Geodynamic Evolution of the Indian Shield: Geophysical Aspects (Society of Earth Scientists Series) 3030405966, 9783030405960

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Society of Earth Scientists Series

Om Prakash Pandey

Geodynamic Evolution of the Indian Shield: Geophysical Aspects

Society of Earth Scientists Series Series Editor Satish C. Tripathi, Lucknow, India

The Society of Earth Scientists Series aims to publish selected conference proceedings, monographs, edited topical books/text books by leading scientists and experts in the field of geophysics, geology, atmospheric and environmental science, meteorology and oceanography as Special Publications of The Society of Earth Scientists. The objective is to highlight recent multidisciplinary scientific research and to strengthen the scientific literature related to Earth Sciences. Quality scientific contributions from all across the Globe are invited for publication under this series. Series Editor: Dr. Satish C. Tripathi

More information about this series at http://www.springer.com/series/8785

Om Prakash Pandey

Geodynamic Evolution of the Indian Shield: Geophysical Aspects

123

Om Prakash Pandey CSIR-National Geophysical Research Institute Hyderabad 500007, India

ISSN 2194-9204 ISSN 2194-9212 (electronic) Society of Earth Scientists Series ISBN 978-3-030-40596-0 ISBN 978-3-030-40597-7 (eBook) https://doi.org/10.1007/978-3-030-40597-7 © Springer Nature Switzerland AG 2020 This work is subject to copyright. All rights are reserved 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. This Springer imprint is published by the registered company Springer Nature Switzerland AG The registered company address is: Gewerbestrasse 11, 6330 Cham, Switzerland

Series Editor Foreword

Indian subcontinent has been the most traveled landmass of the Earth and suffered geodynamically varied conditions affecting crust–mantle evolution. The geological evidences need to be corroborated by geophysical data interpretation to strengthen our understanding of geodynamic evolution of Indian shield. The Indian shield consists of the oldest Archean nucleus, and various cratons were amalgamated together recording episodes of inter-cratonic and intra-cratonic tectonism, magmatism, and sedimentation. The book presents the geophysical data available on Indian shield and their interpretation. Such compilations open the gaps in our knowledge and new domains for future research. My sincere and personal thanks to Prof. J. R. Kayal for critically reviewing the manuscript and suggesting necessary changes to make the book more comprehensive. The mega-event of 36th International Geological Congress 2020 in India opened new chapter on the geology of India. On such an occasion, Society of Earth Scientists Series by Springer decided to bring out 36th IGC Commemorative Volumes on various recent geological and geophysical studies of India. As such, veteran geoscientists were requested to prepare comprehensive account as monographs or edited volumes. I am personally thankful to all the editors and authors for timely submission of high-quality manuscripts for inviting interest of global community of geoscientists. Lucknow, India

Satish C. Tripathi

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Preface

The nature of the evolution of the crust and the mantle lithosphere of the ancient terrains, and their subsequent deformation due to regional plate tectonics and continued intraplate geodynamic activity, has been a subject of immense interest to the geoscientific community. This is especially true for a unique Archean shield like India, which differs considerably from other stable regions of the Earth. The Indian shield remained quite active during its entire course of geologic history, having suffered several episodes of rifting, collision, and continental breakups. It also sustained massive intraplate volcanism, a major K-T boundary impact and interaction with four mantle plumes (Reunion, Marion, Crozet, and Kerguelen) in quick succession, during its super-mobile phase between 130 and 53 Ma. The scars created by these events are clearly visible in the geophysical and geological signatures over all the geotectonic segments of India. The Indian shield contains five major Eoarchean to Paleoproterozoic cratons, which are welded together by rift valleys, Gondwana basins, sutures, and mega-lineaments. Establishing the origin of Indian shield and its time-bound evolutionary nature, especially in relation to other coeval global cratonic shields, remains a challenge. There has been an explosion of geophysical data in last three decades, although these data sets remain insufficient in many geodynamically important areas, like Bastar and Bundelkhand cratons. Thorough interpretation of these data has helped immensely to decipher the preserved seismic, petrophysical, and thermo-geodynamic characters of the shield. At the same time, these studies have posed many new questions like whether (i) Vindhyan sediments occur south of the NSL rift zone, (ii) the repeatedly degenerated Indian crust is denudated so much that the mafic crust is exposed in majority of geological terrains, (iii) Indian cratonic keels are largely destroyed, unlike other global shields where they are still preserved, (iv) cratons like Singhbhum or Aravalli can be considered a craton in strict sense, or they should be termed mobile belts, or (v) mantle metasomatism has created enough crustal seismic inhomogeneity to cause intraplate earthquakes. These aspects need serious scrutiny.

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This book is an attempt to highlight the compositional and evolutionary nature of the crust and mantle lithosphere beneath different geotectonic terrains of India and adjacent western offshore region, based on in-depth and up-to-date interpretation of integrated geophysical, geological, and petrophysical studies, carried out in the last few decades. The book is arranged in ten chapters. Chapter 1 deals with the fundamental aspects of the Earth’s constitution and its dynamic nature. It also includes a brief summary of major constituents of Indian shield, their broad geophysical characters, and its critical position in Paleo-Supercontinent Assembly. This is followed by an in-depth analysis of geophysical data over the Dharwar craton and Cuddapah basin in Chap. 2, Singhbhum and Bastar cratons in Chap. 3, and Aravalli and Bundelkhand cratons in Chap. 4. The diamondiferous Vindhyan basin, which is considered one of the largest Proterozoic basins in the world, is dealt in Chap. 5. Western continental margins and associated offshore region (discussed in Chap. 6) played a key role in degenerating Indian lithosphere during Cretaceous. This region, bestowed with unusual geophysical characters, suffered several geodynamic upheavals, including Deccan volcanism and K-T mass extinction that killed dinosaurs. Since it is not possible to interpret the geological and geophysical data without the knowledge of seismic, elastic, and petrophysical properties, such aspects are covered in Chap. 7, where special emphasis is given to multi-parametric borehole sample studies. In the last fifty years, a number of disastrous intraplate earthquakes also took place in India, killing thousands of people. Chapter 8 focusses on cause and seismotectonics of some of these earthquakes. Earth’s heat flow is another important geophysical field that provides valuable information about the thermal state of the Earth’s crust and upper mantle. Its fundamental nature, and how to use such data in the realm of the lithospheric structure, is discussed in Chap. 9. Chapter 10 describes in detail our current understanding of the seismic nature and prevailing P-T condition in the Indian crust. Many geoscientists do not have a fair idea about this. This book is primarily intended to benefit a good cross section of geoscientific community, as it covers almost all the aspects of geophysical data and its interrelationship with crust–lithosphere structure and prevailing geodynamic scenario. I am sure the readers will be benefited and feel encouraged. I would like to express my sincere gratitude to Prof. V. P. Dimri, Prof. R. S. Sharma, and Dr. Satish C. Tripathi, who encouraged me all the time to write a book on ‘Geodynamics of India.’ This book contains a large number of studies carried out at CSIR-National Geophysical Research Institute, Hyderabad. I am thankful to all the former directors of NGRI, who encouraged and supported me to pursue such studies. During my stay at this institute, I got tremendous support from the present director Dr. V. M. Tiwari, my co-authors, and colleagues, like Dr. Nimisha Vedanti, Dr. Ravi Srivastava, Dr. T. R. K. Chetty, Dr. S. Masood Ahmad, Dr. K. J. P. Lakshmi, Dr. M. V. M. S. Rao, Dr. D. Srinivasa Sarma, Dr. G. Parthasarathy, Dr. S. S. H. Jafri, Dr. U. Raval, Dr. A. Keshav Krishna, Dr. M. Satyanarayanan, Dr. K. Chandrakala, Sri. G. Koti Reddy, Dr. A. Vasanthi, and Dr. Prantik Mandal. I also take this opportunity to

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sincerely acknowledge the support provided by Dr. Mukund Sharma (BSIP), Dr. B. Sreedhar (CSIR-IICT), Dr. H. K. Sachan (WIGH), Shri. V. Rajagopalan (AMD), Prof. J. P. Srivastava (DU), Dr. Sesa Sai (GSI), Prof. S. P. Singh (BU), and Prof. N. V. Chalapathi Rao (BHU) at various stages of the scientific studies. Thanks are also due to Dr. H. C. Tewari, Dr. K. Sain, Dr. Prakash Kumar and Dr. L. Behera for their supportive help. Dr. Priyanka Tripathi, also helped me a lot during the writing of this book. Many figures of the book were made by Sri. M. Vittal, Sri. Sujeet Dwivedi, Dr. K. M. Bhatt, Dr. B. Mandal, Dr. Sandeep Gupta, Dr. P. Karuppannan, Dr. K. K. Abdul Azeez, and Dr. S. S. Ganguli. Further, I am greatly indebted to Prof. J. R. Kayal, who in spite of being very busy took so much pain to review thoroughly all the chapters of this book and suggested needful changes. Publication of this book was expertly handled by Sudhany Karthick of Springer. I am also extremely thankful to Geological Society of India (Bengaluru), Elsevier, Springer, JESS and JIGU, for permitting to reproduce some of the figures used in this book. CSIR (New Delhi) is thanked for granting me Emeritus Scientist Scheme. All the support extended by the Indian Geophysical Union is also thankfully acknowledged. My deepest gratitude remain to Late Professor Dr. J. G. Negi, who was my mentor at NGRI. Professor Negi introduced me to the wide spectrum of geophysical sciences and guided me to grow as an individual and to carry out curiosity-driven basic research. I was also fortunate to have Late Professor F. F. Evison at the Victoria University of Wellington, New Zealand, as my Ph.D. supervisor, who taught me how to become an independent researcher. At this occasion, it is my wish to recall sweet remembrance to my wife late Kamala, my brothers late N. N. Pandey and S. N. Pandey, and my parents late Indrasan Pandey and Radhika Pandey, who shaped my career to what I am today. I also express my sincere appreciation toward my son Manas, daughter Manisha, her husband Anand, and daughter-in-law Sudha, without whose support, love, and blessings, this kind of work would not have been possible. Hyderabad, India

Om Prakash Pandey

Contents

1

Geodynamic and Geologic Evolution of Indian Continent: A Brief History . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.2 Plate Tectonic Processes . . . . . . . . . . . . . . . . . . . . . . . . 1.2.1 Convergent Plate Margin . . . . . . . . . . . . . . . . . 1.3 Indian Subcontinent and Paleo-Supercontinent Assembly 1.3.1 Lateral Geologic Correlation . . . . . . . . . . . . . . . 1.3.2 Paleo-Super Dharwar Craton and Madagascar Breakup . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.4 Geodynamic Evolution of Indian Subcontinent Since Early Cretaceous . . . . . . . . . . . . . . . . . . . . . . . . . 1.4.1 Super Mobility . . . . . . . . . . . . . . . . . . . . . . . . . 1.5 Major Geological Segments of the Indian Peninsular Shield . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.5.1 Geological Time Scale . . . . . . . . . . . . . . . . . . . 1.5.2 Indian Cratons . . . . . . . . . . . . . . . . . . . . . . . . . 1.5.3 Proterozoic Basins . . . . . . . . . . . . . . . . . . . . . . 1.5.4 Gondwana Basins . . . . . . . . . . . . . . . . . . . . . . . 1.5.5 Tertiary Basins . . . . . . . . . . . . . . . . . . . . . . . . . 1.5.6 Deccan Traps . . . . . . . . . . . . . . . . . . . . . . . . . . 1.6 Regional Geophysical Studies . . . . . . . . . . . . . . . . . . . . 1.6.1 Gravity Studies . . . . . . . . . . . . . . . . . . . . . . . . 1.6.2 Deep Crustal Seismic, Tomographic and Receiver Function Studies . . . . . . . . . . . . . 1.6.3 Heat Flow and Lithospheric Studies . . . . . . . . . 1.6.4 Magnetotelluric Studies . . . . . . . . . . . . . . . . . . 1.6.5 Seismicity . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

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Dharwar Craton . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.2 Geological Settings . . . . . . . . . . . . . . . . . . . . . . . . 2.3 Geophysical Studies . . . . . . . . . . . . . . . . . . . . . . . 2.4 Western Dharwar Craton . . . . . . . . . . . . . . . . . . . . 2.4.1 Crustal Seismic Structure . . . . . . . . . . . . . 2.4.2 Gravity Field . . . . . . . . . . . . . . . . . . . . . . 2.4.3 MT Studies . . . . . . . . . . . . . . . . . . . . . . . 2.4.4 Heat Flow and Lithosphere . . . . . . . . . . . . 2.5 Eastern Dharwar Craton . . . . . . . . . . . . . . . . . . . . 2.5.1 Crustal Seismic Structure . . . . . . . . . . . . . 2.5.2 MT Studies . . . . . . . . . . . . . . . . . . . . . . . 2.5.3 Regional Gravity Field . . . . . . . . . . . . . . . 2.5.4 Magnetic Studies . . . . . . . . . . . . . . . . . . . 2.5.5 Heat Flow and Lithosphere Structure . . . . . 2.5.6 Evolution of Cuddpah Basin and Adjacent Continental Terrain . . . . . . . . . . . . . . . . . 2.6 Southern Granulite Terrain . . . . . . . . . . . . . . . . . . 2.6.1 Crustal Seismic Structure . . . . . . . . . . . . . 2.6.2 MT Studies . . . . . . . . . . . . . . . . . . . . . . . 2.6.3 Gravity Field . . . . . . . . . . . . . . . . . . . . . . 2.6.4 Heat Flow and Lithosphere Structure . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

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Singhbhum and Bastar Cratons . . . . 3.1 Singhbhum Craton . . . . . . . . . . 3.1.1 Introduction . . . . . . . . . 3.1.2 Geological Setting . . . . 3.1.3 Geophysical Signatures . 3.2 Bastar Craton . . . . . . . . . . . . . . 3.2.1 Geological Settings . . . 3.2.2 Geophysical Studies . . . References . . . . . . . . . . . . . . . . . . . . .

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Aravalli and Bundelkhand Cratons . . . . . 4.1 Aravalli Craton . . . . . . . . . . . . . . . . 4.1.1 Introduction . . . . . . . . . . . . 4.1.2 Regional Geology . . . . . . . 4.1.3 Geophysical Studies . . . . . . 4.2 Bundelkhand Craton . . . . . . . . . . . . 4.2.1 Geophysical Characteristics . 4.3 A Composite Geodynamic Model . . References . . . . . . . . . . . . . . . . . . . . . . . .

