which is in good agreement with its stratigraphic position (Yaseen et al. 2013). Furthermore, the Saramuj conglomerate is dissected by a swarm of NE–SW striking rhyolitic and trachy-andesitic dykes dated to ca. 561 Ma. The maximum age of deposition of  615 Ma for the Saramuj conglomerate means it is broadly coeval with the volcano-sedimentary successions in Sinai (620–590 Ma; Samuel et al. 2011; Moghazi et al. 2012), the Elat area in Israel (605–580; Morag et al. 2012) and the Eastern Desert of Egypt (630–585; Wilde and Youssef 2000, 2002; Bezenjani et al. 2014). The Saramuj conglomerate shows other similarities with the aforementioned basins: Each has a prominent 625–650 Ma zircon age peak and a distinct gap between ca. 650 Ma and 700 Ma, the regional magmatic lull in the northernmost ANS (e.g., Samuel et al. 2011; Morag et al. 2012). The main difference between the Saramuj conglomerate and other basins in the north ANS is the absence of ages older than 750 Ma in the conglomerate; older ages are recorded in the basins in the Northeastern Desert (Wadi Umm Tawat and Wadi Igla), in Sinai (Rutig and Kid) and in the Elat area. This difference supports the later onset of subsidence in the Saramuj basin and local sourcing of the sediments. 20.8.1.5 Elat Conglomerate In southern Israel, the final stages of late Neoproterozoic evolution of the ANS include the deposition of a subaerial volcano-sedimentary sequence above exhumed deep-seated plutonic–metamorphic rocks. The sediments in this sequence, known collectively as the Elat conglomerate unit, consist primarily of immature polymict conglomerates whose composition closely reflects their underlying basement (Morag et al. 2012). Uranium–Pb dating and Hf

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Fig. 20.14 Geologic map showing the extent of the Saramuj Conglomerate (after Jarrar et al. 1992), the insert map showing a general map of Jordan

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isotopic analysis of single zircons in two volcanic samples from Elat provide evidence for two major magmatic cycles, separated by a lull; early island arc magmatism occurred at 880–760 Ma, whereas post-collisional granitoid intrusions and associated volcanic eruptions were superimposed on the amalgamated arc crust between 660 and 580 Ma (Morag et al. 2012). Both the Elat conglomerates and associated volcanic units are essentially coeval with the younger conglomerates of Wadi Rutig (middle and upper) and are assigned an age of ca. 600 Ma (Bentor 1985; Garfunkel et al. 1999; Be’eri-Shlevin 2008; Be’eri-Shlevin et al. 2011).

20.8.2 Late Alkaline/Peralkaline Volcanic Sequences 20.8.2.1 Attala Felsite in the Eastern Desert of Egypt Felsite is a term used to describe any light-colored fine-grained felsic igneous rock (with or without phenocrysts), composed mainly of quartz and feldspars and characterized by felsitic texture (Bates and Jackson 1984). According to Le Maitre et al. (1989), the term felsite is applied to any microcrystalline rock of granitic composition. The felsites in the Eastern Desert were emplaced in pulses over a long time period (Taman 1996). Based on the stratigraphic position of most felsites in the Eastern Desert— above the Dokhan Volcanics, cutting the Hammamat sediments, and intruded by the younger granites—El-Baily (2020) suggested that they probably were emplaced at around 600 Ma. In the Eastern Desert, “true” Katharina Volcanics have not been described, although alkali rhyolite, alkali microgranite and alkaline felsite have been recognized by many authors at several localities (Essawy 1972; El-Mahallawi and Ahmed 1993; El-Mahallawi 1999a; Abu El-Ela 2001). Akaad (1959) and Akaad and Noweir (1969) reported several discrete felsite bodies in the Eastern Desert that postdate the Hammamat sediments. The felsites occur as effusive bodies as well as sheets, dykes and cone-like intrusions. They are exposed at several localities, including at Gabal Atalla, Wadi Igla, Rasafa-Shimiyia area, Wadi Ranga, Wadi Shait, Wadi El Miyah and Wadi Sodmein (Akaad 1957, 1959; Akaad and El-Ramly 1958; El-Mahallawi and Ahmed 1993; El-Mahallawi 1999a; Abu El-Ela 2001; Asran et al. 2005). The majority of these felsites crosscut the Hammamat sediments. El-Ramly (1972) and Akaad and Noweir (1977, 1980) coined the name “post-Hammamat felsites” for the felsites in the Eastern Desert, the main representatives of which are the Atalla, Igla and Rasafa-Shimiyia felsites. The stratigraphic position of these felsites has been a matter of debate (Dardir and Abu Zeid, 1972; Saleeb-Roufaiel 1975; El-Gaby 1994;

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Takla and Hussein 1995; Asran et al. 2005), but now a general agreement prevails that the so-called postHammamat felsites were emplaced over a protracted time span. Based on field observations and chemical compositions of the felsites in the Eastern Desert, Greenberg (1981) indicated they might be directly related to, or comagmatic with, the younger granites. El-Mahallawi and Ahmed (1993) and El-Mahallawi (1999a) concluded that the Igla felsites are calc-alkaline (I-type felsites), whereas the Atalla felsites are alkaline (A-type felsites). The Atalla felsites are the best example of postaccretionary alkaline rocks in the Eastern Desert and are comparable to the typical Katharina Volcanics in south Sinai (Hassan 1998a; Abu El-Ela 2001). They represent the largest felsite intrusion recorded in the basement complex of the Eastern Desert (Essawy and Abu Zeid 1972), where they occur as an elongate mass (23 km long and 1–3 km wide) with rugged relief, in marked contrast to the surrounding rocks. The Atalla felsites occur north of the Qift–Quseir Road (*10 km from the entrance to W. Atalla) and represent a subvolcanic intrusion. They are buff or pink to brown in color with intense jointing, decorated with iron and manganese oxides. They are fine-grained to aphanitic and porphyritic in texture. They intrude ophiolitic rocks, island arc metavolcanic rocks and Hammamat sediments. In turn, they are intruded by post-orogenic granitic intrusions (Fig. 20.15). The contacts of the felsite with the country rocks are knife-sharp and intensely mylonitized. It seems that the felsite body was emplaced along an early strike-slip fault, although this fault was also reactivated after the emplacement of the felsite. Most outcrops of Atalla felsite are strongly sheared and cataclastic, exhibiting stretching lineations and slickensides, especially at their margins, due to the effect of the Atalla Shear Zone (El Kalioubi and Osman 1996), which consists of a series of thrust faults dipping toward the NE, with a transport direction toward the SW (Akawy 2003, 2007). Essawy and Abu Zeid (1972) concluded that the Atalla felsite intrudes the rhyolite tuffs and flows of the Dokhan Volcanics, which in this area are overlain unconformably by sediments of the Hammamat Group. They considered the Atalla felsite and associated rhyolite to be derived from the same parent magma that gave rise to the Dokhan Volcanics, even though the rhyolites predate the deposition of the Hammamat sediments and the felsite postdates them. Also, Essawy and Abu Zeid (1972) suggested that the Atalla felsites were emplaced along faults or fractures within the Hammamat basins. On the other hand, Hassan and Hashad (1990) considered that the felsites might be directly related to, or comagmatic with, the younger granites. Taman (1996) divided the Atalla felsites into rhyolite, rhyodacite, spherulitic rhyodacite, dacite and garnet-bearing

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Fig. 20.15 Geologic map of the Attala felsites in the Eastern Desert of Egypt, after Abu El-Ela (2001)

felsite. Abu El-Ela (2001) adopted a different scheme and divided the felsites into comendite, pantellerite and rhyolitic tuff. The felsites yield a Rb–Sr isochron age of 588 ± 12 Ma (Hassan 1998a), suggesting they most probably are equivalent to the Katharina Volcanics of south Sinai. The initial 87Sr/86Sr ratio of 0.7043 ± 0.0005 for these rocks implies derivation from the lower crust or upper mantle (Hassan 1998a). Geochemically, they represent the subvolcanic or extrusive equivalents of alkaline to peralkaline granites. Taman (1996) concluded that the Atalla felsites might be equivalent to the Um Had granite and therefore have geochemical characteristics comparable with the Egyptian younger granite of El-Gaby (1975). The Atalla felsites are alkaline rocks that were emplaced in a

within-plate environment (Taman 1996; El-Mahallawi and Ahmed 1993; El-Mahallawi 1999a; Abu El-Ela 2001).

20.8.2.2 Katharina Volcanics in Sinai The last phase of post-accretionary magmatism in the ANS produced alkaline to peralkaline volcanic rocks, known as the Katharina Volcanics (Agron and Bentor 1981; Bentor 1985). These volcanic units are restricted in their occurrence to Sinai, and few equivalent rocks have been identified elsewhere in the ANS, except for the alkaline felsites of Gabal Atalla in the Eastern Desert, which were equated with the Katharina Volcanics by Abu El-Ela (2001). Similar alkaline volcanic rocks are recorded in different localities in Sinai, including at Gebel Ma’ain, Gebel Abu Durba,

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Fig. 20.16 a Schematic geological cross-section showing the relation between the post-collisional volcano-sedimentary successions in the Gebel Tarbush–Wadi Rufaiyil area and their country rocks in Sinai (after Azer and Farahat 2011). b Schematic diagram showing the relation between the post-collisional volcano-sedimentary successions in the Saint Katharina area and their country rocks in Sinai (after Azer et al. 2014). Ages from Be’eri-Shlevin et al. (2009a, 2011) and Moreno et al. (2012)

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a

b

Wadi Mahash, Wadi Khileifiya and Gebel El-Homra (Agron and Bentor 1981; Samuel et al. 2007; El-Bialy 1999; Abu El-Leil and Abdel Razek 2001; El-Bialy and Hassen 2012). The Aheimir volcanic suite in southwest Jordan is also equivalent to the Katharina Volcanics in Sinai and underwent the same K-metasomatism (Jarrar and Ghanem, this volume). The typical Katharina Volcanics are exposed at Gebel Saint Katharina and Gebel Tarbush (Fig. 20.11a). They reach a thickness of about 1000 m, overlying the Rutig Volcanics (>2000 m) across an angular unconformity (Eyal and Hezkiyahu 1980). Eyal et al. (1994, 1995) assigned the name Jurjunia Formation to the Katharina Volcanics at their type locality. El-Masry et al. (1992) stated that the Katharina

Volcanics at the type locality represent only the preserved remnants of the ring dyke roof (or the caldera subsidence intrusion) and were previously misidentified as volcanic. Abdel Maksoud et al. (1993) and Abdel Khalek et al. (1994) disagreed, concluding that they represented discrete volcanic units and differentiating them in the Saint Katharina area into subduction- and rift-related volcanic sequences; each volcanic sequence was further subdivided into older and younger groups. The older and younger subduction-related volcanic sequences are mostly of rhyolitic composition, and they correlated them with the OMV and YMV of Stern (1981). They added that the older rift-related volcanic units are alkali basalts, whereas the younger ones are dominated by rhyolites, alkali rhyolites and rhyolitic ignimbrites.

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Fig. 20.17 Photomicrographs showing a sodic amphibole in the Katharina Volcanics, b sodic pyroxene in the Katharina Volcanics, c porphyritic rhyolite with K-feldspar phenocrysts in the Biq’at Hayareah area, d porphyritic trachyte with sanidine phenocrysts in the Biq’at Hayareah area

The Katharina Volcanics are associated with the Katharina Ring Complex in the central part of south Sinai (Fig. 20.11a). The ring complex is of late Neoproterozoic age (605–580 Ma; Katzir et al. 2007; Moreno et al. 2012; Eyal et al. 2014a, b) and was developed during the final tectonomagmatic stage of the Pan-African in the northern ANS. The country rocks of the Katharina Ring Complex include the Wadi Solaf gneisses, the Rutig Volcanics, and gabbro–diorite and calc-alkaline granitoids. The field relationships between the Katharina Volcanics and the country rocks are clarified in the geological sketch (Fig. 20.16a) and schematic diagram (Fig. 20.16b). At Gebel Katharina and Gebel Tarbush, the Rutig Volcanics are unconformably overlain by the Katharina Volcanics and intruded by the ca. 583 Ma Katharina pluton (Katzir et al. 2007; Moreno et al. 2012). The Katharina Ring Complex includes the Katharina Volcanics, subvolcanic bodies, ring dykes and the granitic Katharina pluton. The Katharina Volcanics represent the earliest stage of the ring complex, which was subsequently followed by large-scale cauldron subsidence and injection of ring dykes into the ring fractures. This was followed by

the intrusion of an alkali feldspar granite pluton at 596 ± 3 Ma (Moreno et al. 2012), which is geochemically comparable to the preceding volcanic units (Bentor and Eyal 1987; Katzir et al. 2007). The Katharina subvolcanic body forms the highest peak in G. Musa; it was dated at 593 ± 2 Ma (Moreno et al. 2012). There are inner and outer ring dykes, with the former exposed around the G. Musa area having a semicircular shape. The inner ring dykes intrude the Rutig and Katharina Volcanics, as well as some subvolcanic rocks, with sharp and subvertical contacts, but are cut by the Katharina pluton. The composition of the inner ring dykes varies from quartz syenite to quartz trachyte. The outer ring dykes were emplaced along a semicircular fault system that outlines the ring complex (Eyal et al. 1980); their width ranges from a few meters to 2.5 km. The contacts with country rocks are always sharp, and both inner and outer contacts dip inward. The outer ring dyke is composite in nature and consists of successive sheets of quartz syenite and quartz trachyte, with minor trachybasalt. The Katharina Volcanics form stratified volcanic sequences of rhyolitic alkaline lavas (alkali rhyolite,

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comendite and pantellerite) with abundant ignimbrites and pyroclastic deposits, associated with subvolcanic intrusions of rhyolite–granite porphyry and granophyre. The pyroclastic deposits include volcanic breccia with lithic and crystal tuffs. The rhyolites occur as sheets in the upper part of the Katharina volcanic section or as roof pendants above the subvolcanic bodies. The ignimbrites occur as successively intercalated sheets among the volcanic flows. Volcanic breccia is exposed in the central part of the section and is composed of angular rock fragments, 1–50 cm in size, embedded in crystal-rich rhyolitic tuff. Mineralogically, the peralkaline Katharina Volcanics contain sodic amphiboles and aegirine (Farahat and Azer 2011; Gahlan et al. 2016; Azer et al. 2014) (Fig. 20.17 a, b). The rocks were emplaced during a non-orogenic period under tensile stresses, accompanied by block faulting and differential uplift (Eyal and Hezkiyahu 1980; Agron and Bentor 1981; Farahat and Azer 2011; El-Bialy and Hassen 2012; Azer et al. 2014; Eyal et al. 2014a, b). Neodymium isotope data show some variations among the members of the Katharina Ring Complex (Katzir et al. 2007; Eyal et al. 2010). The Katharina pluton has eNd(t) between 2.6 and 3.9, while the Katharina Volcanics and ring dykes have higher values (4.2–5.5). The differences in Nd isotope ratios within the Katharina Ring Complex can be attributed to generation from heterogeneous sources. Azer et al. (2014) attributed the mantle signature in the Katharina Ring Complex to the juvenile nature of the ANS crustal sources.