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5

Vindhyan Basin: Anomalous Crust-Mantle Structure . 5.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.2 Geotectonic and Geologic Features . . . . . . . . . . . 5.2.1 Geochronology and Stratigraphy . . . . . . . 5.2.2 Jabera-Damoh Region . . . . . . . . . . . . . . 5.3 Geophysical Characteristics . . . . . . . . . . . . . . . . . 5.3.1 Crustal Seismic Structure from DSS . . . . 5.3.2 Broadband Seismic Studies . . . . . . . . . . . 5.3.3 Gravity Investigations . . . . . . . . . . . . . . . 5.3.4 Magnetotelluric (MT) Studies . . . . . . . . . 5.3.5 Heat Flow Studies . . . . . . . . . . . . . . . . . 5.4 Geodynamic Evolution of Vindhyan Basin . . . . . . 5.4.1 1.1 Ga Super Plume Interaction . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

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Western Continental Margin and Adjacent Oceanic Regions . . . 6.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.2 Super Mobility and Geodynamic Events . . . . . . . . . . . . . . . . 6.2.1 Madagascar Breakup . . . . . . . . . . . . . . . . . . . . . . . 6.2.2 K-T Impact and Deccan Volcanism . . . . . . . . . . . . . 6.3 Thermal Structure . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.3.1 Cambay Basin . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.3.2 Geothermal Springs . . . . . . . . . . . . . . . . . . . . . . . . 6.3.3 West Coast Thermal Anomaly Zone . . . . . . . . . . . . 6.4 Gravity Field, MT Studies and Seismicity . . . . . . . . . . . . . . 6.5 Deep Crustal Seismic Studies . . . . . . . . . . . . . . . . . . . . . . . 6.5.1 DSS Studies . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.5.2 Receiver Function Studies . . . . . . . . . . . . . . . . . . . 6.6 Impact Induced Rifting and Deccan Volcanism . . . . . . . . . . . 6.7 Seychelles Breakup and Initiation of Laxmi and Calsberg Ridges . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.7.1 Laxmi Ridge Evolution . . . . . . . . . . . . . . . . . . . . . 6.7.2 Crust-Mantle Structure in Arabian Sea and Western Continental Region . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

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7

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Seismic, Elastic and Petrophysical Properties of Crustal Rocks: Deccan Volcanic Province . . . . . . . . . 7.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.2 Deep Scientific Drilling . . . . . . . . . . . . . . . . . . . . . 7.3 Measurement of Seismic, Elastic and Petrophysical Properties . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.3.1 Experimental Techniques . . . . . . . . . . . . . 7.3.2 Integrated Geological Studies . . . . . . . . . .

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7.4

Crystalline Basement . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.4.1 Geological Nature of Crystalline Basement . . . . . . . 7.4.2 Elastic and Petrophysical Properties . . . . . . . . . . . . . 7.4.3 Mineralogical Effects on Density, Velocity and Elastic Moduli . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.4.4 Mantle Metasomatism and Anomalous Velocity Drop in Iron and Biotite-Rich Rocks . . . . . . . . . . . . . . . . 7.5 Deccan Volcanics . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.5.1 Geological Nature of Basaltic Column . . . . . . . . . . . 7.5.2 Seismic and Petrophysical Properties . . . . . . . . . . . . 7.5.3 Compressional and Shear Wave Attenuation . . . . . . 7.5.4 Characteristic Density and P- and S-Wave Velocity for Deccan Basalts . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

8

Seismic Instability and Major Intraplate Earthquakes . . . . . . 8.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8.2 1993 Killari Earthquake . . . . . . . . . . . . . . . . . . . . . . . . . 8.2.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8.2.2 Geophysical Studies . . . . . . . . . . . . . . . . . . . . . . 8.3 2001 Bhuj Earthquake . . . . . . . . . . . . . . . . . . . . . . . . . . . 8.3.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8.3.2 Seismic Activity . . . . . . . . . . . . . . . . . . . . . . . . . 8.3.3 Seismic Reflection Studies . . . . . . . . . . . . . . . . . 8.3.4 Gravity Field . . . . . . . . . . . . . . . . . . . . . . . . . . . 8.3.5 MT Studies . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8.4 1967 Koyna Earthquake . . . . . . . . . . . . . . . . . . . . . . . . . 8.4.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8.4.2 Scientific Deep Drilling and Nature of Crystalline Basement . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8.4.3 Crustal Seismic Studies . . . . . . . . . . . . . . . . . . . 8.4.4 Gravity Field . . . . . . . . . . . . . . . . . . . . . . . . . . . 8.4.5 Thermal Regime and Lithosphere Structure . . . . . 8.4.6 MT Studies . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8.4.7 Aeromagnetic Study . . . . . . . . . . . . . . . . . . . . . . 8.5 Major Earthquakes Associated with NSL . . . . . . . . . . . . . 8.6 Other Prominent Earthquakes . . . . . . . . . . . . . . . . . . . . . 8.7 Mantle Metasomatism and Crustal Low Velocity Zones . . 8.7.1 Earthquake Nucleation . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

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Heat Flow and Lithospheric Thermal Structure . . . . . . . . . . . 9.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 9.2 Terrestrial Heat Flow . . . . . . . . . . . . . . . . . . . . . . . . . . . 9.2.1 Temperature Measurements . . . . . . . . . . . . . . . . . 9.2.2 Thermal Conductivity . . . . . . . . . . . . . . . . . . . . . 9.3 Heat Flow Estimation . . . . . . . . . . . . . . . . . . . . . . . . . . . 9.4 Crustal Radioactivity . . . . . . . . . . . . . . . . . . . . . . . . . . . . 9.4.1 Fractal Behaviour in Crustal Radioactivity . . . . . . 9.5 Radioactive Heat Generation and Temperature-Depth Estimation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 9.5.1 Heat Flow-Heat Generation Relationship . . . . . . . 9.5.2 Temperature-Depth Estimation . . . . . . . . . . . . . . 9.6 Heat Flow Studies in India . . . . . . . . . . . . . . . . . . . . . . . 9.6.1 Regional Heat Flow Distribution . . . . . . . . . . . . . 9.7 Lithosphere Thickness Variation . . . . . . . . . . . . . . . . . . . 9.7.1 Previous Studies . . . . . . . . . . . . . . . . . . . . . . . . . 9.7.2 Present Estimates of Lithospheric Thickness . . . . 9.8 Lithospheric Mantle Deformation Beneath Indian Cratons . 9.9 Distribution of Geothermal Springs . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

10 Indian Crust . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 10.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 10.2 What Is Crust? . . . . . . . . . . . . . . . . . . . . . . . . . . . . 10.2.1 Oceanic Crust . . . . . . . . . . . . . . . . . . . . . . 10.2.2 Continental Crust . . . . . . . . . . . . . . . . . . . . 10.3 Crustal Thickness . . . . . . . . . . . . . . . . . . . . . . . . . . 10.4 Continental Crust: A General Perception . . . . . . . . . 10.5 P-T Regime in the Continental Crust . . . . . . . . . . . . 10.6 Indian Crust: A Paradigm Shift . . . . . . . . . . . . . . . . 10.6.1 Deep Borehole Studies . . . . . . . . . . . . . . . . 10.6.2 Widespread Effect of Mantle Metasomatism 10.7 Sub-division of the Indian Crust . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

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About the Author

Dr. Om Prakash Pandey, FNASc, born on May 12, 1950, is an internationally known scientist, who obtained B.Sc. from University of Lucknow, B.Sc. (Hons) and M.Sc. in applied geophysics from Indian School of Mines (IIT-ISM), Dhanbad, and Ph.D. from Victoria University of Wellington, New Zealand. He served CSIR-National Geophysical Research Institute, Hyderabad, for 45 years, where he pursued his research as a CSIR-Emeritus Scientist after superannuating as Chief Scientist. Currently, he is working as Chief Editor, The Journal of Indian Geophysical Union. He has made outstanding contributions in multidisciplinary geoscientific fields like terrestrial heat flow, lithosphere structure, geophysical, geological, geochemical, and geodynamic evolution of continents, catastrophic events, seismogenesis, global volcanism, and Earth resources development. He was the first researcher to advocate that the Indian shield lithosphere is unusually hot and thin, containing non-rigid and deformed mafic crust. He also produced first geophysical evidences of a possible asteroidal impact near offshore Mumbai, triggering Deccan volcanic eruption that led to the K-T boundary mass extinction, killing the dinosaurs. He firmly established Madagascar as a continental fragment of the paleo-super Dharwar craton. He has over 90 widely cited research papers to his credit and has actively participated in IUGG, IGC, ASC, WGC, and Goldschmidt conferences. Most of postulations made by him are widely cited and subsequently endorsed by national and global researchers.

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About the Author

He won many prestigious fellowships like Senior DAAD Fellowship (Germany), Commonwealth Fellowship (New Zealand), and INSA-ASCR Exchange Fellowship of Czech Republic. He is an elected fellow of National Academy of Sciences, India (NASI), apart from several other geoscientific societies, like A.P. Akademi of Sciences, Indian Geophysical Union, Geological Society of India, Indian Society of Earthquake Sciences, and Association of Exploration Geophysicist. He was bestowed with Lifetime Achievement Award from Society of Earth Scientists, India, in 2016.

Chapter 1

Geodynamic and Geologic Evolution of Indian Continent: A Brief History

1.1 Introduction Some 4.6 Ga old Earth is known to be an extremely active physical system, associated with complex evolutionary processes in comparison with other planets. Arguably, it blended out of planetesimals (asteroids), and subsequently underwent differentiation of its interior, which led to development of distinct mantle layers and a core. It was then followed by generation of magnetic field, global plate tectonics and the development of oceans, atmosphere and subsequently, the life. The Earth has a radius of about 6400 km and the equatorial circumference, 40,075 km. The internal parts of the Earth can principally be divided into five major distinguishable regions, viz., inner and outer core, lower and upper mantle and the crust (Fig. 1.1), which are individually characterised by different physical, elastic and geochemical properties. A detailed subdivision of its constituent internal structure is shown in Fig. 1.2. Interestingly, outer core is molten, while the inner core is solid. Out of these layers, the upper mantle is considered buoyant and responsible for the plate-movements. It can be further subdivided into mantle lithosphere, asthenosphere and an underlying transition zone. Above the upper mantle, lies the rigid crust, which are of two different types, oceanic and continental. Oceanic crust is considerably thin, hardly exceeding eight kilometres beneath the ocean floor, while the continental crust is on an average about 35–45 km thick, but it could be as thick as 70 km in orogenic areas, like mountains. It is generated largely by plate tectonic processes.

1.2 Plate Tectonic Processes The plate tectonics concept was formally established by geoscientists around 1960s, although such processes have been operative since Paleo-Mesoarchean era. According to this theory, rigid part of the Earth’s outer layer, known as the lithosphere, which © Springer Nature Switzerland AG 2020 O. P. Pandey, Geodynamic Evolution of the Indian Shield: Geophysical Aspects, Society of Earth Scientists Series, https://doi.org/10.1007/978-3-030-40597-7_1

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1 Geodynamic and Geologic Evolution of Indian Continent: …

Fig. 1.1 Internal divisions of the Earth with approximate depths

Fig. 1.2 Generalised Earth’s deep structure below continental regions

1.2 Plate Tectonic Processes

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Fig. 1.3 Location of major geotectonic plate boundaries, mid oceanic ridges and trenches

is typically about 100–300 km thick, sits and glides over the underlying asthenosphere in the partially molten upper mantle. The lithosphere is broken into several large and smaller-sized regional plates (Fig. 1.3), which move relative to each other at the rates typically around 10–50 mm/year. Plate boundaries are commonly dotted with the occurrence of multiple earthquake events, mostly of smaller magnitudes, along with the topographic features like mountains, mid-ocean ridges, and oceanic trenches. The majority of the world’s active volcanoes occur close to these plate boundaries, for example Pacific ring of fire.

1.2.1 Convergent Plate Margin The tectonic plates, after being pushed from the generation of newer crust at the mid oceanic ridges, are consumed in the upper mantle through ‘subduction process’, operating near the active continental margins, which involves the descent of a cold hydrated oceanic lithosphere (containing a thin cap of oceanic sediments), under another continental tectonic plate (Fig. 1.4) caused by downward convecting limb of a major mantle convection cell (Fig. 1.5) along the deep sea trenches. Downward going slab, can sink even down to a depth of 700 km, before ultimately getting consumed into the upper parts of the mantle. The material lost through this process is roughly balanced by the formation of new oceanic crust along mid oceanic ridges

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Fig. 1.4 Physiographic and geologic features around convergent margins, showing oceanic subduction related processes

Fig. 1.5 Convection currents in the Earth’s upper mantle, related to plate tectonic processes

through seafloor spreading. The path of descent is defined by numerous earthquakes along a plane that is typically inclined between 30° and 60° into the mantle, and is called the Wadati-Benioff zone, or mostly by Benioff zone. Unfortunately, there are not many areas, where the effect of subduction can be seen on the surface, barring a few examples like New Zealand and Japan. A detailed plate tectonic setting of the New Zealand’s active margin is shown in Fig. 1.6 and a section normal to the Bennioff zone across the north island of New Zealand, is shown in Fig. 1.7. It is clear from the figure that the descending plate, bends enormously

1.2 Plate Tectonic Processes

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Fig. 1.6 Regional plate-tectonic settings around New Zealand. Bathymetry is in meters adopted from Lawrence (1967). Boundary between Pacific and Indian plate is shown by dotted line (Walcott 1978). Arrows indicate Pacific plate velocities (in mm/yr) relative to Indian plate based on Chase (1978). Modified after Fig. 1, Reyners (1980)

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Fig. 1.7 A vertical cross section along a micro-earthquake traverse, showing well determined earthquake and micro-earthquakes (Reyners 1978), together with the geometry of of the upper surface of subducted Pacific plate beneath North Island of New Zealand. The open triangles denote earthquakes relocated by Adams and Ware (1977). Modified after Fig. 9, Reyners (1980)

before it subducts into the upper mantle. Magmatism associated with the melting of subducting slabs is responsible for building island arcs and occurrence of volcanoes, which are found over the magmatic arcs. Over these arcs, observed surface heat flow and geothermal gradients are quite high, due to frictional melting of the slab underneath, which is in contrast to characteristically low values over the forearc region. Some essential features of the subduction regime (Fig. 1.4) are the following.

1.2.1.1

Trench

It is a linear topographically steep depression in the ocean basins, typically 5–11 km deep, which marks the boundary between subducting and overriding plates (Fig. 1.6). They are the characteristic morphological feature of convergent plate boundaries, where usually oceanic subducting slab begins to descend under another lithospheric slab. Trenches are normally located at around 200 km away from the volcanic arc (Fig. 1.7).

1.2.1.2

Forearc

This region is located between an oceanic trench and the associated magmatic (volcanic) arc. They lie close to convergent margins and include accretionary wedge and forearc basins. The sediments to these basins are usually derived from the volcanic arc.

1.2 Plate Tectonic Processes

1.2.1.3

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Volcanic Arc

They are also known as magmatic arc that can contain volcanoes. When the downward-moving slab reaches to a depth of about 100–150 km, it gets sufficiently heated, thereby stimulating partial melting of mantle in the plate above the subduction zone (or the mantle wedge), which produces magma, that are predominantly basaltic in composition with largely calc-alkaline affinity. This buoyant magma, after rising through the surface, gives birth to a chain of volcanoes in the overriding plate, known as a volcanic arc.

1.2.1.4

Back Arc

Collapse of the dense slab into the asthenosphere, sometimes gets ‘roll back’ ocean ward and cause extensional stretching into the overlying tectonic plate behind the island arc. This ultimately results into back-arc spreading, where small ocean basins (also known as marginal basins) may be formed. In this region, the crust becomes progressively thinner, and the decompression of the underlying mantle causes the crust to melt, initiating seafloor-spreading processes, somewhat similar to those that occur at ocean ridges.

1.3 Indian Subcontinent and Paleo-Supercontinent Assembly Inter-continental geologic correlation studies have indicated that during the Proterozoic, Indian subcontinent has been an integral part of the two paleo-supercontinents, Columbia and Rodinia. The former is understood to have existed somewhere between 2.5 and 1.6 Ga, while the latter, from approximately 1.3 to 0.9 Ga. The Columbia supercontinent has been associated with subduction related crustal growth during 1.8 and 1.3 Ga via accretion at the continental margins that resulted into formation of magmatic accretionary belts. This supercontinent started fragmenting after 1.6 Ga. In fact, a further accretion and collision of the segments produced by this fragmentation, led to the formation of Rodinia supercontinent, which was fully amalgamated by 1.3 Ga and may have possibly existed till about 0.9 Ga. This was also the period during which collision between East Antarctica and east coast of India took place (around 0.95 Ga), after cessation of the prolonged supra-subduction regime that lasted between 1.6 and 0.95 Ga (Dharma Rao and Reddy 2009; Dharma Rao et al. 2011; Vijaya Kumar et al. 2010; Saha 2011; Saha et al. 2015). It gave rise to collisional features that can still be seen lying east of Cuddapah basin in southern India. Interestingly, during both the supercontinental assembly periods, Indian continent witnessed several accretional, collisional, rifting and continental breakup episodes, apart from massive intraplate magmatism that took place during NeoarcheanPaleoproterozoic period, specially in the Singhbhum and Dharwar cratons. Subsequently at around 750–600 Ma, Rodinia supercontinent too started breaking up and

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Fig. 1.8 Approximate location of the continental segments of West and East Gondwanaland, before their separation during early Jurassic period

then got re-amalgamated at the beginning of the Paleozoic era (somewhere around 570–510 Ma) to form Gondwana Supercontinent. It contained most of the continental land masses, which are presently situated in the southern hemisphere (Fig. 1.8). Interestingly, during all these three assembly periods, eastern coast of India lay close to east Antarctica and thus considered a key element in India-east Antarctica correlation within the Gondwana land assembly (Yoshida et al. 2003; Collins and Pisarevsky 2005; Pandey and Agrawal 2008; Chetty 2014).

1.3.1 Lateral Geologic Correlation Earlier studies revealed that the south Indian shield has been associated with segmentation, and progressive lateral zonation around Dharwar cratonic nucleolus in the Gondwana assembly period (Pandey and Agrawal 2008). The structural and geotectonic patterns across the south Indian shield and some of its neighboring continental masses, are shown in Fig. 1.9, which clearly reveals progressive lateral zonation around the cratonic nucleus of the erstwhile Dharwar craton. In fact, a broad similarity between the major rock units across India, Madagascar, Sri Lanka and east Antarctica, can also be seen. In this figure, zone I, II, III and IV respectively represent cratonic nucleus, late Archean-early Proterozoic crystalline terrain, late Archean early Proterozoic granulite terrain, and Pan-African metamorphic terrain.

1.3 Indian Subcontinent and Paleo-Supercontinent Assembly

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Fig. 1.9 Nature of lateral crustal zonation around western Dharwar craton nucleus. I: Cratonic nuclei of Dharwar and Madagascar, II: Neoarchean—Paleoproterozoic crystalline terrain, III: Neoarchean—Paleoproterozoic granulite terrain, IV: Pan-African metamorphic terrain, 1: granitegneiss and metamorphic complexes, 2: Godavari graben, 3: Eastern Ghats Belt, 4: Mahanadi graben, 5: Lambert rift, 6: Napier complex, 7: Reyner complex, 8: Southern granulite terrain, 9: Achankovil shear zone, 10: Archean—Paleoproterozoic granulites, 11: Archean Dharwar nucleii, 12: Cuddapah basin, 13: Antongil cratonic block, 14: granite-gniess-greenstone block of central Madagascar,15: Ranotsara shear zone, 16: Pan-African granulites of south Madagascar. After Fig. 2, Pandey and Agrawal (2008)

It appears that in Gondwana assembly, the Betsimisaraka Suture Zone of the Madagascar (Collins 2000) extended across the Dharwar craton, and separated Deccan Plateau from the Karnataka Plateau. Cratons are usually considered regions of extreme tectonic stability, containing cool cratonic roots, reaching as deep as 300–350 km (Chapman and Pollack 1977; Polet and Anderson 1995; Artemieva and Mooney 2001; Pandey and Agrawal 1999). But not all the cratons could safeguard themselves against the thermal upwelling and extensional tectonism, for example the Indian shield, which have been under the influence of extensional mode since at least 1500 Ma (Rogers and Callahan 1987). Another example is North China craton. It appears that thermally induced crustal

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reactivations, rise in mantle solidus (Pandey 2016), resulted into large scale influx of volatiles and magmatic material into the lithosphere, thereby making it rheologically weak, warm and less viscous. In presence of already existing weak zones, it made Dharwar craton venerable for the break up in case, it confronted with a rising mantle plume.

1.3.2 Paleo-Super Dharwar Craton and Madagascar Breakup Around 90 Ma, Marion plume hit the SE margin of the Madagascar (Curray and Munasinghe 1991; Storey et al. 1995; Raval and Veeraswamy 2003) (Fig. 1.10), which hosts number of mega shear zones and mobile belts. Below this region, sudden magma upwelling triggered the ridge jump, which was then active between Africa and Madagascar, to the east of Madagascar, thus creating a new rifting phase along India’s western margin. Convective processes, associated with this rifting phase, degenerated and sheared the once thick cratonic root beneath south Indian shield, which ultimately led to the breakup of the Paleo-super Dharwar craton (Agrawal et al. 1992; Pandey and Agrawal 2008; Gibbons et al. 2012; Chetty 2017), part of which is now located as Antongil block in northeastern part of Madagascar. Some authors (Tucker et al. 2011) called the erstwhile Dharwar craton assembly, as “greater Dharwar craton”. Fig. 1.10 Kinship of Madagascar and India before 90 Ma. Star shows the postulated location of Marion Plume (M) before its outburst (Curray and Munasinghe 1991). Circle with dots indicate the area influenced by Marion plume. After Fig. 6, Pandey and Agrawal (2008)

1.3 Indian Subcontinent and Paleo-Supercontinent Assembly

1.3.2.1

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Sri Lanka Connection

Although, the kinship of Madagascar and India in erstwhile Gondwanaland is now well established (Agrawal et al. 1992), the paleo affinity between India and Sri Lanka did not attract much attention. Geographically, Sri Lanka lay quite close to India relative to Madagascar most of the time during the geologic history. However, it has always been questioned whether, (i) Sri Lanka got detached from the Indian subcontinent after the breakup of Gondwanaland, or (ii) underwent rotational and translational movements, or alternatively, (iii) it was always attached to India as a single land mass and travelled together to its present position (Burke et al. 1978; Katz 1978a, b; Crawford 1974; Yoshida et al. 1992). Based on long-wavelength gravity-magnetic, seismological, geological, satellite imagery, geotectonic and geomorphological informations, Agrawal and Pandey (1999a) established that no intra-cratonic rift could ever develop between India and Sri Lanka. Even a faulted boundary or a tectonic rift between southeast Indian peninsular shield and northwest Sri Lanka, appear absent, as revealed by the continuity of long wavelength gravity field (Fig. 1.11) and other geological and geophysical signatures. These two land masses, always accompanied each other after the breakup of Gondwanaland. Possibly a weak mantle upwelling did take place between the two land-masses, which resisted the formation of oceanic crust between them. Before the breakup of Gondwanaland, Sri Lanka rested in the Lutzow-Holm Bay, and its paleo-position relative to India appear to have remained almost same, as it is today, although small rotational movement may be quite possible.

1.4 Geodynamic Evolution of Indian Subcontinent Since Early Cretaceous The Indian subcontinent is bordered by Himalayan mountain belt in the north, Arabian Sea in the west, Indian Ocean in the south and Bay of Bengal in the east. It remained quite active during its entire course of geologic history, having undergone several cycles of deformation, degeneration and restructuring due to underlying geologic processes, specially the geodynamic phenomena, which took place within the Indian ocean during Cretaceous. During this period, Earth passed through an extremely dynamic phase, which was marked by extensive global plate reorganisation and significant increase in spreading rates. This was also the time when the formation of the Indian Ocean took place due to progressive disintegration of the Gondwanaland. It is believed that a persistent thermal anomaly existed below the Gondwana lithosphere for an extremely long period of about 150 Ma, before the actual dispersal of continental fragments took place, which led to massive continental stretching and rifting (Rogers 1993). Apart from this, during 130–65 Ma period, the Indian subcontinent traversed over four mantle hotspots (Reunion, Kerguelen,

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Fig. 1.11 Bouguer anomalies (in mGal) over southeastern part of Indian shield and Sri Lanka (Bouguer Anomaly Map 1976). Extension of gravity anomalies on two sides are shown by dashed lines. 1: Nartamali, 2: Tiruchirappalli, 3: Ginge, 4: Nagapattinam, 5: Jaffna peninsula, 6: Pondicherry. A and B represent gravity anomalies over Palk strait and eastern Vijayan terrain. After Fig. 5, Agrawal and Pandey (1999a)

Crozet, and Marion) in quick succession, which extensively damaged the Indian continental lithosphere, as they were in close proximity to Indian cratonic areas and were also involved with the separation of India from its Gondwanan neighbours (Storey et al. 1995).

1.4 Geodynamic Evolution of Indian Subcontinent Since Early …

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1.4.1 Super Mobility This subcontinent displayed extraordinary mobility after its breakup from Antarctica during early Cretaceous. In fact, it remains the most mobile continent on the surface of the Earth. Between 90 and 53 Ma, its mobility rate reached 20 cm/year, which is the highest among the geotectonic plates, containing continental lithosphere (Jurdy and Gordon 1984). It was also the time, when this continent interacted with two mantle plumes, Marion and Reunion and was also inflicted by large scale magmatic extrusions (Deccan and Rajmahal Traps), besides a possible K-T boundary asteroidal impact offshore near Bombay (Negi et al. 1993), which induced large-scale rifting in nearby regions (Fig. 1.12). Persistent thermal reactivations caused by the presence of higher radioactivity and lithophilic elements since at least 1.5 Ga (Rogers and Callahan 1987), coupled with Late Cretaceous geodynamic events, as stated above, have resulted into higher mantle heat flows, large-scale lithospheric mantle root deformation and extensive degeneration of the Indian continental lithosphere. Various

Fig. 1.12 Location of K-T boundary impact site (solid circle) during super mobile phase of Indian continent, initiation of the Carlsberg ridge and Seychelles break up from India’s west coast. Dotted area represent known extent of Deccan volcanics. Modified after Fig. 9, Pandey et al. (1995)

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studies carried out in the past have confirmed that the Indian lithosphere may be quite thin; only about 100 km (Negi et al. 1986, 1987; Kumar et al. 2007, 2013; Pandey and Agrawal 1999; Pandey 2016; Mandal 2017). This finding was in total departure from the erstwhile thinking that the continental shields are cold, stable and could not be deformed. Currently, the most of Indian cratons may not contain ancient old roots, as they were not strong enough to resist the extensional tectonism. A large number of geotectonic segments in peninsular shield are experiencing repeated uplifts, which has exposed the peneplanation surfaces of earlier cycles to new cycles of erosion (Radhakrishna 1993).

1.5 Major Geological Segments of the Indian Peninsular Shield Earth’s continental crust is known to contain several ancient cratonic blocks, varying in age from Eoarchean to Paleoproterozoic. Besides, there are several Proterozoic and Phanerozoic structures, mobile belts and sedimentary basins, which dot the entire Earth. These ancient cratonic blocks, in due course of time, got merged together and formed Precambrian shield terrains across the globe. It is believed, that by and large, these terrains remained almost stable and undeformed since their inception, barring a few exceptions like the Indian shield and the north China craton, as mentioned before. In order to decipher the evolutionary nature of different geologic terrains and portray their time-bound integrated evolutionary history, one needs to understand chronostratigraphy, for which the knowledge of geologic time scale is essential.

1.5.1 Geological Time Scale It is a global chronostratigraphic scale, which is largely based on Global Standard Stratigraphic age, compiled on the basis of isotopic dates of global geologic events and numerous dating’s of suitable rocks from the geologically well studied sections. A simplified internationally accepted chronostratigraphic chart based on Cohen et al. (2013), is shown in Fig. 1.13. As can be seen, a large part of the Earth’s history (about 88 %) is covered by Precambrian Eon, out of which Archean Era spans for about 1500 Ma; Eoarchean (4000–3600 Ma), Paleoarchean (3600–3200 Ma), Mesoarchean (3200–2800 Ma) and Neoarchean (2800–2500 Ma). Almost all the Indian cratonic nuclei were formed and stabilized during the Archean Era. End of this era was marked by large scale continental crustal production. The Archean Era was followed by the Proterozoic Era, divided into three distinct periods, Paleoproterozoic (2500–1600 Ma), Mesoproterozoic (1600–1000 Ma) and Neoproterozoic (1000–541 Ma). This era has been associated with large scale

1.5 Major Geological Segments of the Indian Peninsular Shield

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Fig. 1.13 A simplified global chronostratigraphic scale, based on compiled global standard stratigraphic ages of geologically well studied sections (Cohen et al. 2013), adopted by International Commission on Stratigraphy (v. 2014/10)

continental rifting, formation of mobile belts, intra-cratonic Proterozoic basin formations, continent-continent collision and mafic dyke swarms intrusion at number of places. It was further succeeded by Phanerozoic Eon, consisting three Era, Paleozoic (541–252.17 Ma), Mesozoic (252.17–66 Ma) and Cenozoic. Several biological mass extinction episodes, took place during this period, resulting into wiping out substantial number of species from the Earth’ surface. Out of these extinction events, Permo-Triassic (~250 Ma) and Cretaceous-Tertiary, K-T (65 Ma) were most severe,

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the latter being responsible for the demise of giant creatures like dinosaurs, whose remains are found in intratrappean layers of Deccan volcanics that cover a large part of western India and surrounding offshore region.

1.5.2 Indian Cratons Indian shield consists five major Eoarchean to Paleoproterozoic cratons (Naqui and Rogers 1987; Ramakrishnan and Vaidyanadhan 2008; Sharma 2009) that include, Dhrawar, Bastar and Singhbhum cratons, situated south of the Narmada–Son Lineament (N-S rift in Fig. 1.14), while the Bundelkhand and Aravalli cratons, are located north of it. This rift structure, situated in the middle of central India, runs for over 1000 km in length and almost divides the Indian continent into two parts. All the major cratonic segments are bounded by rift valleys, Gondwana basins, sutures and mega lineaments, which are repeatedly rejuvenated since Mesoproterozoic times. Continued reactivation/rifting over such a long period of time is not evident anywhere in the world (Rogers and Callahan 1987). Many of its cratonic segments too are undergoing constant rejuvenation, like Singhbhum and Aravalli. The Singhbhum craton is arguably the oldest recognizable cratonic nuclei among all the Indian cratons (Chaudhuri et al. 2018) and perhaps, also one of the most unusual Archean segment on the surface of the Earth. Detailed geological evolution of the individual cratons have been discussed in forthcoming chapters.