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20.8.2.3 Biq’at Hayareah Volcanics in Sinai The Biq’at Hayareah area is located on the border between Sinai (Egypt) and the Negev (Israel). In the published literature, the volcanic rocks of this area have been assigned different names, including Biq’at Hayareah (Agron and Bentor 1981), Gebel El-Homra (EGPC/CONOCO geologic map 1987) and Neshef Massif (Mushkin et al. 1999). In this work, we adopt the name “Biq’at Hayareah,” which includes rocks in both Egypt and Israel. The volcanic sequence at Biq’at Hayareah has been correlated with the Katharina Volcanics (Agron and Bentor 1981; Samuel et al. 2007). The volcanic rocks of the Biq’at Hayareah area (Fig. 20.18) occur as ridges and conical hills with steep slopes, drained by almost dead wadis filled by alluvial accumulations and capped in places by a thin layer of loess. The northern ridges are mainly aligned in an E-W direction parallel to the Themed fault (Tethyan trend), which juxtaposes the volcanic units against the Late Cretaceous Sudr Chalk Formation. The ridges attain a height of 928 m (a. s. l.), about 160 m above the surrounding plain (Azer 2004). The volcanic rocks at Biq’at Hayareah include sequences of late Precambrian alkaline rhyolites (lavas, tuffs and ignimbrites) and their hypabyssal equivalents (rhyolite porphyries). In addition, quartz trachytes and amygdaloidal quartz trachytes define an E-W trending belt, up to 0.75 km wide, surrounded by rhyolites in the eastern and western parts of the area. The rhyolite yielded a Rb–Sr isochron age of 548 ± 5 Ma at Gebel El-Homra (Bielski 1982, in Bentor

Fig. 20.18 Geologic map showing the distribution of alkaline volcanics in the Biq’at Hayareah area (after Eyal et al. 2015)

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1985). The rhyolites can be correlated with those at Wadi Araba (SW Jordan) that yielded a Rb–Sr isochron age of 553 ± 11 Ma (Jarrar et al. 1992). Unfortunately, there are no available U–Pb zircon ages. Rhyolite lavas are the most widespread rocks in the area and are fine-grained with sparse phenocrysts of quartz, K-feldspar and sodic amphibole (Fig. 20.17c). The rhyolites are orange, brown or purple in color. Pyroclastic deposits are less abundant and occur as thin sheets (4–6 m thick) intercalated with the rhyolites. They include lapilli tuffs, banded tuffs and ash tuffs, together with ignimbrites. The subvolcanic rhyolite intrusions form separate hills in the southwestern part of the area and are strongly porphyritic and coarser-grained than the lavas. The quartz trachytes (both massive and amygdaloidal) are younger than the rhyolites, with chilled margins along their contacts. They show trachytic texture with sanidine phenocrysts (Fig. 20.17d). Thus, two volcanic events are now recognized in the Biq’at Hayareah area. Agron and Bentor (1981) related the quartz trachytes to the early Cretaceous volcanic rocks of the Ramon Province. They can be correlated with plutonic equivalents that constitute nearby Gebel Amram. At Gebel Amram, small elongate intrusions of monzonite and quartz syenite with an E-W orientation are found intruding the earlier rhyolites (Mushkin et al. 1999). The quartz syenite of Gebel Amram yielded a Rb–Sr isochron age of 526 ± 22 Ma (Mushkin et al. 1999). Geochemically, the rhyolitic lavas and the pyroclastic deposits at Biq’at Hayareah are alkaline in character, with the presence of sodic amphiboles supporting their peralkaline nature. The overall chemical characteristics of the volcanic units are consistent with a within-plate tectonic setting. These rocks erupted during a non-orogenic extensional period, shortly after the end of the collisional phase of the Pan-African orogeny, and represent a sudden and radical change from typical subduction-related calc-alkaline magmatism to post-tectonic alkaline (peralkaline) magmatism. Stabilization of the platform led to peneplanation and sedimentation of early Paleozoic sandstones at Taba–Nuweiba (Samuel et al. 2007). The rhyolitic lavas and pyroclastic deposits at Biq’at Hayareah are enriched in K2O (up to 10.1 wt%) and depleted in Na2O (around 0.08 wt%). The origin of these potash-enriched and soda-depleted rhyolites presents a challenge, and several hypotheses have been proposed. Primary magmatic processes, such as fractional crystallization or partial melting of K-rich sources, fail to explain the origin of these unusual rhyolites. They possess almost constant total alkali contents, but with potash appearing to increase at the expense of soda. Therefore, many authors favor either

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low- or high-temperature K-metasomatism. Agron and Bentor (1981) stated that about half of the alkali rhyolites (lavas, tuffs and ignimbrites) have undergone post-emplacement metasomatic alteration, expressed by exchange of K+ for Na+ and replacement of albite and primary quartz by K-feldspar. They concluded that potash metasomatism was the result of an interaction between the originally partly glassy rhyolites and hydrothermal solutions or, more probably, supercritical fluids. Low-temperature replacement of soda by potash by groundwater alteration is not feasible, since the rhyolitic rocks are always fresh and no adularia has ever been detected. Yet, high-temperature metasomatism also fails to explain the exchange of K2O for Na2O, because there is no sign of formation of alteration phases; there are only feldspars and quartz. The salient features of the potash-enriched, ultrapotassic rhyolitic rocks in the Biq’at Hayareah area are: (1) They are associated with abundant pyroclastic deposits (suggesting an important role for the accumulation and escape of gases); (2) the extreme depletion in Na2O is accompanied by an increase in normative corundum; (3) the amount of K2O approaches the amount of total alkalis in the normal rhyolites; (4) the K-feldspars are poor in albite, indicating that they crystallized from a magma already deficient in soda. A feasible mechanism for the genesis of the Biq’at Hayareah ultrapotassic rhyolitic rocks depends on the fact that Na2O dissolves more readily than K2O in supercritical hydrous fluids, which may accumulate in the upper part of a magma chamber and lead to explosive eruption and the formation of pyroclastic material. Given ample time for the volatiles to accumulate, the greater part of the Na2O (and a smaller part of K2O) is lost to the accumulating gases (Samuel et al. 2007), which subsequently escape. Alumina remaining in the magma cannot combine to form albite (due to the loss of soda) or muscovite (due to deficiency in H2O) and so potash still available in the deeper levels of the magma chamber forms more K-feldspars. The increased normative corundum reflects the retention of alumina as potash and soda are lost to the fluid phase. This adequately explains the increased potash content (seemingly at the expense of lost soda) in lavas that coexist with normal K2O/Na2O rocks in parts of the system where accumulation of volatiles and pyroclastic explosions did not occur.

20.8.2.4 Elat Volcanics The Precambrian volcanic rocks of the Elat region belong to the Katharina Group, as defined in southern Sinai, which is part of the alkaline magmatic stage (Gass 1977; Bentor 1985; Eyal et al. 1980; Bentor and Eyal 1987). The country rocks of the Elat Volcanics include the Taba quartz dioritic

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Fig. 20.19 Geologic map of the Elat area (after Peltz and Eyal 1994)

gneiss in the south, the Elat granitic gneiss in the east and the Elat granite in the northwest (Peltz and Eyal 1994). Most of the Elat Volcanics are covered by Neogene to recent fluviatile conglomerates and sandstones, but they are well exposed in the Ramat Yotam, Hatabor Hanozel, Container Farm and Mizpe Elat areas (Fig. 20.19). The Garof Tuff of the Yotam Caldera is a member of the post-accretionary Elat Volcanics, erupted in the latest Precambrian (Peltz and Eyal 1994). This volcanic suite is located in the southern part of the Ramat Yotam Plain, about 2 km west of the town of Elat. It consists of a sequence of densely welded silicic ash-flow tuffs (ignimbrites) of dacitic or rhyodacitic composition (Eyal et al. 1990; Eyal and Peltz 1992; Peltz and Eyal 1994). The remnants of the Garof Tuff are about 200 m thick and are entirely restricted to the inside

of the caldera (Peltz and Eyal 1994). A Rb–Sr isochron yielded an age of 545 ± 5 Ma for this tuff (Bielski 1982). The volcanic sequence of the Yotam Caldera consists of 17 cooling units and 23 ash flows. The various units all have rhyolitic–rhyodacitic compositions, and the major difference between the units is the degree of welding, crystallization, alteration, or amount of fiamme and lithic fragments. The basal layers are densely welded tuffs that are fiamme-poor and exhibit a fluidal planar structure. The central layers are also composed of densely welded tuff, but with fluidal to eutaxitic structures and more abundant fiamme and lithic fragments. Tuffs in the upper layers of the Yotam Caldera are rich in vesicles and/or open cavities and densely welded, fiamme-rich and lithic-poor, with a eutaxitic planar structure. They are also highly devitrified and altered (Peltz and Eyal 1994).

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Fig. 20.20 Geological map of central Wadi Araba, showing the distribution of the Wadi Araba Volcanics (Fig. 2b: Jarrar and Ghanem, this volume)

The Timna igneous complex, about 10 km to the north of the Amram–Neshef volcanic field, includes many small, shallow intrusive bodies that have textures similar to subvolcanic intrusions (Eyal et al. 2019). In the Elat area and southern Negev of Israel, the volcano-sedimentary sequence comprises coarse-grained clastic sediments, collectively known as the Elat conglomerate, and associated volcanic

members (Garfunkel 1980). This sequence overlies 630– 605 Ma granitoids (Be’eri-Shlevin et al. 2009a; Morag et al. 2011) and is crosscut by late rhyolitic dykes dated at 590– 585 Ma (Katzir et al. 2007; Morag et al. 2011). In the southern Negev, the alkaline volcanic units are unconformably overlain by early Paleozoic sandstones (Parnes 1971).

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Fig. 20.21 Enlarged geological map of the southern part of central Wadi Araba (Powell et al. 2015)

20.8.2.5 The Aheimir Volcanic Suite in Jordan The Wadi Araba Volcanics are bimodal and were generated after development of a regional unconformity and deposition of the Saramuj conglomerate. The Wadi Araba Volcanics are exposed at several localities, including Wadi Museimir, Wadi Abu Sakakin and Wadi Bourwas. They span the time interval 550–540 Ma (Jarrar et al. 1992). The Wadi Araba

Volcanics (Figs. 20.20 and 20.21) crop out in large masses and stocks, forming a 70 km NNE–SSW trending belt that extends from Feinan in the north to Gharandal in the south. The maximum width of the belt rarely exceeds 3 km (Jarrar et al. 1992). The volcanic rocks of Wadi Araba include rhyolitic lava flows, with subordinate trachybasalt and trachyandesite. In addition, subvolcanic activity is represented

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by numerous dykes and subordinate sills. The rhyolites are classified as alkali rhyolites and comendites. The rhyolitic lava flows exhibit various flow textures, whereas trachyandesites, andesites and basaltic trachyandesites display vesicular and amygdaloidal textures (Al Bakri 1990). The basalts are mantle-derived, whereas the alkali rhyolites and comendites were generated by partial melting of continental crust (Jarrar et al. 1992). The most famous succession of Wadi Araba Volcanics occurs at Wadi Museimir, which is located *65 km north of Aqaba. The exposed thickness of the Wadi Museimir succession is approximately 300 m and consists of the following units from base to top (Jarrar et al. 1992): (1) agglomeratic tuff and volcanic breccia; (2) laminated rhyolitic lava flows with dolomite and calcite-bearing spherules; (3) a finely laminated rhyolite with flattened lenses filled with chalcedony; (4) massive pink-colored rhyolite; (5) violet-reddish, massive, salt-bearing tuff; and (6) laminated spherulitic rhyolite. The initial 87Sr/86Sr ratio of the alkali feldspar rhyolites at Wadi Museimir is 0.7123 ± 0.0028 (Jarrar 1992), indicating that sialic crust participated in the generation of the magma. The peralkaline rhyolites (comendites) are distinctly porphyritic and exposed in the northern part of the Wadi Araba volcanic belt (at Wadi Abu Sakakin and Wadi Bourwas). SIMS data for zircons from these rhyolites record an age of 598 ± 3 Ma (Jarrar et al. 2012). The green color of the comendites is due to the occurrence of tiny crystals of aegirine. In a few occurrences (i.e., Wadi Kusheiba and Wadi Abu Barqa), ignimbrites (*6 m thick) of rhyolitic composition are present with a densely welded texture. The plano-linear arrangement of the glass shards in the ignimbrites is evidence of welding. The largest exposure of volcanic rock with mafic to intermediate composition is exposed as a 5 km by 3 km porphyrite stock at Feinan. Abundant subparallel NE–SW dyke swarms are associated with the Wadi Araba Volcanics (Bender 1974); these are the so-called Ghuweir Volcanics (Jarrar et al. 2008). They occur either as simple or as composite dykes (Wachendorf et al. 1985). The composite dykes are best exposed in Wadi Rahma and Wadi Turban (Salameh 1987; Abdullah 1989). The simple dykes are 2–15 m wide and up to a few hundred meters in length, whereas the composite dykes are up to 30 m wide and can be traced for several kilometers. The composition of the dykes varies from basalt to trachyandesite and rhyolite, with rare dacite and lamprophyre (spessartite). Ghanem et al. (2020) dated a composite dyke with latite margins and a rhyolite core (607 ± 6 Ma, U–Pb), a biotite rhyolite dyke (600 ± 4 Ma, 40Ar/39Ar age of

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biotite), an andesite dyke (594 ± 3 Ma, 40Ar/39Ar age of amphibole) and a dolerite dyke (*579 Ma, 40Ar/39Ar whole -rock total gas age). The rhyolites of Wadi Araba are ultrapotassic (Wachendorf et al. 1985; Jarrar 1992), similar to the alkaline/peralkaline Katharina Volcanics in the neighboring areas of Sinai and Elat (Agron and Bentor 1981; Mushkin et al. 1999; Samuel et al. 2007). Jarrar et al. (1992) described both normal and ultrapotassic alkali rhyolites in Wadi Museimer, central Wadi Araba, and indicated that these rhyolites exhibit strong chemical similarity and probable genetic relationship to those at Biq’at Hayareah. Wachendorf et al. (1985) investigated the late Proterozoic dykes that represent the latest magmatic activity of the Pan-African orogeny in southwest Jordan. Among these dykes, two high-level near-surface alkali rhyolitic dykes at Wadi Umm-Rachel and Wadi Kusheiba are characterized by extremely high K2O/Na2O ratios (35 and 47, respectively). They concluded that these two dykes, with an originally hyaline texture, were subjected to K-metasomatism, enhanced by continuous hydrothermal activity during pressure release. They also noted that the high potassium content of the near-surface dykes grades downward to normal K/Na ratios with increasing depth. Furthermore, a peculiar geochemical feature of the devitrified rhyolites of the Aheimir Volcanic Suite is the extreme K-enrichment and Na depletion (K/Na *1000; Jarrar and Ghanem, this volume).

20.9

Geodynamic Significance of ANS Post-accretionary Volcano-Sedimentary Successions

The distribution in space, time and composition of post-accretionary volcanic rocks and intercalated sedimentary members provides insight not only for the final Pan-African events in the northernmost ANS but also for earlier events. Detrital components in the sedimentary units include both pre-ANS (870–1100 Ma) and Pan-African (600–870 Ma) material (Avigad et al. 2007; Bea et al. 2009; Stern et al. 2010a; Be’eri-Shlevin et al. 2009b, 2011; Samuel et al. 2011). The presence of latest Mesoproterozoic to earliest Neoproterozoic (1.1–0.9 Ga) clastic material that eroded into post-accretionary basins indicates that pre-ANS crust of this age was more widespread in the ANS than has previously been appreciated. This in turn has significant implications for the early stages of construction of the northern ANS. In fact, 1.1–0.9 Ga zircons are also found as inherited xenocrysts in an *850 Ma intrusion now engulfed by the Katharina Ring Complex (e.g., Bea et al. 2009). One

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plausible explanation for the occurrence of such pre-ANS zircons in post-accretionary volcanic rocks is earlier transport of sediments hundreds of kilometers from their source regions via glacial processes. Such material could then be sampled at surficial levels by ca. 850 Ma lavas and could also be later reworked into ca. 620–590 Ma sediments. However, this idea suffers from a number of serious shortcomings. First, if glacial transport were to account for the occurrence of pre-Pan-African zircons in these rocks, older components (Archean to Paleoproterozoic) should be expected as well, such as those documented in diamictites of the central Eastern Desert of Egypt (Ali et al. 2009b). Second, Bea et al. (2009) found 1050 Ma zircons in an intrusive quartz diorite, not a volcanic or even hypabyssal unit. Third, some of the 1.1–1.0 Ga zircons from the Sa’al schist in south Sinai are euhedral, leading Be’eri-Shlevin et al. (2009b) to argue that crustal tracts of this age must have been located very close to the ANS when it formed. The results of Bea et al. (2009), Stern et al. (2010a) and Be’eri-Shlevin et al. (2011) strongly support this assertion. Although no in situ crust of 1.1–0.9 Ga has been found, such crustal components could have formed the substrate within which early magmatism in the northernmost ANS evolved (e.g., Be’eri-Shlevin et al. 2010). Many of the post-collisional calc-alkaline volcanic rocks of the ANS have an adakitic affinity (Thiéblemon 2000; Teklay et al. 2001; Hargrove et al. 2006; Eliwa et al. 2006; Be’eri-Shlevin et al. 2011; Obeid and Azer 2015; Abdelfadil et al. 2018). Adakitic rocks have been reported in quartz diorite–granite associations in southern Israel (Bogoch et al. 2003; Katz et al. 2004), in post-collisional calc-alkaline volcanic rocks in south Sinai (Be’eri-Shlevin et al. 2011; Abdelfadil et al. 2018), in Dokhan Volcanics in the Eastern Desert of Egypt (Eliwa et al. 2006; Obeid and Azer 2015) and in post-accretionary units in Saudi Arabia (Hargrove et al. 2006; Cox et al. 2019). There are no direct U–Pb ages for post-accretionary eruptive adakites, but they would have formed between 630 and 590 Ma (e.g., Wilde and Youssef 2001; Breitkreuz et al. 2010; Be’eri-Shlevin et al. 2011; Abdelfadil et al. 2018). Aside from what these adakitic rocks may tell us about the tectonic evolution of the shield, both the volcanic and plutonic outcrops of adakitic rocks are of considerable interest as proven hosts of gold, silver, base metal, molybdenum and tungsten mineralization (Pease and Johnson 2013). The origin of adakitic rocks has been debated extensively over the last few decades, and many models for their formation have been suggested (Castillo 2006). Although the classical interpretation of adakites put forward by Defant and Drummond (1990) invokes slab melting in an active