1.5.3 Proterozoic Basins The Indian peninsular shield contains number of Proterozoic basins with thick sedimentary succession. They occupy almost 20% area of the Indian Precambrian terrain (Fig. 1.15). They are also popularly known as Purana basins, meaning old basins. These basins are the witness to the early crustal evolution and prelude to the explosion of life during Proterozoic period. In fact, profound microbial activity is documented in Proterozoic siliciclastic environments from the Vindhyan, as well as the Marwar basins in the northern part of India (Kumar and Sharma 2010; Chakraborty 2006). In these basins, the sediments are largely un-metamorphosed and tectonically least disturbed and overlie the metamorphosed Archaean/Palaeoproterozoic basement. Sedimentation in these basins began sometime during the second half of the Paleoproterozoic (around 1.9 Ga) and continued till the end of Neoproterozoics, thus spanning a period of more than 1000 million years. Older sequences of Paleoproterozoics include, sedimentary basins like Bijawar, Gwalior, Abujhmar and Papaghni sub basin of Cuddapah. In comparison, Vindhyan, Chhattisgarh, Pranhita-Godavari, Kaladgi and Cuddapah etc. are considered of mainly Meso to Neoproterozoic age. In some of the basins, thickness of sediments exceeds 5 km. A large part of these basins (specially, Vindhyans, Bhima etc.) can be seen covered under thick Deccan volcanic

1.5 Major Geological Segments of the Indian Peninsular Shield

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Fig. 1.14 Geotectonic segments of the Indian Precambrian shield, showing location of cratonic blocks and associated rift valleys. N-S rift: Narmada-Son Lineament rift system; WDC: Western Dharwar craton; EDC: Eastern Dharwar craton and SGT: Southern granulite terrain. Modified after Fig. 1, Pandey and Agrawal (2008)

rocks. Out of all the Proterozoic basins, Vindhyan and Cuddapah basins are by far the largest and their geological and geophysical characteristics will be discussed in relevant chapters. They are large depositories of limestone, dolomite, phosphorites, barytes and many base metals, including diamonds. Even presence of hydrocarbons are also being looked into these basins.

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1 Geodynamic and Geologic Evolution of Indian Continent: …

Fig. 1.15 Proterozoic basins of India. Modified after Fig. 1, Raha and Sastry (1982).

1.5.3.1

East Coast Sedimentary Basin

A new Proterozoic sedimentary terrain, named as ‘East Coast Sedimentary Basin’, has recently been delineated between the Nellore Schist Belt and the adjoining east coast that may extend further into the offshore region. This basin, detected through DSS studies along the Kavali–Udupi profile (Chandrakala et al. 2017), is associated with velocities 5.3 and 5.5 km/s, which is similar to that of upper and lower Cuddapah sediments (Chandrakala et al. 2013). This Proterozoic basin is separated from the Cuddapah basin by an exhumed horst-like feature that exists below the Nellore Schist Belt. Receiver function studies at PMR (15.10°N, 79.37°E) and PDR (14.30°N,

1.5 Major Geological Segments of the Indian Peninsular Shield

19

79.64°E) seismic stations (Borah 2015), indicated presence of 2–3 km thick low velocity layer (Vs: 2–3 km/s) below the surface which conforms with consolidated sediments. These sediments are also conductive (300–500  m; Naganjaneyulu and Harinaryana 2004) and magnetic anomalies are flat over them. Proterozoic sedimentation thus may have occurred over a much larger area beyond the boundaries of Cuddapah basin during Columbia/Rodinia assembly periods. As mentioned before, during all the supercontinent assembly periods, eastern part of India lay close to east Antarctica. It may appear that the early Proterozoic sedimentation took place simultaneously in both Cuddapah basin, as well as East Coast Sedimentary Basin, when it was a part of erstwhile East Antarctic terrain. Later on, they accreted to India’s eastern margin, during Meso-Neoproterozoic in the form of a deformed accretionary belt like Nellore Schist Belt and Eastern Ghats Belt, as indicated by integrated geological and geophysical studies (Henderson et al. 2014; Chandrakala et al. 2015; Saha et al. 2015). Recent study of Sengupta et al. (2015) concur with this proposition. They indicated that the deposition of similar aged metasedimentary rocks possibly extended further south till Palghat-Cauvery Shear zone in SGT block of south India. This would also mean that the sedimentary rocks of the Ongole Domain may not have been directly derived from the Dharwar Craton alone, but may have affinity with another terrain, like East Antarctica.

1.5.4 Gondwana Basins After wide-spread deposition of Proterozoic sediments, there was a prolonged hiatus in sedimentation for almost 300 Ma. After this interlude, sedimentation again started in a number of sedimentary basins during the Gondwana period, which stretched from early Permian to early Cretaceous (Fig. 1.16). It occurred mostly along linear belts, major river valleys (Damodar, Son, Mahanadi, Godavari, Cauvery etc.), lineaments and reactivated suture zones that bordered various cratonic blocks (Biswas 1999; Ramakrishnan and Vaidyanadhan 2008; Mukhopadhyay et al. 2010). Their extent were possibly far bigger than known today (Chandrakala et al. 2017). These basins preserved a thick succession of un-metamorphosed coal bearing predominantly siliciclastic sediments mostly of terrestrial origin, deposited under a wide range of depositional environment from glacial, glacio-fluvial, glacio-marine, fluvial, lacustrine to shallow marine over a prolonged period of time. As mentioned earlier, these sediments were mostly of terrestrial origin, barring a few sequences of marine intercalations. At some places, these clastic sediments are found to be as thick as 4 km. Major portion of these sediments are confined to Koel-Damodar, Son-Mahanadi and Pranhita-Godavri basins (Fig. 1.16). To begin with, the older sequences (Permo-Triassic) of the Gondwana rocks, such as Talchir-Barakar, got deposited in initial sag basins, formed over the rifted sites having graben or half-graben geometry, even before the actual separation of the Gondwanaland occurred. More than 99% of the total coal reserves of the country are present in such basins. In comparison, the younger equivalents were deposited

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Fig. 1.16 Present day extent of Gondwana sediments in Peninsular India (Dutta and Mitra 1982; Biswas 1999, and recent studies as mentioned in text). 1: Cauvery trough, 2: Palar trough, 3: Vinjamuru Gondwana, 4: Godavari-Krishna Trough, 5: Godavari-Wardha valley, 6: Killari Gondwana sediments, 7: Koyna graben, 8: Aigarh Trough, 9: Mahanadi valley, 10: Damodar-Son valley, 11: Galsi-Malda- Rajmahal Gondwanas, 12: Bengal basin, 13: Frontal zone, Eastern Himalaya, 14: Rewa province, 15: Satpura Province, 16,17: Late Gondwana continental sediments, 18: Saurashtra, 19: Kachchh, 20: Cambay graben

mainly in rift basins like Kachchh, Saurashtra and Cambay located in western India and also some parts of Narmada-Son Lineament zone. The basinal geometry of these basins seems to have been repeatedly modified by tectonic movement during different periods. Recently, some new findings of Gondwana sediments are made from intracratonic settings as well. Untill now, their occurrences are reported mainly from the rifted Gondwana grabens, situated along the rims of Archean-Proterozoic cratonic blocks like Dharwar, Bastar, Singhbhum etc, which were perpetual weak zones, formed consequent to the repeated separation/amalgamation of the various cratonic blocks.

1.5.4.1

Vinjamuru Gondwana Basin

Analysis of shallow seismic data along the Kavali-Udipi DSS profile (as discussed in Chap. 2), indicated presence of 250 m thick Gondwana sediments that stretches for a lateral distance of about 40 km around Vinjamuru region of the Nellore Schist

1.5 Major Geological Segments of the Indian Peninsular Shield

21

Table 1.1 Seismic velocities of Gondwana sediments in different rifted basins of India as derived from deep seismic sounding studies DSS profile

Velocity (km/s)

Source

West Bengal basin

3.8–4.0

Kaila et al. (1992)

Mahanadi basin

3.2–4.0

Kaila et al. (1987), Behera et al. (2002)

Krishna-Godavari basin

2.5–3.5

Kaila et al. (1990)

Between Satpura and Pranhita-Godavari basin

3.20

Mall et al. (2002)

Vinjamuru basin (Nellore Schist Belt)

4.2

Chandrakala et al. (2017)

Belt. Its estimated velocity (4.2 km/s), closely conforms with the seismically derived velocities, associated with the other Gondwana grabens like Bengal, Mahanadi, Krishna-Godavari and central India (Table 1.1).

1.5.4.2

Infratrappean Gondwana Sediments Below Killari

Similar to Vinjamuru Gondwanas, another find of Gondwana sediments has recently been made below Deccan volcanic covered 1993 Killari earthquake region, where they occur as 8 m thick sequence of infratrappean sediment (Parthasarathy et al. 2019) below 338 m thick Deccan basalts in KLR-1 borehole (Gupta et al. 2003). These sediments directly overlie Neoarchean amphibolite to granulite facies basement. Initial palynological studies on the shaly section at the top of the infratrappean sequence, indicated them to be of Precambrian-Cambrian boundary age, and possibly a northward continuation of Neoproterozoic Bhima Group of sediments, found in the close vicinity of Deccan volcanic cover in south India (Ramanujam et al. 1998; Gupta et al. 2003). However, detailed palynological and other integrated geological studies (Pandey et al. 2015; Parthasarathy et al. 2019), revealed presence of characteristic Gondwana palynomorphs in these infratrappeans, which have close proximity with early Permian lower Gondwana sediments of India. These sediments are of non-marine fresh water origin, constituting mainly undeformed calcite-rich carbonate rock (sometimes microcrystalline in nature), with intraclasts of quartz and fine grained phyllosilicates.

1.5.4.3

Cauvery Basin Gondwanas

Prasad and Phor (2009) reported presence of 7 m thick lower Gondwana sedimentary succession, resting directly over the Archean basement in PD-B well, drilled in Tanjore sub-basin of the Cauvery basin. It is shown as patch 1 in Fig. 1.16. As per drilled information, rapid denudation seems to have almost totally removed the deposited Gondwana sediments from this area, as well other regions like Killari and Koyna.

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1.5.5 Tertiary Basins Most of the Tertiary basins in India are located along the margins of the peninsular shield, or adjacent offshore regions, where they are mainly formed due to marine transgression, resulting into the deposition of thick sedimentary successions (Misra 1987). There are large number of such basins, prominent ones include Cambay, Kachchh-Saurashtra and Barmer situated in the northwest, Bombay offshore, KeralaKonkan near the western coast, Ganga basin in the foothills of Himalaya in the north, Upper Assam basin in Northeast and Cauvery, Krishna-Godavari, Mahanadi and west Bengal basins, near the eastern margin. Many of them produce hydrocarbons on commercial basis like upper Assam, Cambay, Mumbai offshore and KrishnaGodavari basins. On continents, most of these basins are located over the rift and graben structures, developed due to extensional tectonics, which were very pervasive since the breakup of Gondwana assembly. In these basins, deposition of sediments took place after the subsidence of basement blocks, associated with nearly vertical faults. Many times, they have been deposited over the top of Cretaceous as well as older Gondwana sediments (Misra 2018). Salient features of some of the geotectonically prominent basins are given below.

1.5.5.1

Cambay Graben

This narrow intra-cratonic petroliferous basin is situated in alluvial plains of Gujarat and covers an area of about 56,000 km2 . It is one of the most extensively explored sedimentary basin in India that extends between latitude 21°N–25°N and longitude 71°30 E to 73°30 E and runs for about 425 km (Mohan et al. 2008). It is formed by discontinuous normal faults running in an approximately NNW-SSE direction, before taking a swing towards NNE-SSW direction near the Gulf of Cambay. The geological sequence comprises mainly of volcanic conglomerates, greywackes, dark grey to black grey shales, siltstones, fine to medium grained sandstone and claystones, with minor Jurassic and Cretaceous outcrops on the basin margin, indicating that the initial rifting may have began sometimes during the Mesozoic period (Sen Gupta 1967; Mathur et al. 1968; Biswas 1987).

1.5.5.2

Mumbai Offshore Basin

This pericratonic basin, situated over the submerged Bombay platform located just south of the Surat depression, came into existence during end of Cretaceous due to faulting in the Deccan basalt basement. Geotectonically, this basin may well be a continuation of the oil/gas bearing Cambay graben in north, having similar thermal history (Pandey and Agrawal 2000; Ganguli et al. 2018). Towards south, it continues up to about 18°S. It contains a large number of giant oil/gas fields. Among them,

1.5 Major Geological Segments of the Indian Peninsular Shield

23

Bombay high is the most prominent oil/gas producer in India, which is a broad regionally upwarped region, formed during Paleogene-Neogene times. This region was probably hit by a massive asteroid around 65 Ma, which caused, Seychelles breakup and triggered Deccan volcanic eruption over the western India as well as adjacent offshore regions, besides inducing K-T boundary biological mass extinction and demise of mighty dinosaurs (Negi et al. 1992, 1993; Pandey et al. 1995; Chatterjee et al. 2006). Measured temperature gradients are very high in the Bombay offshore region, sometimes exceeding 70 °C/km (Pandey and Agrawal 2000).

1.5.5.3

Gangetic Plain

This foreland basin is almost 1000 km in length and extends from the ArravalliDelhi ridge in the west to Rajmahal hills in the east, occupying an area of about 250,000 km2 . Its formation was probably initiated due to a fore-deep sag, formed synchronously around the same time as Himalayan uplift. Sedimentation in the basin started around early Miocene, for which sediments were provided by rising and eroding Himalaya, most of which have been carried away to Bengal fan. No marine transgression affected this basin throughout its depositional history (Singh 1996). Tertiary sediments of Ganga, Kachchh, Konkan-Kerala and Rajasthan, have shown poor source rock development for hydrocarbons.

1.5.5.4

Upper Assam Basin

This is a poly-history basin, which is producing hydrocarbons since more than a century now. It is bounded by three major thrust faults viz. Himalayan orogenic belt in the north, Mishmi Thrust in the east and Schuppen Belt in south (Mandal and Dasgupta 2013). The resultant compressional force is probably still active due to relative motion of the plates, which appears to have influenced the neotectonic regime in the fore-land part of the basin. This basin is primarily comprised of sand and shale alterations of Paleocene/Eocene to Recent age. These Tertiary sequences rest over the granitic basement (Misra 1987). It contains number of oil fields where Miocene Tipam formation situated in the northeastern part, is a prominent hydrocarbon potential rock.

1.5.6 Deccan Traps About 66 Ma old tectonically active Deccan Large Igneous Province (LIP), constitutes one of the largest flood basaltic eruptions on surface of the earth (Courtillot et al. 1988; Duncan and Pyle 1988; Pande et al. 1988; Renne et al. 2015; Schoene et al. 2015). It covers almost half a million sq. km on both onshore and offshore, across the western margin and central India (Fig. 1.17). It consists of a number of basalt flows with a cumulative thickness exceeding 2 km in Western Ghats region.

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1 Geodynamic and Geologic Evolution of Indian Continent: …

Fig. 1.17 Extent of Deccan Volcanics in western and central India, erupted across K-T boundary

The constituent flows (as many as 46), are highly differentiated, partially eroded at the surface and characterized mainly by tholeiitic composition. They seem to have erupted in a quick succession during less than one million years (Courtillot et al. 1988; Renne et al. 2015) across Cretaceous–Tertiary (K-T) boundary. Palaeomagnetic data put these formations in chron 29 R (Schoene et al. 2015). Subsequent studies (Radhakrishna et al. 2019) based on paleomagnetic data, confirmed that these volcanics erupted only in a single geomagnetic polarity transition period. Further, its eruption nearly coincides with the K-T mass extinction, which was possibly most severe in the Earth’s evolutionary history that caused demise of mighty dinosaurs. Interestingly, this region has been experiencing significant intraplate earthquake activity since historical times. Many of them were highly destructive in nature, like the Koyna earthquake of December 10, 1967 (Mw 6.3), Latur-Killari earthquake of September 29, 1993 (Mw 6.3), Jabalpur earthquake of May 21, 1997 (M 6.0) and Bhuj earthquake of January 26, 2001 (Mw 7.7) and many more (Pandey 2009).