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subduction zone, it is now widely acknowledged that this origin is not unique. Rather, published data on post-accretionary adakitic volcanic rocks of the ANS simply require derivation from a source with garnet and possibly amphibole in the residue (Be’eri-Shlevin et al. 2011). In addition, the high Cr and Ni contents of the andesites require interaction with mantle peridotite. A model involving melting of subducted oceanic crust or variations of this model with melting of a hot oceanic ridge were suggested for the adakitic rocks of southern Israel and the post-accretionary adakitic volcanic rocks of the Eastern Desert of Egypt (e.g., Bogoch et al. 2003; Katz et al. 2004; Eliwa et al. 2006). However, in a post-collisional setting, subduction-based models for the formation of adakitic rocks are questionable; if adakites require active subduction, they would be the only evidence for subduction in the region at this time and would conflict with considerable accumulated evidence to the contrary. Eliwa et al. (2006) recognized this problem and suggested that a hot oceanic ridge that had been subducted (pre- or syn-collision) was melted some 20 m.y. later in a post-collisional regime. The lack of evidence for oceanic crust younger than ca. 690 Ma in the northern ANS (Stacey et al. 1984) does not support this view, even if the various adakitic post-accretionary volcanic series are all assumed to be as old as ca. 630 Ma, which is probably not the case. Eliwa et al. (2006) used convincing lines of evidence to preclude both *10% melting of LREE-enriched garnet peridotite (Stern and Gottfried 1986) and melting of thick lower crust as potential models for the post-accretionary adakitic volcanic rocks. Melting of mantle peridotite (*10%) would produce basalts, whereas lower crustal melts would produce andesites only if the temperatures were extreme (>1000 °C). Also, in both cases the high MgO, Cr and Ni content of these magmas and the lack of correlation of these indicators with their silica content are problematic. The compatible element concentrations are too high to be produced by melting of the mafic lower crust, and fractionation trends of mantle-derived basaltic magmas would be expected to reduce these values (Be’eri-Shlevin et al. 2011). A simpler model that would explain the late Neoproterozoic adakitic rocks of the ANS in general, and the post-accretionary volcanism in particular, is melting of delaminated mafic lower crust. Heating of delaminated mafic lower crust in the mantle would explain the high degrees of melting needed to produce andesites (requiring up to 25% melting of eclogite; e.g., Stern and Gottfried 1986), and interaction of the melts with mantle peridotite as they ascend would explain the high MgO, Cr and Ni contents (Kay and Kay 1993). Independent geodynamic considerations supporting delamination in the ANS at this time have been

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recently discussed by many authors (Avigad and Gvirtzman 2009; Moghazi 2003; El-Bialy 2010; Eyal et al. 2010; Obeid and Azer 2015; Abdelfadil et al. 2018). The melting of delaminated lower crust is more consistent with the tectonic setting of the post-accretionary adakitic volcanic sequences than the oceanic ridge model of Eliwa et al. (2006). There are, of course, fundamental differences between these models. They differ in the expected timing of the events that triggered melting;  630 Ma for the subducted ridge model (Avigad and Gvirtzman 2009) versus 610–600 Ma for delamination (Moghazi 2003; Eyal et al. 2010). They differ also in the source components; a broken oceanic slab in one model (Moghazi 2003) versus mafic lower continental arc crust in the other (Avigad and Gvirtzman 2009; Eyal et al. 2010). Focusing on the events recorded in Sinai and southern Israel, we note that the presence of adakitic rocks would favor an early phase of delamination (e.g., Avigad and Gvirtzman 2009), yet large areas of the subjacent lithospheric mantle remained intact. Avigad and Gvirtzman (2009) modeled the next stage in the evolution of the lithosphere as cooling and stabilization of a new lithospheric mantle root. This model, however, offers no explanation for the initiation of the second phase of magmatism with dominantly calc-alkaline character, followed (with some temporal overlap) by widespread, but volumetrically minor, alkaline magmatism. The available data are more consistent with the culmination of delamination at 610 Ma (e.g., Eyal et al. 2010), some 20 m.y. after its initiation. During this stage, upwelling of large portions of asthenospheric mantle would not only cause melting in the remaining lithospheric mantle and in the lower crust, but also trigger an extensional tectonic regime that would open conduits for small-degree alkaline magmas as well as basins to provide accommodation space for volcano-sedimentary accumulation. This model is in good agreement with the record of peak extensional tectonics in the northernmost ANS at ca. 600 Ma (Stern et al. 1984; Garfunkel 1999).

20.10

Conclusions

The Arabian–Nubian Shield (ANS) documents major crustal accretion resulting from the collision between East and West Gondwana during the Pan-African orogeny. Its preservation and exposure provide a world-class example of the products of such events and a basis for testing models of crustal growth and development. The ANS crops out in the western Arabian Plate and the northeastern African Plate, exposing a collage of Neoproterozoic juvenile arcs, younger volcano-sedimentary successions, voluminous granitoid

intrusions and enclaves of pre-Neoproterozoic crust. One of the most striking features of the ANS is the abundance of post-collisional plutons and associated volcano-sedimentary sequences. The post-accretionary Neoproterozoic evolution (after ca. 635 Ma) of the ANS included the voluminous intrusion of granitoids and development of volcanosedimentary basins. Post-accretionary volcano-sedimentary rock associations are widespread in the ANS, but there has been significant controversy regarding their tectonic setting and their ages. In this chapter, we have sought to propose resolutions to some of these debates concerning the age relationships and stratigraphic significance of the post-accretionary sequences in the ANS. In part, the confusion has resulted from a lack of detailed fieldwork and of high-precision geochronology. Rocks of the post-collisional phase have U–Pb zircon ages extending from ca. 635 Ma to 580 Ma. The magmatic history of the post-accretionary stage of the ANS includes two major magmatic episodes: the calc-alkaline phase (635–610 Ma) and the alkaline phase (dominantly 610–580 Ma, with some earlier and later activity). Each of these two magmatic episodes includes plutonic rocks and their volcanic equivalents, and the difference in their character implies a changing geodynamic setting during post-collisional collapse of the orogen. The earlier succession is represented by high-K calc-alkaline volcanic units, whereas the later succession includes alkaline/peralkaline rock series. Both volcano-sedimentary successions are generally unmetamorphosed, except for a few cases where the earlier calc-alkaline volcanic units were weakly metamorphosed and deformed in proximity to later granitoid intrusions. The early post-accretionary volcano-sedimentary sequences commenced with the extrusion of medium- to high-K calc-alkaline volcanic rocks and deposition of immature clastic sediments. They evolved through several cycles of magmatism. Volcanic eruption in each cycle was accompanied by deposition of immature sediments. The volcanic rocks of the early post-accretionary period include intermediate to felsic subaerial lava flows, tuffs, ignimbrites, subvolcanic intrusions and minor basalt. The end of the early post-accretionary stage at *610 Ma is marked by injection of numerous dyke swarms of various compositions. The high-K calc-alkaline character and other traits previously interpreted to indicate arc magmatism in the early post-accretionary stage more likely reflect remelting of arc-related material from earlier tectonic phases in the ANS. Some of the more evolved members of the calc-alkaline volcanic series are transitional to A-types, but this is likely due to extensive fractionation rather than a change in the tectonic regime at this time.

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The clasts in conglomerate units intercalated with the calc-alkaline phase of post-accretionary volcanism include schist, gneiss and syn-contemporaneous granitoids. This strongly suggests major uplift and denudation starting at ca. 630 Ma. The variation in clast composition between the older and younger conglomerates, with dominance of granitoids in the latter, also shows that rapid exhumation continued at least until ca. 600–590 Ma. The preferred model for the presence of adakitic rocks in the early pulse of post-accretionary magmatic activity is delamination and melting of mafic lower crust, perhaps including delamination of the subcontinental lithosphere, which would be entirely consistent with protracted isostatic uplift of the remaining crust. The later cycle of alkaline/peralkaline volcanic activity coincided with the final stage of post-collisional magmatism in the Pan-African orogeny, extending from the Ediacaran (*610 Ma) to the early Cambrian (*541 Ma). Volcanic products of this cycle include alkali rhyolite, comendite and pantellerite flows, along with abundant ignimbrites and pyroclastic deposits. They were erupted during a non-orogenic period in a state of tensile stress, block faulting and differential uplift. During the closing stages of this volcanic phase, large-scale caldera collapse occurred and ring dykes were injected into the bordering fractures. The ages of the youngest calc-alkaline products are similar to, or slightly younger than, the earliest phases of alkaline volcanism. The same overlap has been noted between ages of calc-alkaline and alkaline plutons in the ANS. The implication is that there was a transition between the two phases and that two distinct regional magma sources were being tapped during this complex transitional period from calc-alkaline to alkaline A-type magmatism in the ANS. More isotopic studies, especially of the Sm–Nd, Lu–Hf and Re–Os systems, are needed to fully understand the evolution of the mantle source region for the post-accretionary volcanic series of the ANS. More, and more precise, geochronological constraints on the volcanic rocks are also needed. High-resolution SIMS zircon ages would be especially useful because this method best identifies inherited components, but this technique has rarely been applied to these volcanic rocks. Precise determination of the ages of post-accretionary volcanics in the ANS is also critical for studies of the associated sedimentary successions, which lack much biostratigraphic control; such studies would constitute a rich source of information on environmental changes throughout the Neoproterozoic. Economic considerations also require additional studies of the post-accretionary volcanic sequences of the ANS, especially the adakite occurrences and associated

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mineralization. There has been no substantial economic study of the ore potential of the adakite occurrences in the ANS, despite the common association between similar rocks and potentially economic Cu–Au mineralization worldwide. Acknowledgements The authors thank Qingshang Shi for help in redrafting the geological maps in this chapter. The authors are also grateful to Prof. Zakaria Hamimi for his kind invitation to submit this work and also for editorial handing, as well as the two known reviewers (Prof. Bernard Bonin and Prof. Ghaleb Jarrar) and an anonymous reviewer.

References Abd El-Rahman YA, Seifert T, Gutzmer J, Said A, Hofmann M, Gärtner A, Linnemann U (2017) The South Um Mongul Cu–Mo– Au prospect in the Eastern Desert of Egypt: from a mid-Cryogenian continental arc to Ediacaran post-collisional appinite-high Ba–Sr monzogranite. Ore Geol Rev 80:250–266 Abd El-Rahman YA, Anbar MA, Li XH, Li J, Ling XX, Wu LG, Masoud AE (2019) The evolution of the Arabian-Nubian Shield and survival of its zircon U–Pb–Hf–O isotopic signature: a tale from the Um Had Conglomerate, central Eastern Desert, Egypt. Precambr Res 320:46–62 Abdallah SE, Azer MK, El Shammari AS (2019) Petrological and geochemical evolution of Ediacaran rare-metal bearing A-type granites from Jabal Aja complex, northern Arabian Shield, Saudi Arabia. Act Geol Sin 94:743–762. https://doi.org/10.1111/17556724.13825 Abdel-Karim AA, Ali S, El-Awady A, Elwan W, Zaki MK, Akihiro T (2019) Mineral and bulk–rock chemistry of Shadli bimodal metavolcanics from Eastern Desert of Egypt: implication for tectonomagmatic setting and Neoproterozoic continental growth in the Arabian-Nubian Shield. Lithos 338:204–217 Abdel-Khalek ML, Takla MA, Sehim A, Hamimi Z, El Manawi AW (1992) Geology and tectonic evolution of Wadi Beitan area, southeastern Desert, Egypt. In: International conference on geology of the Arab World, pp. 369–394 Abdel Khalek ML, Abdel Maksoud MA, Abdel Tawab MA, Oweiss KA (1994) Geochemistry of the St. Catherine basement rocks, Sinai, Egypt. Sci J Qatar Univ 14:134–145 Abdel Maksoud M, Khalek M, Oweiss K (1993) Geologic setting of the St. Catherine basement rocks, Sinai, Egypt. Sci J Qatar Univ 13:308–318 Abdel Wahed AAA, Ali KG, Khalil MMA, Abdel Gawad AE (2012) Dokhan volcanics of Gabal Monqul area, North Eastern Desert, Egypt: geochemistry and petrogenesis. Arab J Geosci 5:29–44 Abdelfadil KM, Obeid MA, Azer MK, Asimow PD (2018) Late Neoproterozoic adakitic lavas in the Arabian-Nubian shield, Sinai Peninsula, Egypt. J Asian Earth Sci 158:301–323 Abdel-Rahman A-FM (1996) Pan-African volcanism: petrology and geochemistry of the Dokhan Volcanic suite in the northern Nubian Shield. Geol Mag 133(1):17–31 Abdel-Rahman AM, Doig R (1987) The Rb–Sr geochronological evolution of the Ras Gharib segment of the northern Nubian shield. J Geol Soc Lond 144:577–586 Abdelsalam MG, Stern RJ (1996) Sutures and shear zones in the Arabian-Nubian Shield. J Afr Earth Sci 23(3):289–310

524 Abdullah NAT (1989) Geology, geochemistry and petrology of the Precambrian rocks of Wadi Rahma (Darba), Wadi Araba, SW Jordan. M.Sc. thesis. University of Jordan, Amman Abu El-Enen MM, Makroum FM (2003) Tectonometamorphic evolution of the northeastern Kid Belt, Southeast Sinai, Egypt. Ann Geol Surv Egypt VXXVI: 19–37 Abu El-Enen MM, Whitehouse MJ (2013) The Feiran-Solaf metamorphic complex, Sinai, Egypt: geochronological and geochemical constraints on its evolution. Precambr Res 239:106–125 Abu El-Enen MM, Lorenz J, Ali KA, von Seckendorff V, Okrusch M, Schüssler U, Brätz H, Schmitt RT (2018) A new look on Imperial Porphyry: a famous ancient dimension stone from the Eastern Desert of Egypt—petrogenesis and cultural relevance. Int J Earth Sci 107(7):2393–2408 Abu El-Leil I, Hassan MM, Abdel Tawab MM, Abdel Rahman HB (1990) Geological and geochemical studies on the Ferani Group; the proposed late Proterozoic younger volcanic rocks, Sinai, Egypt. Egypt Miner 2:61–80 Abu El-Leil I, Abdel Razek M (2001) A geological study on the basement rocks of Abu Durba-Abu Haswa area, South Western Sinai, Egypt. Ann Geol Surv Egypt 24:219–232 Abu El-Rus MA, Paquette JL, Chazot G, Vannucci R (2019) Bimodal zircon ages from Natash volcanics (southeast Egypt) and the link between eruption mechanisms and Late Cretaceous tectonics. Arab J Geosci 12(9):291 Abu-El Ela FF (2001) Neoproterozoic alkaline magmatism of the Eastern Desert of Egypt: example of Atalla felsites. In: 2nd international conference on the geology of Africa. Assiut University, Egypt, pp 255–291 Abul Gadayel A (1972) Contribution to geology and geochemistry of Wadi Natash lava flows. M.Sc. thesis. Faculty of Science, Cairo University, Egypt Abuamarah BA, Azer MK, Asimow PD, Shi Q (2021) Post-collisional volcanism with adakitic signatures in the Arabian-Nubian Shield: A case study of calc-alkaline Dokhan volcanics in the Eastern Desert of Egypt. Lithos 388–389:106051 Agar A (1986) The Bani Ghayy group; sedimentation and volcanism in pull-apart grabens of the Najd strike-slip orogen, Saudi Arabian Shield. Precambr Res 31:259–274 Agar RA (1988) Geologic map of the Zalm quadrangle, sheet 22FG, Kingdom of Saudi Arabia. Saudi Arabian directorate general of mineral resources geologic map GM 89 Agar RA, Stacey JS, Whitehouse MJ (1992) Evolution of the southern Afif terrane—a geochronologic study. Directorate General of Mineral Resources Open-File Report DGMR-OF-10–15 Agron N, Bentor YK (1981) The volcanic massif of Biq'at Hayareah (Sinai-Negev): a case of potassium metasomatism. J Geol 89 (4):479–495 Ahmed EA, Khudeir AA, El Tokhi M (1989) Sedimentology and geochemical studies on the Hammamat Sediments at Wadi Queih area, Central Eastern Desert, Egypt. Bull Fac Sci Assiut Univ 18:1– 20 Akaad MK (1957) Petrography, structure and time relations of the Igla Formation, Eastern Desert. Egypt J Geol 1(2):93–102 Akaad MK (1959) Dykes, felsite intrusions and intrusion breccia at Wadi Igla, Eastern Desert. Egypt J Geol 3:89–105 Akaad MK, Abu El-Ela AM (2002) Geology of the basement rocks in the eastern half of the belt between latitudes 25° 30′ and 26° 30′ N Central Eastern Desert, Egyptian Geological Survey, Egypt, Paper no. 78, 118pp Akaad MK, El Ramly MF (1958) Seven new occurrences of the Igla formation in the Eastern Desert of Egypt. Geological Survey Egypt. Paper no 3, 37pp