1.6 Regional Geophysical Studies As evident, Indian shield is much more dynamic compared to all the other similar terrains in the world, having undergone large scale deformation and degeneration at crust and lithospheric mantle level, due to sustained thermo-geodynamic perturbations throughout its geological history (Pandey and Agrawal 1999; Pandey et al.

1.6 Regional Geophysical Studies

25

2013; Pandey 2016). Its original texture and composition is no more the same, as revealed by deep scientific drillings at Killari and Koyna. In fact, bulk of the continental crust in the Indian shield appear to be mafic and made up of low to high-grade metamorphic rocks. Historically, it has been associated with moderately active seismicity, not seen elsewhere in geologically stable terrains. Such unusual characters are very well reflected in geophysical fields, which principally deal with the systematic collection of the data to study the structure and composition of the Earth’s interior, including subsurface layers. Various methods come under the gamut of geophysics, which are, seismic, seismological, gravity, magnetic, electrical, electromagnetic and heat flow etc. Since rock density and in situ velocities in subsurface structural features are quite closely related, their dispositional nature can be better studied, when we combine seismic and gravity data together. Both the methods can also help in elucidating density inhomogeneities and composition inside the crust, specially at deeper levels, which are normally not always possible through seismics only. In comparison, electromagnetic studies can provide auxiliary information about the crustal conductivity and presence of fluids. In a Similar manner, deeper thermal regime of the crust and lithospheric mantle can be studied through heat flow and radioactivity.

1.6.1 Gravity Studies Gravity measurements in India has a long history. It possibly started in 1865 by the Survey of India, which made a few measurements of absolute gravity, using two brass pendulums. Systematic survey for geodynamical and Earth resource studies however began only around 1950s. Utilizing about 3000 gravity observations, the first Gravity Map of India, including parts of neighbouring countries, was produced on a scale of 1:12,000,000 at 20 mGal contour interval (Gulatee 1956). Later on after the establishment of reliable N-S gravity base line network (Manghnani and Woollard 1963), many institutions like the Survey of India, CSIR-National Geophysical Research Institute (CSIR-NGRI), Geological Survey of India (GSI), Oil and Natural Gas commission and Hawaii Institute of Geophysics (USA), collected vast amounts of gravity data, which led to the publication of First Gravity Map Series of India by CSIR-NGRI in 1975, using 30,050 gravity observations at the scale of 1:5,000,000, with 10 mGal contour interval. In view of international standardisation and further acquisition of large amount of data, a new Gravity Map Series of India—2006 has now been published by CSIR-NGRI and GSI, on a 1:2,000,000 scale, with 5 mGal contour interval. This new map series, is based on 51,356 gravity station values, sampled at 3 min arc interval from a total of 143,786 stations, distributed all over the India. A recently released Bouguer anomaly image map of India (GSI-NGRI 2006) is reproduced in Fig. 1.18, which shows several prominent gravity anomalies, attributed to density inhomogeneities at different depth that have been described in detail by Mishra et al. (2008). A conspicuous feature of the map is the presence of positive gravity field in areas lying north of Narmada Son Lineament-Godavari graben line, that contain

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1 Geodynamic and Geologic Evolution of Indian Continent: …

Fig. 1.18 Bouguer anomaly map of India. ADMB: Aravalli-Delhi Mobile Belt; BB: Bangladesh; BC: Bundelkhand Craton; DVP: Deccan Volcanic Province; EDC: Eastern Dharwar craton; EGMB: Eastern Ghats Belt; NSL: Narmada-Son Lineament; SGT: Southern Granulite Terrain; SMB: Satpura Mobile Belt; SP: Shillong Plateau; WDC: Western Dharwar craton. After Fig. 3, Tiwari et al. (2013)

cratons like Aravalli, Bundelkhand, Singhbhum and Bastar, which were repeatedly reactivated time and again. Regions south of this dividing line, is solely occupied by the Dharwar craton, associated with gravity low. The terrains lying either side of this dividing line, exhibit distinct geological and geophysical signatures (Pandey and Agrawal 1999). Some major Bouguer gravity features of this map (Tiwari et al. 2013) are, (i) gravity lows due to crustal thickening below the Himalaya and Ganga basin sediments, (ii) linear gravity highs, related to Proterozoic mobile belts like Aravalli-Delhi Mobile Belt, the Satpura Mobile Belt and the Eastern Ghats Belt (Mishra 2006), that represent Proterozoic collision zones. Short wavelength gravity highs, in NW India, west of the Aravalli-Delhi Mobile Belt, possibly represent multiple episodes of volcanic

1.6 Regional Geophysical Studies

27

activity that began in Meso-Neoproterozoic period (Sinha Roy 1988; Ravi Kumar et al. 2013) or alternatively, caused due to the burst of Reunion mantle plume, K-T impact in offshore Mumbai (Negi et al. 1992), Seychelles breakup (Pandey et al. 1995) and young Tertiary intrusions, that followed such events. However, the most conspicuous feature of this gravity map remains the presence of a high order long wavelength gravity low over the elliptical shaped splitted nucleus of the Dharwar craton (Radhakrishna and Naqvi 1986; Agrawal et al. 1992; Tucker et al. 2011), which is a major regional feature. In this region, at some places, measured gravity anomalies reach around −120 mGal with high gradient contours around the western continental margin. This gravity low also coincides with the anomalous geoidal gravity low over the southern India and the adjacent south Indian ocean (Ghosh et al. 2017). It is now well known that until ~90 Ma, India and Madagascar were a single unit as paleo-super Dharwar craton in the erstwhile Gondwanaland (Agrawal et al. 1992; Storey et al. 1995; Raval and Veeraswamy 2003). It is conjectured, that due to rifting, Dharwar cratonic nucleus was broken, a part of which now forms Antongil cratonic block of the northeast Madagascar, as mentioned before. The incomplete closure of the gravity anomalies on the Indian side apparently join well with the long-wavelength contours on the Madagascar side (discussed in Chap. 6). The high-gradient contours on both sides represent rifted continental blocks, possibly uplifted on either side.

1.6.2 Deep Crustal Seismic, Tomographic and Receiver Function Studies 1.6.2.1

DSS Studies

Understanding the geodynamic evolution of any terrain, requires detailed knowledge of the structural and compositional variations inside the crust and underlying mantle. In this context, Deep Seismic Sounding (DSS) or controlled source seismic studies, can be considered one of the most effective geophysical tool to delineate deeper as well as shallow crustal structures using suitable data sets. In such studies, wide angle reflections/refractions data are used to delineate structural and regional velocity variations in the crust. The first such profile was shot in 1972 through a collaboration between CSIR- NGRI (Hyderabad, India) and the Ukranian Academy of Sciences in the former USSR. It traversed between Kavali on the east coast to Udupi on the west coast, encompassing all the major geological units of southern Indian shield like Eastern Ghats Belt (EGB), Nellore Schist Belt (NSB), Cuddapah basin (CB), Closepet granite in the eastern part of the Dharwar craton and Chitradurga Schist Belt in the western part (Kaila et al. 1979; Mall et al. 2008, 2012; Chandrakala et al. 2013, 2015, 2017; Pandey et al. 2017). Since then, such data have been acquired for about 5000 km along 35 profiles, which criss-cross variety of geologic and tectonic settings (Fig. 1.19), covering Kash-

28

1 Geodynamic and Geologic Evolution of Indian Continent: …

Fig. 1.19 Location of deep seismic profiles across different geological terrains of India

mir Himalaya, Archaean-Proterozoic peninsular shield, Southern Granulite Terrain, Deccan volcanic Province, Narmada-Son lineament and Gondwana and Tertiary basins. These studies revealed number of new subsurface information, which have strong bearing on the time-bound evolution of Indian terrain. Results obtained along these profiles are discussed in various chapters together with other geophysical data sets.

1.6 Regional Geophysical Studies

1.6.2.2

29

Seismic Tomographic and Receiver Function Studies

Seismic tomographic studies in India started sometime in late eighties through a study by Iyer et al. (1989), who analysed P-wave residuals and stated that a high velocity anomaly exists in the depth range of 100–400 km beneath the Deccan Volcanic Province (DVP), thereby implying presence of a coherent, colder and comparatively rigid lithospheric root extending to a depth of about 400 km beneath this region. This was followed by a large number of similar studies like, Srinagesh et al. (1989) and Ramesh et al. (1990, 1993), who also found similar results for the south Indian shield. Several studies that further followed, like Rai et al. (1992) and Ravi Kumar and Mohan (2005), too held similar views. However, such perceptions have been changing slowly due to emergence of new results from recent tomographic and receiver function studis (Mitra et al. 2006; Priestley et al. 2006; Priestley and Mckenzie 2006; Kumar et al. 2007, 2013; Jagadeesh and Rai 2008). Recently Koulakov et al. (2018) also presented a new regional tomography model of the upper mantle based on P-wave travel time tomography, which depicted the lithosphere structure of India. This tomographic study indicated high Vp anomaly beneath the northern part of India extending down to 200 km, which conforms with the presence of a thick continental lithosphere. On the other hand, beneath southern India, the high Vp anomaly was found to be less prominent indicating a thinner lithosphere. Similarly around the western margin also, the lithosphere was found to be thinner (~100 km) due to its degeneration by the Reunion-Deccan hot spot.

1.6.3 Heat Flow and Lithospheric Studies Indian subcontinent is well covered by heat flow studies. So far, heat flow has been measured at about 174 sites, located on geologically distinct terrains, like Archean cratons and fold belts, Proterozoic sedimentary basins, Gondwana and Tertiary grabens and Deccan Volcanic Province, which has been reviewed in detail by Gupta and Rao (1970), Pandey and Negi (1995), Roy and Rao (2000) and Nagaraju et al. (2017). Regionally, the measured terrestrial heat flow varied from as low as 23.0 mW/m2 in Eastern Dharwar craton to as high 107.0 mW/m2 in Son-Mahanadi graben. World-wide Archean terrains are known to be associated with extremely low heat flow values, but quite a few Indian Archean cratons, specially those located in the northern part of India like Singhbum, Bastar and Aravalli, are characterised by reasonably high heat flow values; all of them located in the close vicinity of episodically active mobile belts. Details of regional heat flow distribution, are discussed in various chapters and summarised in Chap. 9. In general almost all the shield terrains in the world, like Canadian, west Australian, west African, Baltic and Siberian, are associated with thick lithosphere of about 250–350 km (Chapman and Pollack 1977; Polet and Anderson 1995; Artemieva and Mooney 2001), but perceptions about thickness of Indian lithosphere has been differing from one study to another. Many researchers (Iyer et al. 1989; Srinagesh

30

1 Geodynamic and Geologic Evolution of Indian Continent: …

et al. 1989; Ramesh et al. 1990, 1993; Gupta et al. 1991; Rai et al. 1992; Gupta 1994; Ravi Kumar and Mohan 2005) feel that the Indian shield cratonic roots are very thick, reaching as much as 400 km, while many others have reported lithospheric thickness of about 100–150 km (Negi et al. 1986, 1987; Pandey and Agrawal 1999; Priestley et al. 2001; Mitra et al. 2006; Priestley and Mckenzie 2006; Kumar et al. 2007, 2013; Mandal 2017). Results from deep drilling at Koyna and Killari, however conform with thin and warm lithosphere (Pandey 2016).

1.6.4 Magnetotelluric Studies The magnetotelluric (MT) method, which is one of the modern techniques in geophysics, is widely used to understand subsurface shallow as well as deeper crustal structure. This technique was initially used in India way back during 1970s and its use has been steadily growing since then to study various geological problems in various parts of India (Harinarayana et al. 2003; Naganjaneyulu and Harinaryana 2004; Harinarayana 2008; Shalivahan et al. 2014; Abdul Azeez et al. 2015, 2018; Malleswari et al. 2019). This included delineation of buried sedimentary structures and geothermal reservoirs below the volcanic covered areas of western margin, Saurashtra, Kutch, central India and western Himalayas. These studies suggest that the upper crustal resistivity could be as high as 30,000–40,000  m, in some terrains of peninsular India (Naganjaneyulu and Harinarayana 2003) in comparison to generally less resistive mid and lower crust. Similarly, major shear and deep faulted zones can be seen associated with anomalous high conductive features (Harinarayna et al. 2003). Besides crustal studies, MT techniques are now also being used to study the fluid related conductive features in seismically active areas, for example, Koyna, Killari and Bhuj (Gupta et al. 1996; Ramaprasad Rao et al. 2003; Sarma et al. 2004; Abdul Azeez et al. 2018). Attempts have also been made to delineate lithosphereasthenosphere boundary (LAB) below different geotectonic blocks, which varied from 95 km to more than 200 km (Gokarn et al. 2004; Naganjaneyulu and Santosh 2012; Shalivahan et al. 2014; Abdul Azeez et al. 2015; Malleswari et al. 2019). Due to recent developments, this method is now being considered at par with other geophysical techniques like gravity and seismics.

1.6.5 Seismicity Earthquakes are usually caused by faulting and sudden lateral or vertical movement of rocks along the ruptured or faulted surface (Fig. 1.20). Indian subcontinent is identified by several such faults, associated with rift valleys, sutures, and mega lineaments (Fig. 1.14), which have been persistently active. Besides, Indian crust is highly evolved, weak, inhomogeneous and non-rigid due to mantle-derived fluid led

1.6 Regional Geophysical Studies

31

Fig. 1.20 Parameters associated with earthquakes along a fault zone

metasomatic alteration (Tripathi et al. 2012a, b; Pandey et al. 2014, 2016; Vedanti et al. 2018). Consequently, GPS and seismological studies indicate occurrence of very high strain rates in the Indian shield terrain (Johnston 1994; Ramalingeswara Rao 2000; Paul et al. 2001; Talwani and Gangopadhyay 2001), in comparison to many other stable continental regions of the world (Johnston 1994). Not surprisingly therefore, it has been experiencing moderate intraplate earthquake activity since historical times (Ramalingeswara Rao 2000, 2015; Agrawal and Pandey 1999b; Kayal 2008; Pandey 2009) (Fig. 1.21), many of them being highly destructive in nature and occurred over Deccan volcanic terrain. As can be seen from the Fig. 1.21, many of the earthquakes are located at the intersection of two or more faults. Although, there has been some progress in understanding the source mechanism of some of these earthquakes (Mandal et al. 2000; Kayal 2007, 2008), the cause of recurring seismic activity in intraplate region of the Indian shield is not much understood, which is often attributed to the NNE compressive forces, generated due to collision of northward moving Indian plate with Eurasia. This could be a major factor in seismogenesis, but not all of them are related to plate tectonic forces only, as discussed in detail in Chap. 8.

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Fig. 1.21 Distribution of earthquake epicenters (1700–1997) on tectonic map of Indian Peninsular shield. Quite a few earthquakes occur close to the intersection of two or three faults or major lineaments. CB: Cuddapah Basin, GG: Godavari graben, EG: Eastern Ghats Belt, T.Z: Transition zone, Sa: Satpura, Sv: Son-Valley, K: Koyna, Ki: Killari, T: Tapti, Ja: Jabalpur, NSL: Narmada-Son Lineament, CG: Closepet granite, GBF: Great Boundary Fault, ADFB: Aravalli Delhi Fold Belt. Modified after Fig. 2, Ramalingeswara Rao (2000)

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Schoene B, Samperton KM, Eddy MP, Keller G, Adotte T, Bowering SA, Khadri SFR, Gertsch B (2015) U-B geochronology of the Deccan traps and relation to the end-Cretaceous mass extinction. Science 347(6218):182–184 Singh IB (1996) Geological evolution of Ganga Plain-An overview. J Palaeontol Soc India 41:99– 137 Sinha Roy S (1988) Proterozoic Wilson cycles in Rajasthan. In: Precambrian of the Aravalli Mountain, Rajasthan, India. Mem Geol Soc Ind 7, 95–108 Srinagesh D, Rai SS, Ramesh DS, Gaur VK, Rao CVR (1989) Evidence for thick continental roots beneath south Indian shield. Geophys Res Lett 16:1055–1058 Storey M, Mahoney JJ, Saunders AD, Duncan RA, Kelley SP, Coffin MF (1995) Timing of hot-spot related volcanism and the breakup of Madagascar and India. Science 267:852–855 Talwani P, Gangopadhyay A (2001) Tectonic framework of the Kachchh earthquake of 26 January 2001. Seism Res Lett 72:336–345 Tiwari VM, Ravi Kumar M, Mishra DC (2013) Long wavelength gravity anomalies over India: crustal and lithospheric structures and its flexure. J Asian Earth Sci 70–71:169–178 Tripathi P, Pandey OP, Rao MVMS, Koti Reddy G (2012a) Elastic properties of amphibolite and granulite facies mid-crustal basement rocks of the Deccan volcanic covered 1993 Latur-Killari earthquake region, Maharashtra (India) and mantle metasomatism. Tectonophysics 554–557:159– 168 Tripathi P, Parthasarathy G, Ahmad SM, Pandey OP (2012b) Mantle derived fluids in the basement of the Deccan traps: Evidence from stable carbon and oxygen isotopes of carbonates from the Killari borehole basement, Maharashtra, India. Int J Earth Sci 101:1385–1395 Tucker RD, Roig J, Delor C, Amelin Y, Goncalves P et al (2011) Neoproterozoic extension in the Greater Dharwar Craton: a reevaluation of the “Betsimisaraka suture” in Madagascar. Can J Earth Sci 48:389–417 Vedanti N, Malkoti A, Pandey OP, Shrivastava JP (2018) Ultrasonic P- and S- wave attenuation and petrophysical properties of Deccan flood basalts, India as revealed by borehole studies. Pure Appl Geophys 175:2905–2930. https://doi.org/10.1007/s00024-018-1817-x Vijaya Kumar K, Ernst WG, Leelanandam C, Wooden JL, Grove MJ (2010) First Paleoproterozoic ophiolite from Gondwana: geochronologic–geochemical documentation of ancient oceanic crust from Kandra, SE India. Tectonophysics 487:22–32 Walcott RI (1978) Present tectonics and Late Cenozoic evolution of New Zealand. Geophys J R Astr Soc 52:137–164 Yoshida M, Funaki M, Vitanage W (1992) Proterozoic to Mesozoic east Gondwana: the juxtaposition of India, Sri Lanka and Antarctica. Tectonics 11:381–391 Yoshida M, Jacobs J, Santosh M, Rajesh HM (2003) Role of Pan-African events in the Circum-East Antarctic Orogen of East Gondwana: a critical overview. In: Yoshida M, Windley BF, Dasgupta S (eds) Proterozoic East Gondwana: Supercontinent assembly and breakup. Geol Soc Lond Spec Publ vol 206, pp 57–75

Chapter 2

Dharwar Craton

2.1 Introduction Out of the five Archean cratons of the Indian peninsular shield (Fig. 1.14), Dharwar craton is one of the oldest and the largest, covering almost entire terrain of southern India. It is largely made up of granite-gneiss-greenstone belts. On its northern side, it is surrounded by Deccan volcanics, in the northeast by the Karimnagar granulite belt, and in the east by Proterozoic Eastern Ghats Belt. It is extensively studied geophysically and geologically, compared to all other Indian cratons. Broadly, this craton can be divided into three individual Precambrian geotectonic blocks (Fig. 2.1); western Dharwar craton (WDC), eastern Dharwar craton (EDC) and southern granulite terrain (SGT), which exhibit distinct geologic, tectonostratigraphic, magmatic and geochronologic characters (Radhakrishna and Naqvi 1986; Naqvi and Rogers 1987; Ramakrishnan and Vaidyanadhan 2008; Sharma 2009; Sunder Raju et al. 2014; Anil Kumar et al. 2015; Jayananda et al. 2018).

2.2 Geological Settings Out of the three geotectonic blocks, WDC of Meso-Neoarchean age, is probably the oldest containing vast exposures of high-grade greenstone schist belts (3.0–3.4 Ga), granites and TTG gneisses, that exhibit low temperature metamorphism (Radhakrishna and Naqvi 1986; Naqvi and Rogers 1987). It is believed that this particular segment of the Indian shield, may be devoid of Proterozoic tectono-thermal events, although, it may not necessarily be true. WDC is separated from the adjacently situated EDC by a steep mylonitic zone along the Chitradurga fault region (CDF in Fig. 2.1), which is a major geotectonic shear zone, dividing the Dharwar craton into two major blocks of differing geologic characters. Unlike the WDC, EDC domain basically evolved during various episodes from 2.7 to 1.1 Ga, the last one being a major kimberlitic event (Anil Kumar et al. 1993). © Springer Nature Switzerland AG 2020 O. P. Pandey, Geodynamic Evolution of the Indian Shield: Geophysical Aspects, Society of Earth Scientists Series, https://doi.org/10.1007/978-3-030-40597-7_2

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Fig. 2.1 Broad geotectonic features of the Dharwar craton, south Indian shield. WDC: Western Dharwar craton, EDC: Eastern Dharwar craton, SGT: Sothern Granulite Terrain, CB: Cuddapah basin, NSB: Nellore Schist Belt, EGB: Eastern Ghats Belt, CDF: Chitradurga Fault. 1: GoaJadcherla MT profile, 2: Chikmagalur–Kavali MT profile, 3: MT profile across WDC and adjoining Coorg block (Fig. 2.6), 4: MT profile across Cuddupah basin and adjacent regions (Fig. 2.14), 5: Chikmagalur-Perur DSS traverse (Fig. 2.3), 6: Kavali-Udupi DSS traverse covering all the geotectonic segments of Dharwar craton

This part of the Dharwar craton is characterised by Neoarchean–early Proterozoic (∼2.5 Ga) cratonic growth, with massive remoblisation of several crustal blocks, containing abundant calc-alkaline to K-rich granitoids along with thin elongated greenstone belt (Harish Kumar et al. 2003; Manikyamba et al. 2008, 2014). It is conspicuously overlain by highly deformed mighty Cuddapah basin, which is one of the largest Proterozoic basin of Peninsular India, characterised by complex geodynamic history (Pandey et al. 2018). This N–S trending crescent-shaped basin is located between Lat. 13° and 17° N and Long. 77° 45 and 80° 15 E and covers an area of almost 44,500 km2 . This basin which extends for over 440 km in length and 150 km in width, is bounded by deep-seated faults, containing thick succession of igneous and Proterozoic sedimentary rocks of the Cuddapah and Kurnool Supergroups, which rest directly over the Neoarchean granitic-gneissic basement (Saha and Tripathy 2012). It evolved through several phases of extension, magmatism (lava flows and sills) and sedimentation, which began around 1.9 Ga or little earlier (Mallikarjuna Rao et al. 1995; Anand et al. 2003; Saha et al. 2015). The eastern half of the basin is made up of folded rocks, commonly known as Nallamalai Fold Belt. This fold belt is bounded on its eastern side by a 300 km long Nellore Schist Belt, which contains deformed Paleoproterozoic to Mesoproterozoic volcano-sedimentary succession, divided into four major units: (i) Vinjamuru Group, (ii) 1.9 Ga Kandra Ophiolite Complex, (iii)

2.2 Geological Settings

43

1.3 Ga Kanigiri Ophiolite Melange, and (iv) Udayagiri group (Saha et al. 2015). Out of these units, the origin of both the Kandra Ophiolite Complex and Kanigiri Ophiolite Melange, have been attributed to supra-subduction settings (Dharma Rao and Reddy 2009; Vijaya Kumar et al. 2010; Dharma Rao et al. 2011; Saha et al. 2015) active during Mesoproterozoic. This schist belt, which underwent multiple deformation and metamorphism till as late as 500 Ma (Dobmeier et al. 2006), can be seen thrusted over the Cuddapah basin. Nellore Schist belt is separated from the Nallamalai Fold Belt by the Vellikonda thrust front (Venkatakrishnan and Dotiwala 1987; Saha 2002; Saha et al. 2010). Recently, a 40 km long thin stretch of Gondwana sediments (Vinjamuru Gondwanas) are also discovered around Vinjamuru, indicating a late stage rifting of this region (Chandrakala et al. 2017). In comparison, Eastern Ghats Belt, lying further east of the Nellore Schist Belt, is a 1000 km long deformed granulite facies terrain that parallels India’s east coast. Its southern part, consisting of Ongole domain, is made up primarily of poly-cyclic 1.7–1.6 Ga old granulite facies rocks (Dobmeier and Raith 2003; Henderson et al. 2014), that accreted to Indian land mass during Neoproterozoic to early Cambrian (Dobmeier and Raith 2003; Saha et al. 2015). Quite likely, Eastern Ghats Belt terrain is absent between Nellore Schist Belt and Kavali town near the east coast, however, a deformed Proterozoic basin may be present in lieu of the Eastern Ghats Belt. The region that lies east of the Cuddapah basin, have played a major role in the evolution of the southern Indian shield. It used to be an active continental margin during Paleoproterozoic to Mesoproterozoic era, having undergone oceanic subduction, continent-continent collision and multi-stage accretional growth (Saha 2002, 2011; Dobmeier and Raith 2003; Chandrakala et al. 2013, 2015; Dasgupta et al. 2013; Henderson et al. 2014; Pandey et al. 2018). Consequent to such a long sustained geodynamic upheaval, it houses many diversified and deformed geological terrains, which have a special place in India-east Antarctica correlation within the Gondwana land assembly (Yoshida et al. 2003; Collins and Pisarevsky 2005; Pandey and Agrawal 2008; Chetty 2014). In fact, the entire eastern margin is marked by tectonomagmatic features that record the events related to the fragmentation of Columbia and Rodinia assembly (Rogers and Santosh 2002; Zhao et al. 2004). In comparison to both WDC and EDC, the SGT, which is situated at the southernmost end of Indian peninsula (below latitude 12.5°N), has undergone extensive regional metamorphism during late Archean and Pan African times (550 ± 30 Ma). It exposes one of the largest Precambrian lower crustal sections in the world, corresponding to P–T regime of 5–12 kb and 700–900 °C (Newton and Perkins 1982; Rao et al. 2003). It consists mainly of deformed Archean/Neoproterozoic high-grade metamorphic and magmatic rocks, that include tonalitic gneisses, migmatites with high-grade assemblages of garnet-bearing felsic-mafic granulites, charnockites and granitoids (Yellappa and Rao 2018). Although, it has traditionally been divided into six distinct crustal blocks namely, (i) the Northern Block, (ii) the Nilgiri Block, (iii) Salem–Madras Block, (iv) Cauvery Suture Zone (CSZ), (v) the Madurai Block, and (vi) the Trivandrum Block, but primarily, it can be considered having only two distinct terrains, situated either side of the CSZ. North of this suture zone, the terrain is made up of Neoarchean–early Proterozoic granulites, consisting charnockites and

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retrograde gneisses, while the southern block south of CSZ, largely contains PanAfrican (550 ± 30 Ma) rocks, like charnockites, khondalites, granites and gneisses. Some segments of this terrain, are tectonically active having been associated with large uplift, like Nilgiri domain.

2.3 Geophysical Studies Divergent views persist regarding the nature of crustal evolution and composition of the Archean and Proterozoic shield terrains world-wide (Durrheim and Mooney 1991, 1994; Rudnick and Fountain 1995; Rudnick and Gao 2003; Julia et al. 2009). In general, it is believed that the Archean crust is largely stable and have a thinner crust due to absence of high velocity layer at the bottom of the lower crust than the Proterozoic terrains, which are often reworked. Numerous studies have now indicated that the present-day continental crust may have evolved through arc magmatism and accretion, rifting, continent-continent collision and basaltic magma extrusion, but whether these processes were operative in the Archean era, remains a subject of considerable speculation. Since, Indian shield terrains too underwent sustained accretion, magma infusion and massive remoblisation of crustal blocks, its subsurface crustal configuration and composition have been probed through various geophysical studies.

2.4 Western Dharwar Craton 2.4.1 Crustal Seismic Structure Two deep seismic profiles, (i) Kavali-Udupi and (ii) Chikmagalur-Perur (Fig. 2.1), have earlier been shot across the Dharwar craton to study underlying crustal structure. Both of them, covered western as well as eastern parts of this craton.

2.4.1.1

Kavali-Udupi Transect

The first such study was conducted along Kavali (east coast) to Udupi (west coast) in southern part of the Indian Peninsular shield (Fig. 2.1) way back in 1972, through a collaboration between CSIR- NGRI (Hyderabad, India) and the Ukranian Academy of Sciences in the former USSR. It was completed in three successive seasons. The purpose of this study was to delineate deep seated crustal/structural configuration using wide angle reflections/refractions data. This profile traversed through all the major geological units of southern India, like Eastern Ghats Belt, Nellore Schist Belt, Cuddapah Basin, Closepet granite and Chitradurga Schist Belt, thereby covering the

2.4 Western Dharwar Craton

45

entire stretch of WDC and EDC. Initial seismic results (Kaila et al. 1979), identified number of crustal and sub crustal reflectors in the seismic section up to a depth of ~70 km. They also identified number of deep seated faults and variations in Moho depths from one segment to the another. Subsequently, acquired data along this profile have been reprocessed by many workers like Reddy et al. (2000, 2004), Sarkar et al. (2001), Mall et al. (2008, 2012) and Chandrakala et al. (2010, 2013, 2015, 2017). Since the earlier studies were based on the analog data, Mall et al. (2012) utilized the digitized data to study the seismic structure of WDC. They modeled refraction and wide-angle reflection data to obtain 2-D (isotropic) P-wave crustal velocity structure, using travel time inversion code, described in detail by Zelt and Smith (1992) and Zelt (1999). Detailed procedure of data acquisition, data reduction and interpretation, are given in Kaila et al. (1979), Mall et al. (2012) and Chandrakala et al. (2010, 2013, 2015). The derived crustal seismic velocity model for WDC is shown in Fig. 2.2. It revealed a six layered crust having velocities 5.90, 6.20, 6.40, 6.60, 6.90 and 7.30 km/s above the Moho, recorded at 45 km depth in central part, with a slight upwarp towards the east, where it reaches 42 km. The bottom- most high velocity (7.3 km/s) layer above the Moho, which is about 10 km thick, would corresponds to emplaced solidified magma, below which velocity jumps from 7.3 to 8.35 km/s. Such a high Pn velocity (8.35 km/s) at a depth of only 45 km is rather unusual for any Archean crustal terrain. An unusually occurring low velocity Zone (LVZ), characterized by the velocity of 7.40 km/s, is also obtained in the upper mantle between depths 47 and 59 km. Such a shallow subcrustal lithospheric LVZ is rarely found in stable Archean cratons. Mall et al. (2012) ascribed the LVZ to metasomatic alterations, caused by rising mantle fluids from the upper mantle (Pandey et al. 2016). The crustal section derived by Mall et al. (2012) obtained a thicker crust (~45 km) in WDC than ~41 km derived earlier (Kaila et al. 1979; Reddy et al. 2000; Sarkar et al. 2001).