M. K. Azer et al. Akaad MK, El-Ramly MF (1960) Geological history and classification of basement rocks of the Central Eastern Desert of Egypt. Geological Survey Egypt, Paper no 9, 24pp Akaad MK, Noweir A (1969) Lithostratigraphy of the Hammamat–Um Seleimat District, Eastern Desert, Egypt. Nature 223(5203):284 Akaad MK, Noweir AM (1977) The post Hammamat felsites and their age relation with the younger granites of Egypt. Proc Acad Sci 30:163–168 Akaad MK, Noweir AM (1980) Geology and lithostratigraphy of the Arabian desert orogenic belt of Egypt, between latitudes 25° 35′ and 26° 30′ N. Bull Inst Appl Geol (Jeddah, Saudi Arabia) 3:127– 134 Akawy A (2003) Fault characterization and paleostress analysis in the Precambrian rocks west of the Meatiq Dome, Central Eastern Desert, Egypt. Neues Jahrb Geol P-A 30:1–35 Akawy A (2007) Geometry and texture of quartz veins in Wadi Atalla area, Central Eastern Desert. J Afr Earth Sci 47:73–87 Al Bakri NM (1990) Geochemistry and petrography of the Feinan porphyrite stock, Central Wadi Araba, Jordan. M.Sc. thesis. University of Jordan, Amman Alaabed S, Amin BE (2010) Late Precambrian metavolcanics of Um Anab North Eastern Desert, Egypt, geochemistry and tectonic environment studies. Eur J Sci Res 42(3):507–524 Alaabed S, El-Tokhi M (2014) Mineralogical and geochemical aspects of the petrogenesis of Pan African Dokhan Volcanics at Esh El Mellaha Area-NE Desert, Egypt. Arab J Geo 7:5553–5568 Aleinikoff JN, Stoeser DB (1988) Zircon morphology and U–Pb geochronology of seven metaluminous and peralkaline post-orogenic granite complexes of the Arabian Shield, Kingdom of Saudi Arabia. US Geological Survey Ali K, Andresen A, Manton WI, Stern RJ, Omar SA, Maurice AE (2012) U–Pb zircon dating and Sr–Nd–Hf isotopic evidence to support a juvenile origin of the *634 Ma El Shalul granitic gneiss dome, Arabian–Nubian Shield. Geol Mag 149(5):783–797 Ali BH, Wilde S, Gabr M (2009a) Granitoid evolution in Sinai, Egypt, based on precise SHRIMP U–Pb zircon geochronology. Gondwana Res 15(1):38–48 Ali KA, Stern RJ, Manton WI, Kimura JI, Khamees HA (2009b) Geochemistry, Nd isotopes and U–Pb SHRIMP zircon dating of Neoproterozoic volcanic rocks from the Central Eastern Desert of Egypt: New insights into the  750 Ma crust-forming event. Precambr Res 171(1–4):1–22 Ali KA, Azer MK, Gahlan HA, Wilde SA, Samuel MD, Stern RJ (2010) Age constraints on the formation and emplacement of Neoproterozoic ophiolites along the Allaqi-Heiani Suture, South Eastern Desert of Egypt. Gondwana Res 8(4):583–595 Ali KA, Kröner A, Hegner E, Wong J, Li S-Q, Gahlan HA, El Ela FF (2015) U–Pb zircon geochronology and Hf–Nd isotopic systematic of WadiBeitan granitoid gneisses, South Eastern Desert, Egypt. Gondwana Res 27:811–824 Ali-Bik MW, El Rahim SH, Wahab WA, Abayazeed SD, Hassan SM (2017) Geochemical constraints on the oldest arc rocks of the Arabian-Nubian Shield: the late Mesoproterozoic to late Neoproterozoic (?) Sa’al volcano-sedimentary complex, Sinai, Egypt. Lithos 284:310–326 Andresen A, El-Rus MAA, Myhre PI, Boghdady GY, Corfu F (2009) U–Pb TIMS age constraints on the evolution of the Neoproterozoic Meatiq Gneiss Dome, Eastern Desert, Egypt. Int J Earth Sci 98:481–497 Andresen A, El-Enen MM, Stern RJ, Wilde SA, Ali KA (2014) The Wadi Zaghra metasediments of Sinai, Egypt: new constraints on the late Cryogenian-Ediacaran evolution of the northernmost Arabian-Nubian Shield. Int Geol Rev 56:1020–1038

20

Volcanism During the Post-accretionary Stage …

Aranda-Gómez JJ, Luhr JF, Housh TB, Connor CB, Becker T, Henry CD (2003) Synextensional Pliocene-Pleistocene eruptive activity in the Camargo volcanic field, Chihuahua, México. Geol Soc Am Bull 115(3):298–313 Asran AMH, Azer MK, Aboazom AS (2005) Petrological and geochemical investigation on Dokhan and felsite in Wadi Sodmein Area, Central Eastern Desert, Egypt. Egypt J Geol 49:1–20 Augland LE, Andresen A, Boghdady GY (2012) U–Pb ID-TIMS dating of igneous and metaigneous rocks from the El-Sibai area: time constraints on the tectonic evolution of the Central Eastern Desert. Egypt Int J Earth Sci 101:25–37 Avigad D, Gvirtzman Z (2009) Late Neoproterozoic rise and fall of the northern Arabian-Nubian shield: the role of lithospheric mantle delamination and subsequent thermal subsidence. Tectonophysics 477:217–228 Avigad D, Stern RJ, Beyth M, Miller N, McWilliams MO (2007) Detrital zircon U–Pb geochronology of Cryogenian diamictites and lower Paleozoic sandstone in Ethiopia (Tigrai): age constraints on Neoproterozoic glaciation and crustal evolution of the southern Arabian-Nubian Shield. Precambr Res 154(1–2):88–106 Azer MK (2004) Petrological and geochemical studies of some volcanic rocks in south Sinai, Egypt. Ph.D. thesis. Faculty of Science, Cairo University, 130pp Azer MK (2007) Tectonic significance of Late Precambrian calc-alkaline and alkaline magmatism in Saint Katherine area, South Sinai, Egypt. Geol Acta 5(3):255–272 Azer MK, Asimow PD (2021) Petrogenetic evolution of the Neoproterozoic igneous rocks of Egypt. In: The geology of the Egyptian Nubian Shield. Springer, Cham, pp 343–382 Azer MK, Farahat ES (2011) Late Neoproterozoic volcano-sedimentary succession of Wadi Rufaiyil, southern Sinai, Egypt: a case of transition from late- to post-collisional magmatism. J Asian Earth Sci 42:1187–1203 Azer MK, Stern RJ, Kimura J-I (2010) Origin of a late Neoproterozoic (605±13 Ma) intrusive carbonate–albitite complex in Southern Sinai, Egypt. Int J Earth Sci 99(2):245–267 Azer MK, El-Ela FA, Ren M (2012) The petrogenesis of late Neoproterozoic mafic dyke-like intrusion in south Sinai, Egypt. J Asian Earth Sci 54:91–109 Azer MK, Obeid MA, Ren M (2014) Geochemistry and petrogenesis of late Ediacaran (580–605 Ma) post-collisional alkaline rocks from Katharina Ring complex, south Sinai, Egypt. J Asian Earth Sci 93:229–252 Azzaz SA, Blasy M, Moharem A, Amer O (2015) Late Neoproterozoic Dokhan volcanics, around G. Esh, North Eastern Desert, Egypt: geochemistry and relation of K2O and Nb abundances with crustal thickness. Arab J Geosci 8:3551–3564 Baldridge WS, Eyal Y, Bartov Y, Steinitz G, Eyal M (1991) Miocene magmatism of Sinai related to the opening of the Red Sea. Tectonophysics 197:181–201 Basta F (1997) Petrology, geochemistry and Rb–Sr geochronology of Wad1 Meknas Volcantcs, South Eastern Sinai, Egypt. In: 3rd conference on geochemistry, Alexandria University, pp 171–l83 Basta EZ, Kotb H, Awadallah MF (1980) Petrochemical and geochemical characteristics of the Dokhan formation at the type locality, Jabal Dokhan, Eastern Desert, Egypt. Bull Inst Appl Geol King Abdul Aziz Univ Jeddah 3:121–140 Basta FF, Maurice AE, Bakhit BR, Ali KA, Manton WI (2011) Neoproterozoic contaminated MORB of Wadi Ghadir ophiolite, NE Africa: geochemical and Nd and Sr isotopic constraints. J Afr Earth Sci 59(2–3):227–242 Bates RL, Jackson JA (1984) Dictionary of geological terms, 3rd edn. American Geological Institute

525 Bea F, Abu-Anbar M, Montero P, Peres P, Talavera C (2009) The  844 Ma Moneiga quartz-diorites of the Sinai, Egypt: evidence for Andean-type arc or rift-related magmatism in the Arabian-Nubian Shield? Precambr Res 175(1–4):161–168 Be’eri-Shlevin Y (2008) The origin and evolution of Neoproterozoic magmatism in the northern Arabian-Nubian Shield (Sinai Peninsula, Egypt and southern Israel): evidence from stable and radiogenic isotope record. Unpublished Ph.D. thesis. Ben-Gurion University, Israel, p 134 Be’eri-Shlevin Y, Katzir Y, Whitehouse M (2009a) Post-collisional tectonomagmatic evolution in the northern Arabian-Nubian Shield (ANS): time constraints from ion-probe U–Pb dating of zircon. J Geol Soc (Lond) 166:71–85 Be’eri-Shlevin Y, Katzir Y, Valley JW (2009b) Crustal evolution and recycling in a juvenile continent: oxygen isotope ratio of zircon in the northern Arabian Nubian Shield. Lithos 107:169–184 Be’eri-Shlevin Y, Katzir Y, Valley JW (2009c) Crustal evolution and recycling in a juvenile continent: oxygen isotope ratio of zircon in the northern Arabian Nubian Shield. Lithos 107(3-4):169–184. Be’eri-Shlevin Y, Katzir Y, Blichert-Toft J, Kleinhanns I, Whitehouse M (2010) Nd–Sr–Hf–O isotope provinciality in the northernmost Arabian–Nubian Shield: implications for crustal evolution. Contrib Miner Petrol 160(2):181–201 Be’eri-Shlevin Y, Samuel MD, Azer MK, Rämö OT, Whitehouse MJ, Moussa HE (2011) The late Neoproterozoic Ferani and Rutig volcano-sedimentary successions of the northernmost Arabian-Nubian Shield (ANS): new insights from zircon U–Pb geochronology, geochemistry and O–Nd isotope ratios. Precambr Res 188:21−44 Be’eri-Shlevin Y, Eyal M, Eyal Y, Whitehouse MJ, Litvinovsky B (2012) The Sa’al volcano-sedimentary complex (Sinai, Egypt): a latest Mesoproterozoic volcanic arc in the northern Arabian Nubian Shield. Geology 40(5):403–406 Bender F (1974) Explanatory notes on the geological map of the Wadi Araba, Jordan: scale 1:100,000; 3 sheets. Geol Jahrb Reihe B 10:1–62 Bentor YK (1983) The Arabo-Nubian crust in the Late Proterozoic– From marine tholeiites to continental alkali-granites. Geol Soc Am Bull Abstr with Progr 15:525 Bentor YK (1985) The crustal evolution of the Arabo-Nubian Massif with special reference to Sinai Peninsula. Precambr Res 28:1–74 Bentor YK, Eyal M (1987) The geology of Sinai, its implication for the evolution of the Arabo-Nubian Massif, Jebel Sabbagh sheet. Israel Academy of Sciences and Humanities, Jerusalem, 484p Beyth M, Grunhagen H, Zilberfarb A (1978) An ultramafic rock in the Precambrian of eastern Sinai. Geol Mag 115(5):373–378 Beyth M, Stern RJ, Altherr R, Kröner A (1994) The late Precambrian Timna igneous complex, southern Israel: evidence for comagmatic-type sanukitoid monzodiorite and alkali granite magma. Lithos 31(3–4):103–124 Bezenjani RN, Pease V, Whitehouse MJ, Shalaby MH, Kadi KA, Kozdroj W (2014) Detrital zircon geochronology and provenance of the Neoproterozoic Hammamat Group (Igla Basin), Egypt and the Thalbah Group, NWSaudi Arabia: implications for regional collision tectonics. Precamb Res 245:225–243 Bielski M (1982) Stages in the evolution of the Arabian-Nubian Massif in Sinai. Doctoral dissertation. Hebrew University of Jerusalem, 155p Blasband B, White S, Brooijmans P, De Boorder H, Visser W (2000) Late Proterozoic extensional collapse in the Arabian-Nubian shield. J Geol Soc 157(3):615–628 Bogoch R, Avigad D, Weissbrod T (2003) Geochemistry of the quartz diorite-granite association, Roded area, southern Israel. J Afri Earth Sci 35:51–60

526 Breitkreuz C, Eliwa H, Khalaf I, El Gameel K, Sergeev S, Larionov A, Hoffmann U (2008) Neoproterozoic SHRIMP U–Pb zircon ages of Dokhan Volcanics in the northern part of the Eastern Desert, Egypt. In: 33rd interational geological congress, Oslo, Norway (Abstract), 6–14 Aug 2008 Breitkreuz C, Eliwa H, Khalaf I, El Gameel K, Bühler B, Sergeev S, Larionov A, Murata M (2010) Neoproterozoic SHRIMP U–Pb zircon ages of silica-rich Dokhan Volcanics in the North Eastern Desert, Egypt. Precambr Res 182:163–174 Brown GF, Schmidt DL, Huffman AC Jr (1989) Geology of the Arabian Peninsula; shield area of western Saudi Arabia. U.S. geological survey professional paper 560-A, 188p Bühler B, Breitkreuz C, Pfänder JA, Hofmann M, Becker S, Linnemann U, Eliwa HA (2014) New insights into the accretion of the Arabian-Nubian Shield: depositional setting, composition and geochronology of a Mid-Cryogenian arc succession (North Eastern Desert, Egypt). Precambr Res 243:149–167 Calvez JV, Delfour J, Feybesse JL (1985) 2000-million-year old inherited zircons in plutonic rocks from the Al Amar region: New evidence for an Early Proterozoic basement in the eastern Arabian Shield? Saudi Arabian Deputy Ministry for mineral resources open file report BRGM-OF-05–11, p 28 Castillo PR (2006) An overview of adakite petrogenesis. Chin Sci Bull 51:257–268 Castillo PR (2012) Adakite petrogenesis. Lithos 134:304–316 Castillo PR, Janney PE, Solidum RU (1999) Petrology and geochemistry of Camiguin Island, southern Philippines: insights to the source of adakites and other lavas in a complex arc setting. Contrib Mineral Petrol 134(1):33–51 Cole JC, Hedge CE (1986) Geochronologic investigations of late Proterozoic rocks in the northeastern shield of Saudi Arabia: Saudi Arabian Deputy Ministry for mineral resources technical record USGS-TR-05-5, 42p Coleman RG, Gregory RT, Brown GF (1983) Cenozoic volcanic rocks of Saudi Arabia (Vol. 83, No. 788). US Department of the Interior, Geological Survey. Coulon C, Megartsi MH, Fourcade S, Maury RC, Bellon H, Louni-Hacini A, Cotten J, Coutelle A, Hermitte D (2002) Post-collisional transition from calc-alkaline to alkaline volcanism during the Neogene in Oranie (Algeria): magmatic expression of a slab break off. Lithos 62(3–4):87–110 Cox GM, Foden J, Collins AS (2019) Late Neoproterozoic adakitic magmatism of the eastern Arabian Nubian Shield. Geosci Front 10 (6):1981–1992 Cox GM, Lewis CJ, Collins AS, Halverson GP, Jourdan F, Foden J, Nettle D, Kattan F (2012) Ediacaran terrane accretion within the Arabian–Nubian Shield. Gondwana Res 21(2–3):341–352 Dardir AA, Abu Zeid KM (1972) Geology of the Basement rocks between latitudes 27° 00′ and 27° 30′ N, Eastern Desert. Ann Geol Surv Egypt II:129–159 Darwish ME, El-Tohamy AM (2014) Geochemical and mineralogical characterization of Natash Volcanics, South Eastern Desert, Egypt. Nucl Sci J 3(1):107–122 Defant MJ, Drummond MS (1990) Derivation of some modern arc magmas by melting of young subducted lithosphere. Nature 347:662–665 Defant MJ, Richerson PM, De Boer JZ, Stewart RH, Maury RC, Bellon H, Drummond MS, Feigenson MD, Jackson TE (1991) Dacite genesis via both slab melting and differentiation: petrogenesis of La Yeguada Volcanic Complex, Panama. J Petrol 32:1101– 1142 Dilek Y, Ahmed Z (2003) Proterozoic ophiolites of the Arabian Shield and their significance in Precambrian tectonics. In: Dilek Y,

M. K. Azer et al. Robinson PT (eds) Ophiolites in earth history. Geological society London special publications, vol 218, pp 685–700 Dimalanta CB, Yumul GP Jr (2008) Crustal thickness and adakite occurrence in the Philippines: is there a relationship? Island Arc 17 (4):421–431 Dixon TH (1979) The evolution of continental crust in the late Precambrian Egyptian Shield. Ph.D. thesis. California, San Diego University, 231p Dixon TH (1981) Gebel Dahanib, Egypt: a late Precambrian layered sill of komatiitic composition. Contrib Mineral Petrol 76:42–52 Doebrich JL, Al-Jehani AM, Siddiqui AA, Hayes TS, Wooden JL, Johnson PR (2007) Geology and metallogeny of the Ar Rayn terrane, eastern Arabian shield: evolution of a Neoproterozoic continental-margin arc during assembly of Gondwana within the East African orogen. Precambr Res 158:17–50 Downs DT, Stelten ME, Champion DE, Dietterich HR, Nawab Z, Zahran H, Hassan K, Shawali J (2018) Volcanic history of the northernmost part of the Harrat Rahat volcanic field, Saudi Arabia. Geosphere 14(3):1253–1282 El Habaak GH (2005) Petrogenesis and tectonic implications of the rock succession hosting banded iron formation at Um Anab Area, North Eastern Desert of Egypt. In: 4th international conference on the geology of Africa. Assiut-Egypt 2:479–513 El Kalioubi BA (1988) Deformation events, mineral facies and metamorphic condition in the contact aureoles of the Hammamat group around Umm Had pluton, Central Eastern Desert, Egypt. MERC Ain Shams Univ Earth Sci Ser 2:172–190 El Kalioubi BA (1996) Provenance, tectonic setting and geochemical characteristics of the Hammamat molasses sediments around Umm had pluton, Eastern Desert, Egypt. MERC Ain Shams Univ Earth Sci Ser 10:75–88 El Kalioubi BA, Osman AF (1996) Petrological and structural aspects of the mylonitized felsite at Wadi Atalla, Central Eastern Desert, Egypt. Proc Geol Surv Egypt, 159–184 El-Araby A, Abdel-Motelib A (1999) Depositional facies of the Cambrian Araba formation in the Taba region, east Sinai, Egypt. J Afr Earth Sci 29(3):429–447 El-Bialy MZ (1999) Geology of the basement rocks of Abu Durba–Abu Haswa area, southwestern Sinai, Egypt. Unpublished M. Sc. thesis, Suez Canal University, Ismailia, Egypt, 189p El-Bialy MZ (2010) On the Pan-African transition of the Arabian-Nubian Shield from compression to extension: the post-collision Dokhan volcanic suite of Kid-Malhak region, Sinai, Egypt. Gondwana Res 17:26–43 El-Bialy MZ (2020) Precambrian basement complex of Egypt. In: Geology of Egypt. Springer Nature, Switzerland, pp 37–79 El-Bialy MZ, Ali KA (2013) Zircon trace element geochemical constraints on the evolution of the Ediacaran (600–614Ma) post-collisional Dokhan Volcanics and Younger Granites of SE Sinai, NE Arabian-Nubian Shield. Chem Geol 360–361:54–73 El-Bialy MZ, Hassen IS (2012) The late Ediacaran (580–590 Ma) onset of anorogenic alkaline magmatism in the Arabian-Nubian Shield: Katharina A-type rhyolites of Gabal Ma’ain, Sinai, Egypt. Precambr Res 216–219:1–22 El-Gaby S (1975) Petrochemistry and geochemistry of some granites from Egypt. Neues Jahrb Miner Abh 124:147–189 El-Gaby S (1983) Architecture of the Egyptian basement complex. In: Proceedings of the 5th international conference on basement tectonics, pp 1–8 El-Gaby S (1994) Geologic and tectonic framework of the Pan-African orogenic belt in Egypt. In: Proceedings of the 2nd international conference on geology of the Arab World. Cairo University, pp 3–17

20

Volcanism During the Post-accretionary Stage …

El-Gaby S (2007) Integrated classification and evolution of the Neoproterozoic Pan-African Belt in Egypt. In: 5th international conference on the geology of Africa, pp 143–154 El-Gaby S, El-Nady O, Khudeir A (1984) Tectonic evolution of the basement complex in the Central Eastern Desert of Egypt. Geol Rundsch 73(3):1019–1036 El-Gaby S, List FK, Tehrani R (1988) Geology, evolution and metallogenesis of the Pan-African Belt in Egypt. In: El-Gaby S, Greiling RO (eds) The Pan-African Belt of northeast Africa and adjacent areas. Braunschweig Viewig, pp 17−68 El-Gaby S, Khudeir A A, El-Taky M (1989) The Dokhan Volcanics at Wadi Queh area, Central Eastern Desert, Egypt. In: 1st conference on geochemistry. Alex Univeristy Egypt, pp 42−62 El-Gaby S, List FK, Tehrani R (1990) The basement complex of the Eastern Desert and Sinai. In: Said R (ed) The geology of Egypt. Balkema, Rotterdam, Brookfield, pp 175–184 El-Gaby S, Khudier AA, Abdel Tawab M, Atalla RF (1991) The metamorphosed volcano-sedimentary succession of Wadi Kid, southeastern Sinai, Egypt. Ann Geol Sur Egypt XVII:19−35 El-Gaby S, Khalaf IM, Eliwa HA, El-Miligy A, Gomaa RM (2002) The volcano-sedimentary successions of the Wadi Sa’al-Wadi Zaghra area, Southeastern Sinai Egypt. In: 6th international conference on geology Arab world, pp 25–44 Eliwa HA, Kimura J-I, Itaya T (2006) Late Neoproterozoic Dokhan Volcanics, North Eastern Desert, Egypt: geochemistry and petrogenesis. Precambr Res 151:31–52 Eliwa H, Breitkreuz C, Khalaf I, El Gameel K (2010) Depositional styles of Early Ediacaran terrestrial volcanosedimentary succession in the Gebel El Urf area, North Eastern Desert, Egypt. J Afri Earth Sci 57:328–344 Eliwa HA, Breitkreuz C, Murata M, Khalaf IM, Bühler B, Itaya T, Takahashi T, Hirahara Y, Miyazaki T, Kimura JI, Shibata T (2014a) SIMS zircon U–Pb and mica K–Ar geochronology, and Sr–Nd isotope geochemistry of Neoproterozoic granitoids and their bearing on the evolution of the north Eastern Desert, Egypt. Gondwana Res 25(4):1570–1598 Eliwa HA, El-Bialy MZ, Murata M (2014b) Edicaran post-collisional volcanism in the Arabian-Nubian Shield: the high-K calc alkaline Dokhan volcanics of Gabal Samr El-Qaa (592 ± 5 Ma), North Eastern Desert, Egypt. Precambr Res 246:180–207 El-Mahallawi MM (1999a) Geochemistry of rare earth elements in the post-Hammamat felsites from Gebel Atalla and Wadi Igla areas, central Eastern Desert, Egypt: implications for their origin and evolution. In: 4th conference on geochemistry. Alexandria University, Egypt, pp 137–147 El-Mahallawi MM (1999b) Petrology of some Dokhan volcanics from the north Eastern Desert of Egypt. In: GAWA4, international conference on geology of the Arab World. Cairo University, Egypt, pp 315−335 El-Mahallawi MM, Ahmed AF (1993) Petrography, geochemistry and petrogenesis of post-Hammamat felsites of Gebel Atalla and Wadi Igla areas, Eastern Desert, Egypt. Bull Fac Sci Assiut Univ 22:81– 111 El-Masry NN, Hegazi AM (2005) Finite strain analyses of conglomerates, Saint Catherine area, south Sinai, Egypt. Ann Geol Surv Egypt XXVII:181–194 El-Masry NN, El-Kaliouby BA, Khawasik SM, El-Ghawaby MA (1992) Reconsideration of the geologic evolution of Saint Catherine ring dyke, south Sinai. In: Proceedings of the 3rd conference on geology Sinai develop, Ismailia, pp 229–238 El-Nisr SA (1997) Late Precambrian volcanism at Wadi Allaqi, south Eastern Desert, Egypt: evidence for transitional continental arc/margin environment. J Afr Earth Sci 24:301–312

527 El-Ramly MF (1962) The absolute ages of some basement rocks from Egypt. United Arab Republic, Geological Survey, Paper, (15) El-Ramly MF (1972) A new geological map for the basement rocks in the Eastern and Southwestern desert of Egypt: Scale 1: 1,000,000. Annals of the Geological Survey of Egypt El-Sayed MM, Obeid MA, Furnes H, Moghazi AM (2004) Late Neoproterozoic volcanism in the southern Eastern Desert, Egypt: petrological, structural and geochemical constraints on the tectonic-magmatic evolution of the Allaqi Dokhan volcanic suite. Neues Jahrb Mineral Abh 180(3):261–286 El Shazly EM, Hashad AH, Sayyah TA, Bassyuni FA (1973) Geochronology of Abu Swayel area, south eastern Desert. Egypt J Geol 17:1–18 Essawy MA (1972) Petrogenesis of alkaline rhyolites and microgranites from Samadai-Tunduba area, Eastern Desert, Egypt. Ann Geol Surv Egypt II:239–250 Essawy MA, Abu Zeid KM (1972) Atalla felsite intrusion and its neighbouring rhyolitic flows and tuffs, Eastern Desert. Ann Geol Surv Egypt II:271–328 Eyal M, Hezkiyahu T (1980) Katharina pluton: the outline of a petrologic framework. Isr Jo Earth Sci 29:41–52 Eyal M, Peltz S (1992) Evolution of the Ramat Yotam Caldera. Isr Geol Soc Ann Meet Abst, pp 39–40 Eyal M, Bartov Y, Shimron AE, Bentor YK (1980) Sinai geological map, aeromagnetic map. Survey of Israel, scale: 1:500,000, 1 sheet Eyal M, Eyal Y, Peltz S (1990) The Precambrian magmatism in the Elat region. Isr Geol Soc Ann Meet, 42–46 Eyal M, Voznesensky V, Bentor Y (1994) The geological map of the Precambrian of South-Western Sinai, scale 1: 50,000. Geological Society of Israel Eyal M, Voznesensky V, Bentor Y (1995) The geological map of Katharina ring complex, 1: 50,000. Israel Geological Society Eyal M, Litvinovsky B, Jahn BM, Zanvilevich A, Katzir Y (2010) Origin and evolution of post-collisional magmatism: coeval Neoproterozoic calc-alkaline and alkaline suites of the Sinai Peninsula. Chem Geol 269:153–179 Eyal M, Be’eri-Shlevin Y, Eyal Y, Whitehouse MJ, Litvinovsky B (2014a) Three successive Proterozoic island arcs in the Northern Arabian-Nubian Shield: evidence from SIMS U–Pb dating of zircon. Gondwana Res 25:338–357 Eyal M, Zanvilevich AN, Litvinovsky BA, Jahn BM, Ye V, Be’eri-Shlevin Y (2014b) The Katharina ring complex (Sinai Peninsula, Egypt): Sequence of emplacement and petrogenesis. Am J Sci 314:462–507 Eyal Y, Eyal M, Calvo R, Hall JK, Rosensaft M (2015) Geological map of the Elat metamorphic belt, Taba area, NE Sinai, Egypt. Scale 1: 50,000. Ben-Gurion University of Negev and the Geological Survey of Israel Eyal Y, Eyal M, Calvo R, Hall JK (2016) Geological map of the Elat metamorphic Belt, NE Sinai, Egypt. 1: 50,000. Geological Survey of Israel and Ben-Gurion University of the Negev Eyal Y, Eyal M, Litvinovsky B, Jahn BM, Calvo R, Golan T (2019) The evolution of the Neoproterozoic Elat metamorphic Complex, northernmost Arabian-Nubian Shield: Island arc to syncollisional stage and post-collisional magmatism. Precambr Res 320:137–170 Farahat ES (2006) The Neoproterozoic Kolet Um Kharit bimodal metavolcanic rocks, south Eastern Desert, Egypt: a case of enrichment from plume interaction?: Int J Earth Sci 95:275–287 Farahat ES, Azer MK (2011) Post-collisional magmatism in the northern Arabian-Nubian Shield: the geotectonic evolution of the alkaline suite at Gebel Tarbush area, south Sinai, Egypt. Chem Erde 71:247–266

528 Fowler A, Osman AF (2013) Sedimentation and inversion history of three molasse basins of the western Central Eastern Desert of Egypt: implications for the tectonic significance of Hammamat basins. Gondwana Res 23:1511–1534 Fritz H, Abdelsalam M, Ali KA, Bingen B, Collins AS, Fowler AR, Ghebreab W, Hauzenberger CA, Johnson PR, Kusky TM, Macey P (2013) Orogen styles in the East African Orogen: a review of the Neoproterozoic to Cambrian tectonic evolution. J Afr Earth Sci 86:65–106 Frost BR, Frost CD (2008) A geochemical classification for feldspathic igneous rocks. J Petrol 49(11):1955–1969 Frost BR, Barnes CG, Collins WJ, Arculus RJ, Ellis DJ, Frost CD (2001) A geochemical classification for granitic rocks. J Petrol 42 (11):2033–2048 Furnes H, Shimron AE, Roberts D (1985) Geochemistry of Pan-African volcanic arc sequences in southeastern Sinai Peninsula and plate tectonic implications. Precambr Res 29(4):359–382 Gahlan H, Azer M, Khalil AES (2015) The Neoproterozoic Abu Dahr ophiolite, South Eastern Desert, Egypt: petrological characteristics and tectonomagmatic evolution. Mineral Petrol 109:611–630 Gahlan H, Azer M, Asimow P, Al-Kahtany K (2016) Late Ediacaran post-collisional A-type syenites with shoshonitic affinities, northern Arabian-Nubian Shield: a possible mantle-derived A-type magma. Arab J Geosci 9(12):603 Garfunkel Z (1980) Contribution to the geology of the Precambrian of the Elat area. Israel J Earth Sci 29(1–2):25–40 Garfunkel Z (1999) History and paleogeography during the Pan-African orogen to stable platform transition: Reappraisal of the evidence from the Elat area and the northern Arabian-Nubian Shield. Isr J Earth Sci 48:135–157 Gass IG (1977) The evolution of the Pan-African crystalline basement in N.E. Africa and Arabia. J Geol Soc Lond 134:129–138 Genna A, Guerrot C, Deschamps Y, Nehlig P, Shanti M (1999) Les formations Ablah d’Arabie Saoudite (datation et implication géologique). Comp Rend Acad Sci Paris 329:661–667 Ghanem H (2009) Petrology, geochemistry, and thermobarometry of the contact metamorphism Aureole in the Saramuj conglomerate. The University of Jordan, Southeast Dead Sea, Jordan, p 261 Ghanem H, Jarrar GH (2013) Geochemistry and petrogenesis of the 595Ma shoshonitic Qunai monzogabbro, Jordan. J Afr Earth Sci 88:1–14 Ghanem M, Dardir AA, Francis MH, Zalata AA, Abu Zeid KM (1973) Basement rocks in Eastern Desert of Egypt north of latitude 26 40′ N. Ann Geol Surv Egypt III:33–38 Ghanem H, McAleer RJ, Jarrar GH, Whitehouse M (2020) 40Ar/39Ar and U–Pb SIMS zircon ages of Ediacaran dykes from the Arabian-Nubian Shield of south Jordan. Precambr Res 20:105714 Gharib ME, Ahmed AH (2012) Late Neoproterozoic volcanics and associated granitoids at Wadi Ranga, south Eastern Desert, Egypt: a transition from subduction related to intra-arc magmatism. Lithos 155:236–255 Ghobrial MG, Lotfi M (1967) The geology of Gebel Gattar and Gebel Dokhan areas. Geological survey of Egypt special publication, paper no. 40, 26p Greenberg JK (1981) Characteristics and origin of Egyptian younger granites. Geol Soc Am Bull 92:224–232 Greene RC (1993) Stratigraphy of the Late Proterozoic Murdama Group, Saudi Arabia; U.S. Government Printing Office, Washington, DC, USA Greenwood HR, Hadley DG, Anderson RE, Fleck RJ, Shmidt DL (1976) Late Proterozoic cratonization in southwestern Saudi Arabia. Phil Trans Roy Soc Lond 280:517–527