Fig. 2.2 Crustal seismic section across the western Dharwar craton (WDC), covering western part of Kavali-Udupi traverse, derived from deep seismic sounding studies. Layer velocities are in km/s. SP refers to shot points. Modified after Fig. 5, Mall et al. (2012)

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Fig. 2.3 Crustal seismic structure along Chikmagalur-Perur geotransect (Fig. 2.1), cutting across western and eastern Dharwar cratonic terrain based on the studies of Vijaya Rao et al. (2015) and Mandal et al. (2018). WDC: Western Dharwar craton, EDC: Eastern Dharwar craton, BGB: Bababudan Greenstone Belt, OG: Older gneisses, CGB: Chitradurga Green stone Belt, YG: Younger gneisses, CG: Closepet granite, UC: Upper crust, MC: Middle crust, LC: Lower crust, UPC: Underplated magmatic crust, CT: Chitradurga Thrust. Modified after Fig. 7, Mandal et al. (2018)

2.4.1.2

Perur-Chikmagalur Profile

This 200 km long profile runs from Chikmangalur to Perur (Fig. 2.1). It covered entire Archean segments of western part and some parts of adjacent EDC, that includes greenstone schist belts, Chitradurga shear zone, Closepet granites and peninsular gneisses. Travel time inversion and amplitude modeling of seismic data, acquired from seismic refraction/wide-angle reflection experiments, delineated a five-layer crust (Fig. 2.3), which can be clubbed into four zones, characterised by average Vp (i) 5.85–6.0 km/s, corresponding to 3–4 km thick exposed crystalline basement/green stone cover, (ii) 6.3 km/s to 9–10 km thick middle crust, (iii) 6.5–6.7 km/s, representing 20 km thick lower crust, and (iv) 7.0 km/s above the Moho conforming to 10 km thick high velocity magmatic crust (Vijaya Rao et al. 2015; Mandal et al. 2018). Below this region, Moho is 43 km deep, characterised by upper mantle velocity of 8.1 km/s. Such a thick magmatic infusion below WDC could be related to Meso-Neoarchean magmatism, Mesoproterozoic dyke swarms activity, late Cretaceous dyke intrusions, and Madagascar breakup due to marion plume (Anil Kumar et al. 2001; Jayananda et al. 2008), that led to widespread reworking and modification of the lower crust. In comparison to WDC, crust is only 37–38 km thick near Perur located in EDC. Moho between the two adjacent cratonic segment appears gradual, rather than abrupt.

2.4.1.3

Receiver Function Studies

There have been a number of receiver function or similar kind of studies over the Dharwar craton (Rai et al. 2003; Gupta and Rai 2005; Kumar et al. 2007, 2013;

2.4 Western Dharwar Craton

47

Fig. 2.4 Shear wave velocity structure across Dharwar craton based on studies by Saikia et al. (2017). Velocities are in km/s. WDC: Western Dharwar craton, EDC: Eastern Dharwar craton, CDF; location of Chitradurga fault separating WDC and EDC terrains

Jagadeesh and Rai 2008, Kiselev et al. 2008, Julia et al. 2009; Ravi Kumar et al. 2013; Saikia et al. 2017), but many of the results obtained by such studies were of contrasting nature. Similar is the case while deducing the lithospheric thicknesses. For example, Julia et al. (2009) reported a thick felsic upper crust of about 12.5– 17.5 km (having Vs < 3.60 km/s) and deep Moho at 45–52.5 km below WDC, which is in contrast to recent works by Saikia et al (2017), who studied shear wave velocity structure along 660 km long profile (Fig. 2.4), through the joint inversion of receiver functions and Rayleigh wave group velocity. Their profile, consisted of 38 broadband seismic stations that covered both WDC and EDC. It revealed Moho around 40 km depth beneath WDC and 35 km below EDC; similar to that obtained by deep crustal seismic studies. Similarly, Saikia et al. (2017) found very thin upper crust, only 3– 4 km (in comparison to 12.5–17.5 km by Julia et al. 2009) and a thick underplated magma layer (11 ± 3 km) below WDC. Pandey et al. (2013) too analysed data from 32 broadband seismic stations and found that the Indian crust is of intermediate to mafic in composition and only a couple of kilometers of the initially formed upper granitic-gneissic crust now remains in place (Table 2.1). Saikia et al. (2017) found similar results. They also found stacking of about 16 km thick magma layer on either side of the Moho, due to sub-crustal erosion.

2.4.2 Gravity Field Regional gravity field is possibly the best geophysical tool to gainfully understand the nature of concealed deep-seated geologic features, as well as the subsurface mass/density distributions. It helps to decipher whether the gravity field is caused due to sources lying in the crust or beyond in the mantle. Observed Bouguer gravity field over the cratonic part of the Dharwar craton (Fig. 2.5) varys from almost −60 to

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Table 2.1 Inferred compositional structure beneath various segments of the south Indian shield. Thicknesses of different crustal layers are in kilometres WDC Upper granitic-gneissic crust (felsic)

EDC

DVP

SGT 1

3

4

2

Middle crust (intermediate)

21

20

4

9

Original lower crust (mafic)

19

8

28

25

Underplated Magmatic lower crust Moho depth Mantle magma layer Depth to normal ultramafic mantle

5

3

4

9

48

35

38

44

4

1

2

10

52

36

40

54

Modified from Table 2, Pandey et al. (2013) WDC Western Dharwar craton; EDC Eastern Dharwar craton; DVP Deccan volcanic covered areas of Dharwar craton; SGT Southern granulite terrain

−120 mGal. An examination of this map indicates that the regional gravity trend over WDC is NW–SE. WDC is conspicuously associated with negative gravity anomaly as well as several relatively higher anomalies, superimposed over a high order broad regional gravity low (Agrawal et al. 1992), probably reflecting relatively thicker crust (44–45 km) and relatively thicker and colder lithosphere below WDC, compared to other cratons, specially those located in the northern part of the peninsular shield. The gravity field which is little higher near the western margin at around − 60 mGal, drops down sharply to about −120 mGal (anomaly L1) over the central parts of WDC containing supracrustals, volcano-sedimentary sequences of the greenstone/schist belts and Peninsular gneiss, then increases to about −65 mGal over the Shimoga and Chitradurga Schist Belts (Singh et al. 2004). Short wavelength gravity lows can be ascribed to presence of granite outcrops. Similarly, the N–S trending linear positive Bouguer anomalies paralleling the coast, appears to be a regional feature extending far beyond the Dharwar craton and covering the entire west coast of India,. This could be interpreted in terms of highdensity rocks emplaced at base of the crust (Singh and Meissner 1995; Singh 1998, 1999; Singh and Mall 1998). Alternatively, high-gradient contours on the continental coastal region, may represent uplifted and rifted part of the Dharwar cratonic block, related to India –Madagascar breakup at around 90 Ma. Crustal thinning and asthenospheric upwarping, could be another reason (Qureshy 1981).

2.4.3 MT Studies Gokarn et al. (2004), Abdul Azeez et al. (2015) and Malleswari et al. (2019) carried out detailed MT investigations along three nearly east-west traverses over western part of the Dharwar craton (Fig. 2.1). While the southern and central segment was

2.4 Western Dharwar Craton

49

Fig. 2.5 Bouguer gravity anomaly map (in mGal) of Dharwar craton (scale: 1:5000,000) (NGRI, 1975). Possible extension of gravity trends from SGT to EDC, as shown by dotted line, can also be clearly seen. H1–H6 are locations of gravity highs and L1–L3, gravity lows. WDC: Western Dharwar craton, EDC: Eastern Dharwar craton, SGT: Sothern Granulite Terrain. Modified after Fig. 8, Agrawal and Pandey (2004)

covered by Abdul Azeez et al. (2015) and Malleswari et al. (2019), Gokarn et al. (2004) studied northern segment of the WDC along a 400 km long ENE–WSW trending Goa–Dharwar–Jadcherla profile (profile 1 in Fig. 2.1) that covered the granite– greenstone terrain. This MT investigation collected data at 50 stations and revealed presence of (i) 3–5 km thick high resistivity top layer, throughout the profile, and (ii) a prominent suture zone along the Chitradurga–Gadag schist belt, formed due to thrusting of WDC below EDC. They also delineated a low resistivity zone at 40 km depth beneath WDC, which may be a regional feature. It may be extending below Deccan traps region as well. Presence of high conductivity is ascribed to the

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hydrothermal fluids trapped in the mantle during the passage of the Indian Plate over the Reunion hotspot. Low-resistive asthenosphere is detected at 80–100 km depth beneath WDC. Similarly, Abdul Azeez et al. (2015) reported results along a ~250 km east–west MT traverse, covering WDC and adjoining Coorg block (Profile 3 in Fig. 2.1). This study, which was based on 19 long-period MT measurements spaced at around 15 km interval, reported presence of several discrete low-resistivity (around 30–100 m) zones (labeled as crustal conductors C1 to C6 in Fig. 2.6), within the highly-resistive crustal formation. Out of these, C3 to C6 correspond to greenstoneschist belts, while C1 and C2 are located in the northern fringe of the Coorg block over a shear zone that demarcates Coorg block from the WDC terrain. These mid to lower crustal conductive zones are attributed mainly to presence of fluid in sheared zones. Srikantappa et al. (1994) have earlier reported high-grade CO2 -rich fluid metamorphism (7–8.6 kbar, 720–760 °C) for the Coorg granulites, in comparison to amphibolites facies gneissic formations further north in WDC. Besides these crustal conductors, a near vertical, low-resistivity structure (MC1 in Fig. 2.6) is located in the upper mantle beneath the transitional suture zone between the Coorg block (located in the southernmost part of WDC) and the WDC main terrain, which is formed due to collison between these two blocks during early Paleoproterozoic. This prominent conductor is associated with highly negative Bouguer gravity anomaly of about −120 mGal. The substantial gravity anomaly drop (from −50 to −120 mGal) over this zone, is related to the presence of a collision zone underneath. This conductor is underlain by another anomalous upper mantle conductive zone (MC2) below the western Dharwar cratonic nucleous, which may represent subsequent modification and destruction of the cratonic root. Consequently, reported lithosphere is only around 125 km thick below the Coorg block (Fig. 2.6), compared to about 190 km in eastern segment of the WDC. Heat flow studies, as discussed later, would support this conjecture, which may mean that at a few locations in WDC, some parts of cratonic keel may still be well preserved.

Fig. 2.6 Resistivity section from Coorg to Chitradurga schist belt across WDC based on Abdul Azeez et al. (2015). The electrical LAB (e-LAB) from MT studies, is also marked

2.4 Western Dharwar Craton

51

Malleswari et al. (2019) also covered this craton from east to west. They reported results of broad-band and long period MT measurements at 63 locations along a 500 km long profile, that originated from Chikmagalur (west coast) to Kavali on the east coast. Like, Abdul Azeez et al. (2015), this study too reported presence of several moderately conductive features in the crust and underlying uppermost mantle, along the entire traverse. They also reported presence of two conductive features, one occurring at the depths of 100–150 km below western margin of the WDC. It was supported by Kusham et al. (2018), who found a similar conductive zone at depths of 120–150 km below WDC. The other conductor, nearly a vertical conductive feature in the upper mantle, was found located below the Chitradurga Shear Zone that separates the WDC from EDC. LAB was detected at the depth of about 200 km below WDC.

2.4.4 Heat Flow and Lithosphere Entire Dharwar craton is very well studied with respect to heat flow. Altogether, there are 73 measurements available over this terrain, which vary from a low of 23 to a high of 75 mW/m2 (Fig. 2.7). Out of these, 5 heat flow values come from WDC, which are consistently low and vary from 29 to 39 mW/m2 with a mean of 32 ± 3.6 mW/m2 . Measured heat flow over this region can be considered low, reflecting the stable nature of the underlying lithosphere beneath Archean Dharwar nucleii. These low values are similar to those recorded in comparable global terrains like Yilgarn block of western Australia, Baltic shield, Canadian shield and Kaapval craton (Sass and Lachenbruch 1979; Mareschal et al. 1989; Nyblade et al. 1990; Balling 1995). Based on crustal heat production model (Table 2.2), temperature-depth variation is estimated for the Gadag region (Fig. 2.8) using the method described in detail in Chap. 9. Figure 2.8 reveals lithosphere to be about 185 km below WDC, similar to that reported earlier by Negi et al. (1986, 1987). As it would appear, WDC is underlain by thickest lithosphere among all other Indian terrains, but still it is almost half of 300–350 km recorded in other global shields (Chapman and Pollack 1977; Polet and Anderson 1995; Pandey and Agrawal 1999; Pandey 2016).

2.5 Eastern Dharwar Craton 2.5.1 Crustal Seismic Structure Geotectonically, this part of the Dharwar craton is more complex than the WDC, containing several geodynamically important features like Closepet granite in the west to mantle-plume-infected Cuddapah basin in the middle, Nellore Schist Belt (NSB), Ongole domain of the Eastern Ghats Belt (EGB) and possibly Gondwana and Proterozoic basins further east. Out of these geotectonic elements, the Paleoproterozoic

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Fig. 2.7 Heat flow (in mW/m2 ) distribution in the south Indian shield. Data is taken from Verma et al. (1969), Rao (1970), Gupta and Rao (1970), Verma and Gupta (1975), Rao et al. (1976), Gupta et al. (1987, 1991), Sharma et al. (1991), Roy and Rao (2000), Ray et al. (2003), Roy et al. (2003, 2007, 2008). WDC: Western Dharwar craton, EDC: Eestern Dharwar craton, SGT: Southern Granulite Terrain, EGMB: Eastern Ghat Mobile Belt, KKB: Kerala Khondalite Belt, AKSZ: Achankovil Shear zone, PCSZ: Palghat-Cauvery Shear Zone. Modified after Fig. 3, Agrawal and Pandey (2004)

aged intra-cratonic Cuddapah basin, also referred as Purana basins (Holland 1906; Radhakrishnan 1987), is most prominent as mentioned earlier, forming the largest basin in southern India. It also hosts minerals of economic importance. Two deep seismic profiles, (i) Alampur–Koniki and (ii) Kavali-Udipi cuts across this terrain (Fig. 2.9).