M. K. Azer et al. Grothaus B, Eppler D, Ehrlish R (1979) Depositional environment and structural implications of the Hammamat Formation, Egypt. Ann Geol Surv Egypt IX: 564–590. Guo F, Fan W, Li C (2006) Geochemistry of late Mesozoic adakites from the Sulu belt, eastern China: magma genesis and implications for crustal recycling beneath continental collisional orogens. Geol Mag 143(1):1–3 Guo Z, Wilson M, Liu J (2007) Post-collisional adakites in south Tibet: products of partial melting of subduction-modified lower crust. Lithos 96(1–2):205–224 Habib ME, Ahmed AA, El Nady OM (1985) Two orogenies in the Meatiq area of the central Eastern Desert, Egypt. Precambrian Res 30(2):83–111 Hadley DG, Schmidt DL (1980) Sedimentary rocks and basins of the Arabian Shield and their evolution. In: Cooray PG, Tahoun SA (eds) Evolution and mineralization of the Arabian–Nubian Shield. Institute of Applied Geology Bulletin, vol 3. Jeddah, pp 25–50 Halpern M, Tristan N (1981) Geochronology of the Arabian-Nubian Shield in southern Israel and eastern Sinai. J Geol 89:639–648 Hargrove U, Stern R, Kimura J-I, Manton W, Johnson P (2006) How juvenile is the Arabian-Nubian Shield? Evidence from Nd isotopes and pre-Neoproterozoic inherited zircon in the Bi’r Umq suture zone, Saudi Arabia. Earth Planet Sci Lett 252(3):308–326 Harris NBW, Hawkesworth CJ, Ries AC (1984) Crustal evolution in northeast and east Africa from model Nd ages. Nature 309:773–776 Hashad AH (1980) Present status of geochronological data on the Egyptian basement complex. Bull Inst Appl Geol King Abdul Aziz Univ Jeddah 3:31–46 Hashad AH, El Reedy MWM (1979) Geochronology of the anorogenic alkalic rocks south Eastern Desert, Egypt. Ann Geol Surv Egypt IX:81–101 Hassan AM (1998) Geochemistry, Rb–Sr geochronology and petrogenesis of the late pan-African Atalla felsite, central Eastern Desert, Egypt. In: 4th international conference on geology of the Arab World Cairo University, pp 89–90 Hassan MA, Hashad AH (1990) Precambrian of Egypt. In: Said R (ed) Geology of Egypt. A.A Balkerna Publishers, The Netherlands, pp 201–245 Hassan AM, Ghoneim MAF, Abu Anbar MM (1997) Geochemistry and mineral chemistry of alkaline volcanics, Egypt. In: 3rd conference on geochemical Alexandria University Egypt, pp 203– 219 Hassan IS, Ali Bik MW, Gaal K (2001) Mineral chemistry of continental Mesozoic and Tertiary basaltic rocks, Um Bogma area, South Sinai: an integrated comparative approach. In: 2nd international Hegazy HK, Mawas SG (1992) Petrography, geochemistry and tectonic setting of late Proterozoic volcanic sequences of the area west of Ras Gharib, Eastern Desert, Egypt. Bull Fac Sci Assiut Univ 21 (1-f):31–56 Holail HM, Moghazi AKM (1998) Provenance, tectonic setting and geochemistry of greywackes and siltstones of the Late Precambrian Hammamat Group, Egypt. Sedim Geol 116(3–4):227–250 Hubbard HB (1981) Geochemical evolution of the Wadi Natash volcanic field: an alkaline basalt flood in the Eastern desert of Egypt. Doctoral dissertation. University of South Carolina Hume WF (1906) The topography and geology of the Peninsula of Sinai (South Eastern Portion). Geological Survey of Egypt, 280pp Jarrar GH (1985) Late proterozoic crustal evolution of the Arabian-Nubian Shield in Wadi Araba Area, SW-Jordan. Geol Jahrb B61:3–87

20

Volcanism During the Post-accretionary Stage …

Jarrar GH (1992) Geochemistry and petrogenesis of an alkali feldspar suite from Wadi Museimir, Central Wadi Araba, Jordan. Chemie der Erde 52:301–312 Jarrar GH, Dill H, Yaseen N, Whitehouse M (2012) The Aheimir Volcanic Suite: An Ediacaran post-collisional sequence from the northernmost Arabian-Nubian Shield, Wadi Arab, SW Jordan: Age, geochemistry and petrogenesis. In Abstract. North-Central Section-46th Annual Meeting of the Geological Society of America Jarrar G, Wachendorf H, Saffarini G (1992) A Late Proterozoic bimodal volcanic/subvolcanic suite from Wadi Araba, southwest Jordan. Precambr Res 56(1–2):51–72 Jarrar GH, Wachendorf H, Zachmann DG (1993) A Pan-African alkaline pluton intruding the Saramuj Conglomerate, South-West Jordan. Geol Rundsch 82:121–135 Jarrar G, Stern R, Saffarini G, Al-Zubi H (2003) Late-and post-orogenic Neoproterozoic intrusions of Jordan: implications for crustal growth in the northernmost segment of the East African Orogen. Precambr Res 123(2–4):295–319 Jarrar GH, Manton WI, Stern RJ, Zachmann D (2008) Late Neoproterozoic A-type granites in the northernmost Arabian-Nubian Shield formed by fractionation of basaltic melts. Chem Erde 68(3):295– 312 Johnson PR (2003) Post-amalgamation basins of the NE Arabian Shield and mplications for Neoproterozoic tectonism in the northern East African Orogen. Precambr Res 123:321–337 Johnson PR, Andresen A, Collins AS, Fowler AR, Fritz H, Ghebreab W, Kusky T, Stern RJ (2011) Late Cryogenian–Ediacaran history of the Arabian–Nubian Shield: a review of depositional, plutonic, structural, and tectonic events in the closing stages of the northern East African Orogen. J Afr Earth Sci 61(3):167–232 Johnson PR, Kattan F (2007) Geochronologic dataset for Precambrian rocks in the Arabian Peninsula. A catalogue of U–Pb, Rb–Sr, Ar– Ar, and Sm–Nd ages: Saudi geological survey open-file report SGS-OF-2007-3, 21pp Johnson PR, Kattan FH (2012) The geology of the Saudi Arabian Shield; Saudi Geological Survey: Jiddah, Saudi Arabia, pp 1–479 Johnson PR, Stewart ICF (1995) Magnetically inferred basement structure in central Saudi Arabia. Tectonophysics 245:37–52 Johnson PR, Woldehaimanot B (2003) Development of the Arabian-Nubian Shield: perspectives on accretion and deformation in the East African Orogen and the assembly of Gondwana. In: Yoshida M, Windley BF, Dasgupta S (eds) Proterozoic East Gondwana: supercontinent assembly and breakup. Geological Society London Special Publications, pp 289–325 Johnson PR, Andresen A, Collins AS, Fowler AR, Fritz H, Ghebreab W, Kusky T, Stern RJ (2011) Late Cryogenian–Ediacaran history of the Arabian–Nubian Shield: a review of depositional, plutonic, structural, and tectonic events in the closing stages of the northern East African Orogen. J Afr Earth Sci 61(3):167–232 Karsli O, DokuzA UI, Aydin F, Kandemir R, Wijbrans J (2010) Generation of the Early Cenozoic adakitic volcanism by partial melting of mafic lower crust, Eastern Turkey: implications for crustal thickening to delamination. Lithos 114:109–120 Katz O, Beyth M, Miller N, Stern R, Avigad D, Basu A, Anbar A (2004) A Late Neoproterozoic (630 Ma) Boninitic suite from southern Israel: implications for the consolidation of Gondwanaland. Earth Planet Sci Lett 218:475–490 Katzir Y, Eyal M, Litvinovsky B, Jahn B, Zanvilevich A, Valley J, Beeri Y, Pelly I, Shimshilashvili E (2007) Petrogenesis of A-type granites and origin of vertical zoning in the Katharina pluton, Gebel Mussa (Mt. Moses) area, Sinai, Egypt. Lithos 95(3–4):208–228

529 Kay SM, Jones HA, Kay RW (2013) Origin of Tertiary to Recent EM-and subduction-like chemical and isotopic signatures in Auca Mahuida region (37–38 S) and other Patagonian plateau lavas. Contrib Miner Petrol 166:165–192 Kennedy AK, De Laeter JR (1994) The performance characteristics of the WA SHRIMP II ion microprobe. In: Eighth international conference on geochronology, cosmochronology and isotope geology. U.S. geological survey circular, vol 1107, Berkeley, USA, p 166 (Abstracts) Kennedy A, Johnson PR, Kattan FH (2004) SHRIMP geochronology in the northern Arabian Shield, part I: data acquisition: Saudi geological survey. Open-file report SGS-OF-2004–11, 28pp Kennedy A, Johnson PR, Kattan FH (2005) SHRIMP geochronology in the northern Arabian Shield: part II: data acquisition 2004: Saudi geological survey. Open-file report SGS-OF-2005–10, 44pp Kennedy A, Kozdrój W, Kattan FH, Kozdrój MZ, Johnson PR (2010) SHRIMP geochronology in the Arabian Shield (Midyan Terrane, Afif Terrane, Ad Dawadimi Terrane) and Nubian Shield (Central Eastern Desert), part IV. Data acquisition 2008. Saudi geological survey open-file report SGS-OF-2010–10, 101pp Kennedy A, Kozdrój W, Johnson PR, Kattan FH (2011a) SHRIMP Geochronology in the Northern Arabian Shield. Part III. Data acquisition 2006. Saudi geological survey. Open-file report SGS-OF-2007–9, 81pp Kennedy A, Kozdrój W, Kadi K, Kozdrój MZ, Johnson PR (2011b) SHRIMP geochronology in the Arabian Shield (Midyan Terrane, Afif Terrane) and Nubian Shield (Central Eastern Desert Terrane), part V: data acquisition 2009. Saudi geological survey. Open-file report SGS-OF-2010–11, 80pp Khalaf EA (2002) Petrogenesis of Feirani volcanics, southeastern Sinai, Egypt. Egypt J Geol 46:183–211 Khalaf EEDAH (2012) Origin and evolution of post-collisional volcanism: an example from Neoproterozoic Dokhan volcanics at Gabal Nugara area, Northeastern Desert, Egypt. Arab J Geosci 5 (4):663–695 Khalaf IM, Dardir AA, El-Boughdady AA (1994) The basement rocks around Wadi Dahab, southeastern Sinai, Egypt. Egypt Miner 38 (1):35–54 Khalaf EE, Sano T, Tsutsumi Y (2018) Evolution of monogenetic rift-related alkaline magmatism in south Egypt: insight from stratigraphy, geochronology, and geochemistry of the Natash volcanics. J Afr Earth Sci 147:450–476 Khalil AES, Obeid MA, Azer MK, Asimow PD (2018) Geochemistry and petrogenesis of post-collisional alkaline and peralkaline granites of the Arabian-Nubian Shield: a case study from the southern tip of Sinai Peninsula, Egypt. Int Geol Rev 60(8):998–1018 Khudeir AA, Ali MM, El Habaak GH (1988) The metavolcanics at Um Samiuki area, Egypt. Bull Fac Sci Assiut Univ 17(2-F):73–101 Khudeir AA, Abu El-Rus MA, El-Gaby S, El-Nady O, Bishara WW (2008) Sr–Nd isotopes and geochemistry of the infrastructural rocks in the Meatiq and Hafafit core complexes, Eastern Desert, Egypt: evidence for involvement of pre-Neoproterozoic crust in the growth of Arabian-Nubian Shield. Island Arc 17(1):90–108 Kröner A (1984) Late Precambrian plate tectonics and orogeny: A need to redefine the term Pan-African. In: Michot J (ed) Klerkx, J. Geologie africaine, Tervuren, pp 23–28 Kröner A (1985) Ophiolites and the evolution of tectonic boundaries in the late proterozoic Arabian—Nubian shield of northeast Africa and Arabia. Precambrian Res 27(1–3):277–300 Kröner A, Stern, RJ (2005) Pan-African orogeny: encyclopedia of geology, vol 1.Elsevier, Amsterdam, pp 1–12

530 Kröner A, Greiling R, Reischmann T, Hussein IM, Stern RJ, Kruger J, Duur S, Zimmer M (1987) Pan-African crustal evolution in the Nubian segment of Northeast Africa. In: Kröner A (ed) Proterozoic lithosphere evolution. American Geophysical Union, Washington, DC, pp 235–257 Kröner A, Linnebacher P, Stern R, Reischmann T, Manton W, Hussein I (1991) Evolution of Pan-African island arc assemblages in the southern Red Sea Hills, Sudan, and in southwestern Arabia as exemplified by geochemistry and geochronology. Precambr Res 53 (1–2):99–118 Kröner A, Todt W, Hussein IM, Mansour M, Rashwan AA (1992) Dating of late Proterozoic ophiolites in Egypt and Sudan using the single grain zircon evaporation technique. Precambr Res 59:15–32 Kusky TM, Abdel Salam MG, Stern RJ, Tucker RD (2003) Evolution of the East African and related orogens, and the assembly of Gondwana. Precambr Res 123:81–85 Le Maitre RW, Bateman P, Dudek A, Keller J, Lameyre J, LeBas MJ, Sabine PA, Schmid R, Sorensen H, Streckeisen A, Woolley AR (1989) A classification of igneous rocks and glossary of terms: recommendations of the international union of geological sciences subcommission on the systematics of Igneous rocks. Blackwell, Oxford, p 193 Li ZX, Bogdanova SV, Collins AS, Davidson A, De Waele B, Ernst RE, Fitzsimons IC, Fuck RA, Gladkochub DP, Jacobs J, Karlstrom KE (2008) Assembly, configuration, and break-up history of Rodinia: a synthesis. Precambr Res 160(1–2):179–210 Li JW, Zhao XF, Zhou MF, Ma CQ, de Souza ZS, Vasconcelos P (2009) LateMesozoic magmatism from the Daye region, eastern China: U–Pb ages, petrogenesis, and geodynamic implications. Contrib Mineral Petrol 157:383–409 Li X-H, El-Rahman YA, Anbar MA, Li J, Ling X-X, Wu L-G, Masoud AE (2018) Old continental crust underlying juvenile oceanic arc: evidence from northern Arabian-Nubian shield. Egypt Geophys Res Lett 45(7):3001–3008 Liégeois JP, Stern RJ (2010) Sr–Nd isotopes and geochemistry of granite-gneiss complexes from the Meatiq and Hafafit domes, Eastern Desert, Egypt: no evidence for pre-Neoproterozoic crust. J Afr Earth Sci 57(1–2):31–40 Loizenbauer J, Wallbrecher E, Fritz H, Neumayr P, Khudeir AA, Kloetzli U (2001) Structural geology, simple zircon ages and fluid inclusion studies of the Meatiq metamorphic core complex: implications for Neoproterozoic tectonics in the Eastern Desert of Egypt. Precambr Res 110:357–383 Madbouly MI (1991) Petrology and geochemistry of some mafic ultramafic rocks of Sinai, Egypt. M.Sc. thesis. Cairo University, 132p Makovicky E, Frei R, Bailey JC, Karup-Moller S (2016) Imperial Porphyry from Gebel Abu Dokhan, the Red Sea Mountains, Egypt: part II. Geochem N Jb Miner Abh 193:29–44 Martin H (1999) Adakitic magmas: modern analogues of Archaean granitoids. Lithos 46(3):411–429 Martin H, Smithies RH, Rapp R, Moyen J-F, Champion D (2005) An overview of adakite, tonalite–trondhjemite–granodiorite (TTG), and sanukitoid: relationships and some implications for crustal evolution. Lithos 79:1–24 Maurice AE, Basta FF, Khiamy AA (2012) Neoproterozoic nascent island arc volcanism from the Nubian Shield of Egypt: magma genesis and generation of continental crust in intra-oceanic arcs. Lithos 132:1–20 Maurice AE, Bakhit BR, Basta FF, Asimow PD, Wälle M, Azer MK, El-Sobky AF (2018) The last subduction-related volcanism in the northern tip of the Arabian-Nubian Shield: a Neoproterozoic arc

M. K. Azer et al. preceding the terminal collision of East and West Gondwana. Precambr Res 310:256–277 McWilliams MO (1981) Palaeomagnetism and Precambrian tectonic evolution of Gondwana. In: Developments in Precambrian geology, vol 4. Elsevier, pp 649–687 Meert JG (2003) A synopsis of events related to the assembly of eastern Gondwana, Tectonophysics 362: 1−40 Mogahed MM (2020). Genesis and tectonic implications of Kabr El‐ Bonaya ultramafic rocks, Sinai Peninsula, Egypt: constraints from mineralogical and geochemical characteristics. Acta Geol Sin Engl Ed. https://doi.org/10.1111/1755-6724.14406 Moghazi AM (2003) Geochemistry and petrogenesis of a high-K calc-alkaline Dokhan Volcanic suite, south Safaga area, Egypt: the role of Late Neoproterozoic crustal extension. Precambr Res 125:161–178 Moghazi A-KM, Ali KA, Wilde SA, Zhou Q, Andersen T, Andresen A, Abu El-Enen MM, Stern RJ (2012) Geochemistry, geochronology, and Sr–Nd isotopes of the late Neoproterozoic Wadi Kid volcano-sedimentary rocks, Southern Sinai, Egypt: implications for tectonic setting and crustal evolution. Lithos 154:147–165 Mohamed FH (2001) The Natash alkaline volcanic field, Egypt: geochemical and mineralogical inferences on the evolution of a basalt to rhyolite eruptive suite. J Volcan Geoth Res 105(4):291–322 Mohammed FH, Moghazi AM, Hassanen MA (2000) Geochemistry, petrogenesis and tectonic setting of late Neoproterozoic Dokhan-type volcanic rocks in the Fatira area, Eastern Egypt. Int J Earth Sci 88:764–777 Morag N, Avigada D, Gerdesb A, Belousovac E, Harlavand Y (2011) Crustal evolution and recycling in the northern Arabian-Nubian Shield: new perspectives from zircon Lu–Hf and U–Pb systematics. Precambr Res 186:101–116 Morag N, Avigad D, Gerdes E, Harlavan Y (2012) 1000–580 Ma crustal evolution in the Northern-Arabian Shield revealed by U–Pb– Hf of detrital zircons from late Neoproterozoic sediments (Elat area, Israel). Precambr Res 208–211:197–212 Moreno JA, Montero P, Abu Anbar M, Molina JF, Scarrow JH, Talavera C, Cambeses A, Bea (2012) SHRIMP U–Pb zircon dating of the Katerina Ring Complex: Insights into the temporal sequence of Ediacaran calc-alkaline to peralkaline magmatism in southern Sinai, Egypt. Gondwana Res 12:887–900 Moreno J, Molina J, Montero P, Anbar MA, Scarrow J, Cambeses A, Bea F (2014) Unraveling sources of A-type magmas in juvenile continental crust: constraints from compositionally diverse Ediacaran post-collisional granitoids in the Katerina Ring Complex, southern Sinai, Egypt. Lithos 192:56–85 Moufti MR, Németh K (2013) The intra-continental Al Madinah volcanic field, Western Saudi Arabia: a proposal to establish Harrat Al Madinah as the first volcanic geopark in the Kingdom of Saudi Arabia. Geoheritage 5(3):185–206 Moufti MR, Németh K (2016) Harrat Rahat: the geoheritage value of the youngest long-lived volcanic field in the Kingdom of Saudi Arabia. In: Geoheritage of volcanic harrats in Saudi Arabia. Springer, Cham, pp 33–120 Moufti MR, Moghazi AM, Ali KA (2012) Geochemistry and Sr–Nd– Pb isotopic composition of the Harrat Al-Madinah volcanic field, Saudi Arabia. Gondwana Res 21(2–3):670–689 Moussa HE (2002) Mineral chemistry and geochemistry of some mafic-ultramafic intrusions in the South Eastern Desert and Sinai, Egypt. Egypt J Geol 46:213–238 Moussa HE (2003a) Low-pressure metamorphism in volcano-sedimentary succession of Gebel Ferani area, southeastern Sinai, Sinai, Egypt. Egypt J Geol 47(1):41–65

20

Volcanism During the Post-accretionary Stage …

Moussa HE (2003b) Geologic setting, petrography and geochemistry of the volcanosedimentary succession at Gebel Ferani area, southeastern Sinai, Egypt. Egypt J Geol 47:153–173 Moussa EM, Stern RJ, Manton WI, Ali KA (2008) SHRIMP zircon dating and Sm/Nd isotopic investigations of Neoproterozoic granitoids, Eastern Desert, Egypt. Precambr Res 160(3–4):341–356 Murcia H, Lindsay JM, Németh K, Smith IE, Cronin SJ, Moufti MR, El-Masry NN, Niedermann S (2017) Geology and geochemistry of Late quaternary volcanism in northern Harrat Rahat, Kingdom of Saudi Arabia: implications for eruption dynamics, regional stratigraphy and magma evolution. J Geol Soc Lond Spec Publ 446 (1):173–204 Mushkin A, Navon O, Halicz L, Heimann A, Woerner G, Stein M (1999) Geology and geochronology of the Amram massif, southern Negev Desert, Israel. Israel J Earth Sci 48:179–193 Obeid MA (2006) The Pan-African arc-related volcanism of the Wadi Hodein area, south Eastern Desert, Egypt: Petrological and geochemical constraints. J Afr Earth Sci 44(3):383–395 Obeid MA, Azer MK (2015) Pan-African adakitic rocks of the north Arabian-Nubian Shield: petrological and geochemical constrains on the evolution of the Dokhan volcanics in the north Eastern Desert of Egypt. Int J Earth Sci 104:541–563 Obeid MA, Khalil AES, Azer MK (2016) Mineralogy, geochemistry and geotectonic significance of the Neoproterozoic ophiolite of Wadi Arais area, south Eastern Desert, Egypt. Int Geol Rev 58:687– 702 Oweiss KA (1994) A short note on the banded iron formation (BIF) at Gabal Nakhila area, Southern Sinai, Egypt. Ann Geol Surv Egypt VXX:874–876 Parnes A (1971) Late lower Cambrian trilobites from the Timna area and Har Amram (southern Negev, Israel). Isr J Earth Sci 20(4):179– 205 Patchett PJ, Chase CG (2002) Role of transform continental margins in major crustal growth episodes. Geology 30:39–42 Pearce JA (1983) Role of the sub-continental lithosphere in magma genesis at active continental margins Pearce JA, Harris NB, Tindle AG (1984) Trace element discrimination diagrams for the tectonic interpretation of granitic rocks. J Petrol 25 (4):956–983 Pease V, Johnson PR (2013) Introduction to the JEBEL volume of Precambrian research. Precambr Res 239:1–5 Peltz S, Eyal M (1994) Ash Flows and Cooling Units Associated with the Garof Tuff Member: Ramat Yotam Caldera, Elat, Southern Israel. Ministry of Energy and Infrastructure, Geological Survey of Israel Powell JH, Abed A, Jarrar GH (2015) Ediacaran Araba Complex of Jordan. Geo Arabia 20(1):99–156 Ragab AI (1987) On the petrogenesis of the Dokhan Volcanics of the northeastern Desert, Egypt. MERC Ain Shams Univ Earth Sci Ser 1:151–158 Reischmann T, Kröner A (1994) Late Proterozoic island arc volcanics from Gebeit, Red Sea Hills, north-east Sudan. Geol Rundsch 83 (3):547–563 Ressetar R, Monrad JR (1983) Chemical composition and tectonic setting of the Dokhan Volcanic Formation, Eastern Desert, Egypt. J Afr Earth Sci 1:103–112 Reymer A, Schubert G (1984) Phanerozoic addition rates to the continental crust and crustal growth. Tectonics 3:63–77 Rickwood PC (1989) Boundary lines within petrologic diagrams, which use oxides of major and minor elements. Lithos 22:247–263 Ries AC, Shackleton RM, Graham RH, Fitches WR (1983) Pan-African structures, ophiolites and mélange in the Eastern Desert of Egypt: a traverse at 26° N. J Geol Soc Lond 140(1):75–95

531 Ringwood A (1990) Slab-mantle interactions: 3. Petrogenesis of intraplate magmas and structure of the upper mantle. Chem Geol 82:187–207 Rino S, Kon Y, Sato W, Maruyama S, Santosh M, Zhao D (2008) The Grenvillian and Pan-African orogens: world’s largest orogenies through geologic time, and their implications on the origin of superplume. Gondwana Res 14:51–72 Roobol MJ, Ramsay CR, Jackson NJ, Darbyshire DP (1983) Late Proterozoic lavas of the Central Arabian Shield—evolution of an ancient volcanic arc system. J Geol Soc 140(2):185–202 Salameh, SMS (1987) Petrology and geochemistry of the dyke swarm in Wadi Turban, southwest Jordan. M.Sc. thesis, University of Jordan, Amman Saleeb-Roufaiel GS (1975) On the time relation between the Igla Formation and the felsite bodies at Wadi Igla, Eastern Desert. Geol Soc Egypt (Abstract) Saleeb-Roufaiel GS, Samuel MD, Meneisy MY, Moussa HE (1989) K– Ar age determinations of Phanerozoic basaltic rocks in West Central Sinai. N Jb Geol Paläont Mh 11:683–691 Samuel MD (1977) Lithological and sedimentological studies on the red beds of Wadi Igla, Eastern Desert, Egypt. Bull Nation Res Cent Egypt 2:287–297 Samuel MD, Be’eri-Shlevin Y, Azer MK, Whitehouse MJ, Moussa HE (2011) Provenance of conglomerate clasts from the volcano-sedimentary sequence at Wadi Rutig in southern Sinai, Egypt as revealed by SIMS U–Pb dating of zircon. Gond Res 20:450−464 Samuel MD, Moussa HE, El Gharabawi RE, Sadek Ghabrial D (1999) Geochemistry of Mesozoic picritic and basaltic rocks of West Central Sinai. MERC Ain Shams Univ Earth Sci Ser 13:55–71 Samuel MD, Moussa HE, Ali Bik MW (2000) Geochemistry of Oligo-Miocene continental rift basalts, West and Central Sinai, Egypt. Egypt J Geol 44(2):115–129 Samuel MD, Moussa HE, Azer MK (2001) Geochemistry and petrogenesis of Iqna Shar, volcanic rocks, Central Sinai, Egypt. Egypt J Geol 45(2):921–940 Samuel MD, Moussa HE, Mengel K (2002) Mineral chemistry and geochemistry of Gebel Himeiyir basalts, Central Sinai, Egypt. Egypt J Geol 46(1):73–97 Samuel MD, Moussa HE, Azer MK (2007) A-type Volcanics in Central Eastern Sinai, Egypt. J Afr Earth Sci 47:203–226 Samuel MD, Moussa HE, Azer MK, Ghabrial DS (2019) Geochemical constraints of the Ediacaran Volcano-Sedimentary Succession of the Sa’al Metamorphic Complex at Wadi Zaghra, South Sinai, Egypt. Acta Geol Sin Engl Ed 93(1):50–73 Schürmann HME (1966) The Pre-Cambrian along the Gulf of Suez and the northern part of the Red Sea. EJ Brill Shaw JE, Baker JA, Menzies MA, Thirlwall MF, Ibrahim KM (2003) Petrogenesis of the largest intraplate volcanic field on the Arabian Plate (Jordan): a mixed lithosphere–asthenosphere source activated by lithospheric extension. J Petrol 44(9):1657–1679 Shimron AE (1980) Proterozoic island arc volcanism and sedimentation in Sinai. Precambr Res 12:437–458 Shimron AE (1987) Pan-African metamorphism and deformation in the Wadi, Kid Region, SE Sinai Peninsula: evidence from porphyroblasts in the Umm Zariq Formation. Isr J Earth Sci 36(4):173–193 Shimron A (1981) The Dahab mafic–ultramafic complex, southern Sinai Peninsula: a probable ophiolite of Late Proterozoic (Pan-African) age. Ofioliti 6(1):161–164 Shimron AE (1984) Evolution of the Kid Group, southeast Sinai Peninsula: Thrusts, mélange, and implications for accretionary tectonics during the late Proterozoic of the Arabian-Nubian Shield. Geology 12(4):242–247

532 Shimron AE, Furnes H, Roberts D, Bogoch R (1993) Petrogenesis of the Late Proterozoic Sa, al Group, southern Sinai Peninsula. GSI Curr Res 8:24–29 Smith IEM, Worthington TJ, Stewart RB, Price RC, Gamble JA (2003) Felsic volcanism in the Kermadic arc, SW Pacific: crustal recycling in an oceanic setting—intra-oceanic-subduction systems: tectonic and magmatic processes. Geol Soc Lond Spec Publ 219:99–118 Stacey JS, Agar RA (1985) U–Pb isotopic evidence for the accretion of a continental microplate in the Zalm region of the Saudi Arabian Shield. J Geol Soc Lond 142:1189–1203 Stacey JS, Stoeser DB, Greenwood WR, Fischer LB (1984) U–Pb zircon geochronology and geological evolution of the Halaban–AI Amar region of the Eastern Arabian Shield, Kingdom of Saudi Arabia. J Geol Soc Lond 141:1043–1055 Stein M (2003) Tracing the plume material in the Arabian-Nubian Shield. Precambrian Research 123(2–4):223–234 Stein M, Goldstein SL (1996) From plume head to continental lithosphere in the Arabian-Nubian Shield. Nature 382:773–778 Stern RJ (1979) Late Precambrian ensimatic volcanism in the Central Eastern Desert of Egypt. Ph.D. thesis. San Diego University, California, 210pp Stern RJ (1981) Petrogenesis and tectonic setting of Late Precambrian ensimatic volcanic rocks, Central Eastern Desert of Egypt. Precambr Res 16(3):195–230 Stern RJ (1994) Arc assembly and continental collision in the Neoproterozoic East African Orogen: implications for the consolidation of Gondwanaland. Ann Rev Earth Planet Sci 22:319–351 Stern RJ (2002) Crustal evolution in the East African Orogen: a neodymium isotopic perspective. J Afr Earth Sci 34:109–117 Stern RJ (2018) Neoproterozoic formation and evolution of Eastern Desert continental crust—the importance of the infrastructure-superstructure transition. J Afr Earth Sci 146:15–27 Stern RJ, Ali K (2020) Crustal evolution of the Egyptian Precambrian rocks. In: Geology of Egypt. Springer Nature, Switzerland, pp 131– 151 Stern RJ, Gottfried D (1986) Petrogenesis of a Late-Precambrian (575– 600 Ma) bimodal suite in northern Africa. Contrib Miner Petrol 92:492–501 Stern RJ, Hedge CE (1985) Geochronologic and isotopic constraints on late Precambrian crustal evolution in the Eastern Desert of Egypt. Amer J Sci 285:97–127 Stern RJ, Johnson PR (2010) Continental lithosphere of the Arabian Plate: a geologic, petrologic, and geophysical synthesis. Earth Sci Rev 101:29–67 Stern RJ, Gottfried D, Hedge CE (1984) Late Precambrian rifting and crustal evolution in the Northeastern Desert of Egypt. Geology 12 (3):168–172 Stern RJ, Ali KA, Liégeois JP, Johnson PR, Kozdroj W, Kattan FH (2010) Distribution and significance of pre-Neoproterozoic zircons in juvenile Neoproterozoic igneous rocks of the Arabian-Nubian Shield. Amer J Sci 310(9):791–811 Stern RJ, Johnson PR, Kröner A, Yibas B (2004) Neoproterozoic ophiolites of the Arabian-Nubian Shield. In: Kusky TM (ed) Precambrian ophiolites and related rocks. In: Developments in Precambrian geology, vol 13, pp 95–128 Stern RJ, Sellers G, Gottfried D (1988) Bimodal dyke swarms in the North Eastern Desert of Egypt: significance for the origin of late Precambrian A-type granites in Northern Afro-Arabia. In: The Pan-African belt of Northeast Africa and adjacent areas: tectonic evolution and economic aspects of a Late Proterozoic orogen (pp. 147–179) Stoeser DB, Stacey JS (1988) Evolution, U–Pb geochronology, and isotope geology of the Pan-African Nabitah orogenic belt of the

M. K. Azer et al. Saudi Arabian shield. In: El-Gaby S, Greiling RO (eds) The Pan-African Belt of Northeast Africa and adjacent areas: Braunschweig/Wiesbaden, Vieweg und Sohn, pp 227–288 Sturchio NC, Sultan M, Batiza R (1983) Geology and origin of Meatiq Dome, Egypt: A Precambrian metamorphic core complex?. Geology 11(2):72 Sultan M, Chamberlain KR, Bowring SA, Arvidson RE, Abu Zied H, El Kaliouby B (1990) Geochronologic and isotopic evidence for involvement of pre-Pan-African crust in the Nubian Shield, Egypt. Geology 18:761–764 Sultan, M, Tucker RD, El Alfy Z, Attia R, Ragab, AG (1994) U–Pb (zircon) ages for the gneissic terrane west of the Nile, southern Egypt. Geol Rundsch 83(3):514–522 Takla M, Hussein A (1995) Shield rocks and related mineralization in Egypt. In: Eleventh sympoisum on Precambrian and development, Cairo (Abstract) Taman Z (1996) Geology and mineralization of Wadi Atalla area, Eastern Desert of Egypt. M.Sc. thesis. Ain Shams University Teklay M, Kröner A, Mexger K (2001) Geochemistry, geochronology and isotope geology of Nakfa intrusive rocks, northern Eritrea: products of a tectonically thickened Neoproterozoic arc crust. J Afr Earth Sci 33:283–301 Thiéblemont D (2000) A geochemical database for the Proterozoic magmatism of the Arabian-Nubian Shield: final report Arabian-Nubian Shield project, in GIS Arabia, Bureau de Recherches Géologiques et Minières Valley JW, Kinny PD, Schulze DJ, Spicuzza MJ (1998) Zircon megacrysts from kimberlite: oxygen isotope variability among mantle melts. Contrib Miner Petrol 133:1–11 Wachendorf H, Jarrar GH, Zachmann D (1985) The role of pressure in control of potassium, sodium, and copper concentration in hypabyssal intrusives as demonstrated in late Precambrian dykes in SW Jordan. Precambr Res 30:221–248 Wallace CA (1986) Lithofacies and depositional environments of the Maraghan Formation and speculation on the origin of Gold in Ancient Mines, an Najadi Area, Kingdom of Saudi Arabia; Saudi Arabian Deputy Ministry for Mineral Resources: Jiddah, Saudi Arabia; USGS-OF-06-6, pp 1–19 Wang Q, Wyman DA, Xu JF, Jian P, Zhao ZH, Li CF, Xu W, Ma JL, He B (2007) Early Cretaceous adakitic granites in the Northern Dabie Complex, central China: implications for partial melting and delamination of thickened lower crust. Geoch Cosmoch Acta 71:2609–2636 Whalen JB, Currie KL, Chappell BW (1987) A-type granites: geochemical characteristics, discrimination and petrogenesis. Contrib Mineral Petrol 95:407–419 Wilde SA, Youssef K (2000) Significance of SHRIMP U–Pb dating of the imperial porphyry associated Dokhan volcanics, Gabal Dokhan, northern Eastern Desert, Egypt. J Afr Earth Sci 3:403–413 Wilde SA, Youssef K (2001) SHRIMP U–Pb dating of detrital zircons from the Hammamat group at Gebel Umm Tawat, north Eastern Desert, Egypt. Gondwana Res 4(2):202–206 Wilde SA, Youssef K (2002) A re-evaluation of the origin and setting of the Late Precambrian Hammamat group based on SHRIMP U–Pb dating of detrital zircons from Gebel Umm Tawat, North Eastern Desert, Egypt. J Geol Soc 159:595–604 Willis KM, Stern RJ, Clauer N (1988) Age and geochemistry of Late Precambrian sediments of the Hammamat Series from the Northeastern Desert of Egypt. Precambr Res 42(1–2):173–187 Xu J-F, Shinjo R, Defant MJ, Wang Q, Rapp RP (2002) Origin of Mesozoic adakitic intrusive rocks in the Ningzhen area of east China: partial melting of delaminated lower continental crust? Geology 30:1111–1114

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Volcanism During the Post-accretionary Stage …

Zhang HF, Parrish R, Zhang L, Xu W-C, Yuan H-L, Gao S, Crowley QG (2007) A-type granite and adakitic magmatism association in Songpan-Garze fold belt, eastern Tibetan Plateau: implication for lithospheric delamination. Lithos 97:323–335

533 Zimmer M, Kröner A, Jochum KP, Reischmann T, Todt W (1995) The Gabal Gerf complex: a Precambrian N-MORB ophiolite in the Nubian Shield, NE Africa. Chem Geol 123:29–51

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An Overview Study of Zircon Geochronology from Sinai Precambrian Basement: Implications for Crustal Evolution of Northern Arabian-Nubian Shield Mahrous M. Abu El-Enen and Kamal A. Ali

Abstract

Compiled U–Pb and Pb–Pb zircon ages from the Sinai Peninsula Precambrian basement record a crustal history covering of a time span of *450 Ma, which reflects a complex history of magmatism, sedimentation, and metamorphism extending from ca. 1030–578 Ma. Three magmatic main age populations are recognized at 1025– 975, 850–725, and 650–575 Ma and are interpreted to represent three discrete calc-alkaline volcanic arc magmatic episodes for the Sinai basement, succeeded by a phase of post-collisional within-plate alkaline-peralkaline magmatism at the end of the Neoproterozoic time. Detrital zircon ages show three age populations comparable to those of the magmatic zircons, suggesting that the host sedimentary rocks were deposited within basins adjacent to the volcanic arcs. Provenances of the sediments include Neoproterozoic arc assemblages, and limited contributions from old continental crust trapped within the Arabian-Nubian Shield or located at its border with the Saharan Metacraton. Metamorphic zircons within the Sinai metamorphic rocks give ages consistent with two discrete high-grade metamorphic events at 800– 750 Ma and 627–592 Ma in the Kid and Feiran-Solaf metamorphic complexes, respectively. The latter metamorphic event is divisible into two sub-events: a regional mid-crustal level upper amphibolite facies event, and a shallow-crustal level isothermal decompression event, caused by heat released from the youngest arc magmatism in the Sinai. Inherited zircons of Paleoproterozoic, Mesoproterozoic, and older ages are infrequent in the magmatic rocks of the Sinai basement. The common age

M. M. Abu El-Enen (&) Geology Department, Faculty of Sciences, Mansoura University, Mansoura, 35516, Egypt K. A. Ali Department of Geology, United Arab Emirates University, P.O. Box 15551 Al Ain, Abu Dhabi, UAE

of zircon xenocrysts is Neoproterozoic, implying a low contribution of reworked older crust. Keywords








In situ zircon dating Sinai Peninsula Ediacaran Magmatic Metamorphic

21.1





Cryogenian Inherited




Introduction

Absolute age determination of geological events is fundamentally important for constraining models of Earth evolution. Among the most important of the dating techniques are those based on zircon radiogenic element isotopic compositions. Zircon (ZrSiO4) is a common minor or accessory mineral in a wide variety of igneous and metamorphic rocks, which represent the main provenances of siliciclastic sediments. Zircon has played a key role in geochronology, tracing crustal evolution, and constraining petrogenetic models (Harley and Kelly 2007). This is due to its chemical stability and mechanical resilience, and the significant contents of radioactive U and Th isotopes, along with typically low initial Pb contents. Thus, zircon crystals are like time capsules, preserving details of the igneous and metamorphic history of their host rocks. Progressive improvements in situ U–Pb geochronology include the use of small-diameter laser, ion and electron beams, high-precision mass spectrometry, and a variety of microscopic imaging methods. These have enabled the isotopic analysis of tiny volumes of complexly zoned zircon crystals, resulting in more precise dating (Harley and Kelly 2007). Reliable and convenient zircon dating methods, including secondary ion mass spectrometry (SIMS), laser-ablation-induced coupled plasma–mass spectrometry (LA-ICP-MS), and low-blank thermal-ionization mass spectrometry (TIMS), in addition to evaporation techniques, are the most accurate and precise methods for isotopic

© The Author(s), under exclusive license to Springer Nature Switzerland AG 2021 Z. Hamimi et al. (eds.), The Geology of the Arabian-Nubian Shield, Regional Geology Reviews, https://doi.org/10.1007/978-3-030-72995-0_21

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analyses of zircons (Schaltegger et al. 2015). The Pb–Pb evaporation method of single zircons only provides 207 Pb/206Pb ages and cannot reveal whether a zircon has lost or gained Pb (Kober 1986); however, this technique is still suitable for identifying zircon xenocrysts. SIMS zircon ages typically have uncertainties that are significantly greater than those yielded by the best TIMS laboratories (Bowring and Schmitz 2003), though the former technique permits faster dating of larger numbers of zircons. The LA-ICP-MS technique is similar to the SIMS method, except that laser-ablation spots are larger (ca. 30–65 µm against 15– 30 µm for SIMS) and generally deeper, and analyses are faster. Therefore, the LA-ICP-MS technique can rapidly provide a large number of analyses, and for this reason is increasingly used to date detrital zircons. A beneficial feature of modern zircon geochronology is the improved capacity to recognize and individually date multiple events, based on the evidence of growth zoning revealed in the zircons. Moreover, oxygen and hafnium isotopes in zircon can be used to constrain models of crustal growth. The main aims of this study are to review the existing state of knowledge of U–Pb in situ zircon geochronology for the basement rocks of the Sinai Peninsula, and to use this database to assign dates for the major geological events, including magmatic crystallization, maximum depositional age of sedimentary successions, and timing of metamorphic peaks in the Sinai basement. Further aims are to discuss the significance of these events in the evolution of the Precambrian crustal rocks of the Sinai and clarify details of the petrogenesis of the magmatic rocks in the region.

21.2

Regional Geology of the Sinai Precambrian Basement

The exposed crystalline basement of the Sinai Peninsula (Fig. 21.1) comprises voluminous unmetamorphosed post-collisional calc-alkaline granitoid batholiths and less abundant within-plate alkaline granite plutons, associated with slightly to non-metamorphosed volcano-sedimentary successions. The plutons intrude and isolate metamorphic complexes of strongly deformed and metamorphosed sediments, arc volcanics, and calc-alkaline granitoids (Shimron and Zwart 1970; Reymer et al. 1984; Abu El-Enen 1995, 2008, 2011; Cosca et al. 1999; Abu El-Enen et al. 2003, 2004; Brooijmans et al. 2003). The metamorphic complexes are: the Taba (TMC) in the northeast, the Kid (KMC) in the southeast, the Sa’al-Zaghra (SZMC) in the center, and the Feiran-Solaf (FSMC) in the west of the exposed basement, in addition to small fragments of the complexes dispersed within the granitoids (Fig. 21.1). These complexes experienced low-pressure metamorphism at greenschist to

amphibolite facies conditions, accompanying three ductile deformation phases D1 to D3. A single Late Neoproterozoic cycle of regional metamorphism is envisaged, with the prograde stage coinciding with D1, the peak conditions during D2, and a retrograde stage occurring during D3 (Eyal 1980; Abu El-Enen 1995, 2008, 2011; Cosca et al. 1999; Broojmans et al. 2003; Abu El-Enen et al. 2003, 2004, Abu El-Enen and Okrusch 2007; Eliwa et al. 2008; Moghazi et al. 2012; Abu El-Enen and Whitehouse 2013; Elisha et al. 2019). However, Hassan et al. (2014) concluded that metamorphism in Sa’al–Zaghra complex records an earlier stage of metamorphism and deformation during break-up of Rodinia. They further concluded that this complex escaped deep crustal metamorphism during the Pan-African event related to Gondwana collision. The Taba Metamorphic Complex (TMC) is composed of metapelitic schists, migmatitic metapelites, various types of orthogneisses, metagabbros, and metamorphosed dykes (Eyal 1980; Abu El-Enen 1995; Abu El-Enen et al. 1999, 2004, 2009). The protoliths of the metapelitic schists were deposited in an active continental margin or island arc regime and were later metamorphosed at amphibolite facies conditions (Abu El-Enen et al. 2004). The migmatitic metapelites are formed by metamorphic differentiation under upper amphibolite conditions (Amit and Eyal 1976; Abu El-Enen et al. 2009). The six types of orthogneisses are grouped into (1) an older suite of highly deformed rocks of calc-alkaline quartz-diorite and tonalite composition; and (2) younger suite of less deformed rocks of alkaline to calc-alkaline quartz monzonite to alkali granite composition (Abu El-Enen et al. 1999). The protoliths of these gneisses were emplaced in an arc tectonic setting along an active continental margin during the pre- to syn-collision stage of the Pan-African Orogeny (Abu El-Enen et al. 1999). The Feiran-Solaf Metamorphic Complex (FSMC) comprises the Solaf zone in the SE and Feiran zone in the NW, which are separated by a narrow NE-trending elongated diorite body. The two zones consist of metasedimentary and meta-igneous rocks of various parentage, represented now by gneisses, calc-silicates, metapelites, migmatitic metapelites, and migmatitic gneisses (El-Shafei and Kusky 2003; Abu El-Enen 2011; Abu El-Enen and Whitehouse 2013; Eyal et al. 2014a). The complex underwent high-grade metamorphism to upper amphibolite facies, resulting in local migmatization of both metapelites and Feiran-zone orthogneisses (Abu-Alam and Stüwe 2009; Abu-Alam et al. 2010; Abu El-Enen et al. 2009; Abu El-Enen 2011), and reportedly displays a sinistral NW–SE striking Najd transcurrent shear system in an oblique transpressive regime against the surrounding granitoids (Abu-Alam and Stüwe 2009). The metasedimentary rocks occupy the Solaf zone and the

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An Overview Study of Zircon Geochronology …

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Fig. 21.1 Regional geological map of the Neoproterozoic ANS exposures in the Sinai Peninsula (modified after Eyal 1980 and Be’eri-Shlevin et al. 2009b), showing the extent of rock groups and complexes, with ages determined by U–Pb and Pb–Pb analyses of zircon plotted (Locations are given in Appendix 1)

538

extreme northwestern part of the Feiran zone. They are represented by metapsammitic gneisses, calc-silicates, and migmatitic metapelites in the Solaf zone, and metapelites in the extreme northwestern part of the Feiran zone. However, orthogneisses and migmatitic orthogneisses also occupy the Feiran Zone and the northern part of the Solaf zone (Abu El-Enen and Whitehouse 2013). The protoliths of the meta-sediments and orthogneisses were deposited/emplaced in an active continental marginal basin within a continental-arc system, and later experienced upper amphibolite facies metamorphism (Abu-Alam and Stüwe 2009; Abu El-Enen et al. 2009; Abu-Alam et al. 2010; Abu El-Enen 2011; Abu El-Enen and Whitehouse 2013; Eyal et al. 2014a). The Sa’al-Zaghra metamorphic complex (SZMC) encompasses the Sa’al Group (Shimron et al. 1993), which consists of greenschist to amphibolite facies metasedimentary and metavolcanic sequences, gneisses, and locally migmatites (Be’eri-Shlevin et al. 2012; Hassan et al. 2014; Eyal et al. 2014a; Andresen et al. 2014; Fowler et al. 2015). The Sa’al Group incorporates three formations: the Ra’ayan (the oldest), the Agramiya, and Zaghra (the youngest), arranged from northwest to southeast. The Ra’ayan Formation is dominated by greenschist to lower amphibolite facies phyllites and schists at its base, overlain by greywackes with interbedded conglomerates (Shimron et al. 1993; Eyal et al. 2014a; Fowler et al. 2015). The Agramyia Formation is mainly a volcano-sedimentary succession, represented by metabasalt and meta-andesite flows and metarhyolite pyroclastics and ignimbrite (Be’eri-Shlevin et al. 2012; Eyal et al. 2014a, b; Hassan et al. 2014). The Zaghra Formation comprises tuffs, lapilli tuffs, and rhyolitic ignmbrite (northern part) and a succession of interbedded feldspathic sandstones, conglomerate, diamictite, and schists (the southern part at Wadi Zaghra) (Shimron et al. 1993; Eyal et al. 2014a; Andresen et al. 2014; Hassan et al. 2014). The SZMC underwent greenschist facies metamorphism during tectonic extension related to Rodinia break-up, followed by Pan-African amphibolite facies metamorphism (Hassan et al. 2014). The Kid Metamorphic Complex (KMC) consists of the Kid Group (Furnes et al. 1985; Shimron 1984, 1987), a deformed volcano-sedimentary succession with island arc geochemical affinity, including metapelites, meta-turbidites, followed by metavolcanics of andesitic to metarhyolitic composition, metaconglomerates, serpentinites, liptite, gneisses, and migmatites (Shimron 1984; Furnes et al. 1985; Blasband et al. 1997; Abu El-Enen et al. 2003; Abu El-Enen 2008; Fowler et al. 2010a, b). The KMC is divided into several tectonostratigraphic units (the Heib, Umm-Zariq, and Malhaq formations, and the Tarr and the Dahab-Shahira

M. M. Abu El-Enen and K. A. Ali

Complexes). There is considerable debate on their nature, extent, and contacts between these units (e.g., Shimron 1980; Reymer 1983; El-Gaby et al. 1991; Abu El-Enen and Makroum 2003). These rocks suffered low-grade metamorphism, except in some areas in the central and northern KMC, which experienced lower amphibolite facies (Furnes et al. 1985; Blasband et al. 1997; Abu El-Enen 2008; Abu El-Enen et al. 2003). The geochemical signature of KMC volcanic and metasedimentary rocks implies that they were generated in an island arc tectonic regime (Abu El-Enen et al. 2003; Abu El-Enen 2008; Eyal et al. 2014a). Volcanic rocks of the KMC have chemical similarities to the temporally equivalent Dokhan Volcanics (Moghazi et al. 2012). Rutig and Feirani Group volcano-sedimentary associations and scattered small exposures (e.g., at Mahash, Iqna Shar’a and Khashabi) are unmetamorphosed, apart from slight low-grade metamorphism at their base. They are time equivalents of the Dokhan Volcanics and Hammamat clastic sediments in the Eastern Desert of Egypt (Samuel et al. 2011; Moreno et al. 2012). The Rutig Group, exposed in the central part of the southern Sinai (Fig. 21.1), comprises intermediate to acid calc-alkaline lava flows and pyroclastics, alternating with conglomerates and sandstones (Eyal and Hezkiyahu 1980; Bentor 1985; Azer 2007; Katzir et al. 2007; Samuel et al. 2011; Moreno et al. 2012). These volcanics and sediments are intruded by the Katherine alkaline granite and are overlain by the later Katherine ring complex, which forms a huge ring dike of peralkaline granite (Eyal and Hezkiyahu 1980; Katzir et al. 2007). The Ferani Group, exposed in the eastern part of the Sinai (Fig. 21.1), is represented by intermediate to acid lava flows and pyroclastic rocks, alternating with immature sediments that are divided into lower and upper parts separated by coarse conglomerates (Moussa 2003). They were folded and metamorphosed at their base to lower greenschist facies (Bentor and Eyal 1987; Moussa 2003) and intruded by Dahab alkali granites (Be’eri-Shlevin et al. 2009b). The Rutig and Feirani volcano-sedimentary associations are assigned to be deposited in terrestrial basins of magmatically active rift, with rivers winding through desolate lava plains (Stern and Ali 2020). Voluminous granitoids are widespread in the Sinai basement and include calc-alkaline high-K suites and alkaline granites, with a time overlap between the two magmatic groups (Be’eri-Shlevin et al. 2009a). The calc-alkaline group is more dominant, occurs as batholiths, and is commonly identified as Older (Gray) Granitoid or syn- to late-tectonic granitoids. They are diorite, quartz-diorite, granodiorite, and monzogranite, showing little or no ductile deformation. They contain microgranular mafic enclaves of amphibolite or melanocratic diorite composition. The calc-alkaline granitoids were emplaced within a volcanic arc regime. Less abundant alkaline granitoid forms small circular to elliptical

21

An Overview Study of Zircon Geochronology …

plutons (