2.5.1.1

Kavali-Udupi Transect

After the classical work by Kaila et al. (1979), a large number of studies followed (Mall et al. 2008; Chandrakala et al. 2010, 2013, 2015, 2017; Pandey et al. 2018), where the seismic data acquired along the Kavali-Udupi traverse was re-interpreted

2.5 Eastern Dharwar Craton

53

Table 2.2 Crustal heat production models and adopted geothermal parameters for temperaturedepth estimation beneath various geotectonic segments of Western and Eastern Dharwar cratons Depth range (km)

Rock type

Heat production (µW/m3 )

Thermal conductivity (W/m°C)

Gadag (Western Dharwar craton)

Surface heat flow: 29.0 mW/m2

0–3.0

Chitradurga supracrustals

0.4

3.0

3.0–12.0

Amphibolite-granulite

0.4

2.88

12.0–23.0

Granulite

0.2

2.5

23.0–32.0

Mafic granulite

0.16

2.5

32.0–42.0

Magmatic crust

0.02

2.6

>42.0

Ultramafic mantle

0.01

3.0

Data source Gupta et al. (1991), Liu and Zoback (1997), Roy and Rao (1999), Ray et al. (2003), Vijaya Rao et al. (2015), Mandal et al. (2018) Kolar (Eastern Dharwar craton)

Surface heat flow: 39.0 mW/m2

0–4.0

Granite-greenstone

1.5

4.0–14.0

Granulite

0.2

2.5

14.0–32.0

Mafic granulites

0.16

2.50

32.0–34.0

Magmatic crust

0.02

2.6

>34.0

Ultramafic mantle

0.01

3.0

3.0

Data source Gupta and Rao (1970), Rao et al. (1976), Liu and zoback (1997), Ray et al. (2003) Agnigundala (Cuddapah basin, EDC)

Surface heat flow: 75.0 mW/m2

0–1.0

Argillites and Phyllites

2.61

3.0

1.0–17.0

Amphibolite-granulite

0.78

2.88

17.0–27.0

Granulite

0.20

2.5

27.0–44.0

Magmatic crust

0.02

2.6

>44.0

Ultramafic mantle

0.01

3.0

Data source Rao et al. (1976), Liu and Zoback (1997), Roy and Rao (1999), Ray et al. (2003), Pandey et al. (2017), K. Chandrakala (personal communication) Dharmavaram (Eastern Dharwar craton)

Surface heat flow: 40.0 mW/m2

0–3.0

Peninsular gneisses

1.70

3.0

3.0–11.0

Amphibolite-granulite

0.78

2.88

11.0–23.0

Granulite

0.2

2.5

23.0–30.0

Mafic granulite

0.16

2.5

30.0–38.0

Magmatic crust

0.02

2.6

>38.0

Ultramafic mantle

0.01

3.0

Data source Roy and Rao (1999, 2000), Liu and Zoback (1997), Ray et al. (2003), Vijaya Rao et al. (2015), Pandey et al. (2017) (continued)

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Table 2.2 (continued) Depth range (km)

Rock type

Tummalapalli (Cuddapah basin, EDC)

Heat production (µW/m3 )

Thermal conductivity (W/m°C)

Surface heat flow: 51.0 mW/m2

0–2.5

Granite-gneiss

2.20

3.62

2.5–15.0

Amphibolite-granulite

0.78

2.88

15.0–35.0

Mafic granulite

0.16

2.5

> 35.0

Ultramafic mantle

0.01

3.0

Data source Roy and Rao (1999, 2000), Ray et al. (2003), Pandey et al. (2017) Pavagada (Closepet granites, EDC)

Surface heat flow: 39.0 mW/m2

0–3.0

Granites

1.8

3.22

3.0–12.0

Metasomatised granulite

0.35

2.5

12.0–23.0

Granulite

0.2

2.5

23.0–30.0

Mafic granulite

0.16

2.5

30.0–38.0

Magmatic crust

0.02

2.60

>38.0

Ultramafic mantle

0.01

3.0

Data source Liu and Zoback (1997), Ray et al. (2003), Roy et al. (2008), Vijaya Rao et al. (2015)

using modern inversion tools. Observed results were integrated with other geophysical findings in order to decipher subsurface deep crustal structure. These studies provided a totally new picture of crustal seismic configuration across the Cuddapah basin and the regions located east and west of it. Shallow Seismic Structure (i) Sedimentary Thickness in Cuddapah basin There has always been a controversy about the thickness of sediments in the basin, as discussed in detail by Chandrakala et al. (2013). Crustal seismic model (Fig. 2.10) indicates presence of a two layer sedimentary column, having velocities 5.20 and 5.60 km/s, resting over the granite-gneissic basement with velocity 5.90–6.00 km/s. The first layer at the top represents exposed upper Cuddapah sediments, while the second layer at the bottom, corresponds to lower Cuddapah formation, both of them together having a cumulative thickness of about 4 km in the deepest part of the basin that lies below the Nallamalai Fold Belt. (ii) Vinjamuru Gondwana and East coast Proterozoic Basins The shallow seismic section across the NSB (Fig. 2.11), reveals presence of about 250 m thick Gondwana sediments, with distinct velocity of 4.20 km/s below the shot point 40, which was inferred from direct wave arrival times, thus unambiguously representing velocity of the exposed outcrop below the shot point region. In this zone, the resolution of seismic data is quite high; the data have been acquired using

2.5 Eastern Dharwar Craton

55

Fig. 2.8 Estimated temperature-depth distribution beneath chosen geotectonic segments of Western and Eastern Dharwar cratons, south Indian shield, based on crustal heat production models as given in Table 2.2. Locational details are given in Table 9.1, Chap. 9

a small geophone separation of only 100 m. This layer stretches for about 40 km around Vinjamuru region of the NSB and pinches out on its either side. This new find was termed as ‘Vinjamuru Gondwanas’ by Chandrakala et al. (2017). Their study also detected presence of 5.3 and 5.5 km/s velocity sediments between NSB and the eastern coast, which correlate very well with the upper and lower Cuddapah sediments (Chandrakala et al. 2015), suggesting possible occurrence of another Proterozoic sedimentary terrain in the region, besides the Cuddapah basin. This Proterozoic basin is named as East Coast Sedimentary Basin (ECSB in Fig. 2.11). It is separated from the Cuddapah basin by an exhumed horst-like feature which underlies NSB.

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Fig. 2.9 Detailed geological map of the Cuddapah basin and adjoining regions, showing location of DSS profiles with shot points. P–K refers to Parnapalle–Kavali segment of Kavali-Udipi DSS profile, and A–K to Alampur–Koniki profile. KOC and KOM respectively indicate the locations of Kandra Ophiolite Complex and Kanigiri Ophiolite Melange. After Fig. 1, Chandrakala et al. (2015)

Fig. 2.10 Shallow crustal seismic section along Kavali-Parnapalle section of EDC. Numbers in the model represent average P-wave velocities in km/s. Dashed lines represent faults. UC: Upper Cuddapah sediments; LC: Lower Cuddapah sediments: NKSB: Nellore Khammam Schist Belt; EGMB: Eastern Ghat Mobile Belt. Modified after Fig. 10, Chandrakala et al. (2013)

2.5 Eastern Dharwar Craton

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Fig. 2.11 Shallow crustal velocity—depth model below Nellore Schist Belt (NSB) and adjacent areas. CB and ECSB refers to Cuddapah basin and East Coast Sedimentary basin respectively. Extent of Vinjamuru Gondwanas, as characterized by velocity of 4.2 km/s, is also shown. The number indicates the velocity of the corresponding crustal layers. Dashed lines indicate faults. After Fig. 7, Chandrakala et al. (2017)

5.30 km/s velocity layer, which is exposed to the east of this belt, where it may be ~400–800 m thick, is either very thin or almost absent beneath NSB. Deep Crustal Seismic Structure An improved version of the recently obtained crustal seismic section based on the studies of Mall et al. (2008), Chandrakala et al. (2013, 2015, 2017) and Pandey et al. (2018) is shown in Fig. 2.12. This figure reveals presence of a four layer crust above the Moho (8.1–8.3 km/s) below Cuddapah basin. Crystalline basement, exposed outside the basin, is associated with 5.9 km/s velocity, which is found at a depth of

Fig. 2.12 Deep crustal seismic section along Kavali-Parnapalle deep seismic sounding profile, covering Cuddapah basin and adjoining regions. Inverted triangle shows the shot point locations. NSB: Nellore Schist Belt. F1–F6 represent seismically derived faults. Velocities in different crustal layers are given in km/s. Modified after Fig. 8, Chandrakala et al. (2015)

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2 Dharwar Craton

about 1 km below NSB, 3 km below ECSB and at extremely shallow depths at both the ends of Cuddapah basin. This layer up-dips at an angle of ∼7° below NSB and ECSB, and is underlain further by a mid-crustal layer (6.30–6.50 km/s), which dips steeper at about 12º. This layer is followed by mafic lower crust (6.5–6.6 km/s), which is underlain further by an unprecedently thick (15–25 km) accreted high velocity (7.0– 7.4 km/s) magmatic crust above the Moho that may be differentiated. The Moho depth along the profile varies from 40 to 45 km. On the far western side near Parnapalle, a conspicuous signature of an upwelled thermal plume is seen, which coincides with the exposures of 1.1 Ga Kimberlitic magmatism (Anil Kumar et al. 1993). Six major faults are also detected along the profile. They typically coincide almost one to one with major geotectonic boundaries. For example, prominent faults F1 and F2 demarcate the western and eastern boundary of the Cuddapah basin while, faults F2 and F6 coincides with the boundaries of NSB. The exposed Chitravati, Kurnool and Nallamalai formations of the Cuddapah basin are seen bounded by faults F1 and F4, F4 and F5, and F5 and F3 respectively. All the non-magmatic crustal layers in the upper parts, show up-dipping trend down to about 20 km below NSB and ECSB. In contrast, the magmatic layer at the bottom as well as the Moho tend to show down-dip trend.

2.5.1.2

Perur-Chikmagalur Transect

This profile covers Chitradurga boundary fault-Perur segment of EDC. The crustal structure below this segment is almost similar to that found below WDC. Upper crust (6.00 km/s) is about 3 km thick, followed by 12 km thick middle crust (6.25 km/s). Lower crust (6.5 km/s) is about 18 km thick and extends down to a depth of 30 km. This layer is underlain by about 8 km thick magmatic crust with velocity 7.0 km/s. However, the Moho is shallow at 38 km depth, compared to 42–43 km in WDC. The crustal velocity structure (Fig. 2.3) across the EDC–WDC boundary zone, is consistent with the Neoarchean plate convergence which would support (i) subduction of the EDC beneath the WDC, (ii) presence of a collision zone between the two cratonic blocks below the Closepet granites (CG in Fig. 2.3), and (iii) shallowing of Moho below Chitradurga fault shear zone (Fig. 2.3). This profile also revealed distinct differences in seismic properties on either side of Chitradurga fault shear zone, the WDC being associated with a simple structure, compared to EDC which reflects complex reflectivity pattern (Mandal et al. 2018). Besides, an easterly dipping a 10–12 km thick Chitradurga thrust zone (CT) was also mapped below this segment that begins from the Chitradurga Fault zone and ends below the Closepet granites and extends down to Moho (Fig. 2.3). This feature compares well with the Jahazpur thrust zone (Mandal et al. 2014; Vijaya Rao et al. 2000) found below the Hindoli and Mangalwar group of rocks in the Mesoarchean Mewar Terrain of the Aravalli craton, as discussed in Chap. 4. This profile further indicated presence of an west-dipping reflection fabric extending from 34 km to 43 km depth below EDC, which may indicate paleo-subduction, as mentioned above.

2.5 Eastern Dharwar Craton

59

Fig. 2.13 Crustal depth section along a part of the Alampur-Koniki DSS profile (A–K in Fig. 2.9), covering northern part of the Cuddapah basin. NFB: Nallamalai Fold Belt; CB: Cuddapah basin; NSB: Nellore Schist Belt; EGB: Eastern Ghats Belt. Numbers inside the different segments refer to crustal velocities in km/s. Modified after Fig. 3, Kaila and Tewari (1985)

2.5.1.3

Alampur -Koniki Profile

Alampur-Koniki is another DSS profile (A–K in Fig. 2.9), which traversed through northern part of the Cuddapah basin, and seems to contain similar faults (Kaila and Tewari 1985) like those found along Kavali-Parnapalle profile. One of them is located near SP 140 A (Fig. 2.13), but orientation of this fault differs from that reported by Chandrakala et al. (2015); it appears to be a normal fault. Further, upthrusting signatures east of SP 140 can also be clearly seen, where lower crustal layer (6.69 km/s) seems to have moved upwarped by as much as 13 km below NSB. Here, the angle of upthrusting is much steeper than that observed along the Parnapalle-Kavali profile (Fig. 2.11). Upthrusting of the crustal layer, lying east of Cuddapah basin, appears well correlated with the gravity and other geophysical fields (Pandey et al. 2018), as discussed in subsequent sections.

2.5.1.4

Receiver Function Studies

Julia et al. (2009) studied receiver function using 8 broadband seismic stations located in EDC and Cuddapah basin. It was reported that upper crust is 10–15 km thick and the Moho depth, 32.5–35 km. On the other hand, Saikia et al. (2017), based on a much larger data set, reported similar (35 km) Moho depth, but found extremely thin upper crust of only 3–4 km, almost identical to the findings of Pandey et al. (2013) in EDC. Saikia et al. (2017) also reported thick magma underplating of about 11 ± 3 km below the Cuddapah basin. A much thicker underplated magma layer was also contemplated by deep seismic profiles (Fig. 2.12). Based on data acquired from the

60

2 Dharwar Craton

CUD broadband seismic station (14°29 , 78°46 ) that falls over the western part of Cuddapah basin, Kumar et al. (2007) reported shallow Lithosphere- asthenosphere boundary at a depth of only about 80 km.

2.5.2 MT Studies Naganjaneyulu and Harinarayana (2004) carried out MT study between west of Anantpur to Kavali along the Kavali-Udupi DSS transect and estimated conductivity 300–500 m in the upper 4 km strata, lying over NSB, East Coast Proterozoic and Vinjamuru Gondwana basins (Fig. 2.14). They delineated an anomalous conductive feature, associated with very high conductivity (resistivity less than 100 m), below SW part of the Cuddapah basin, which is also characterised by a high gravity anomaly (H1 in Fig. 2.5; C in Fig. 2.15b). This anomaly is ascribed to a paleo mantle thermal plume (Mall et al. 2008). It has a width of about 100 km, extends down to at least 50 km depth and may be of basic/ultra basic in composition. The MT studies carried out in the northeast part of the Cuddapah basin that covered Neoproterozoic succession of the Palnad sub-basin as well as the Nallamalai Fold Belt, indicated presence of a high conductivity (resistivity 40–300 m) in the top 350–2000 m thick layer which contain highly disturbed sediments of the Palnad sub-basin (Konda et al. 2015). As mentioned in earlier sections, the MT profiles carried out by Gokarn et al. (2004) and Malleswari et al. (2019) also covered substantial part of the eastern Dharwar craton. Gokarn et al. (2004) revealed presence of a 3–5 km thick high resistivity top layer beneath the EDC, which is underlain by a zone of low resistivity at a Fig. 2.14 2-D Geoelectric depth section along DSS profile below Cuddapah basin and adjoining areas. EDC: Eastern Dharwar craton, CB: Cuddapah basin, EGB: Eastern Ghats Belt. Numbers within the figure refer to velocities in Km/s. Modified after Fig. 5, Naganjaneyulu and Harinaryana (2004)

2.5 Eastern Dharwar Craton

61

Fig. 2.15 a Bouguer gravity contour image map of the Cuddapah basin and surroundings, based on new gravity map series (GSI-NGRI 2006). Proposed paleo boundary of the basin, is indicated by yellow dashed lines (After Fig. 5, Pandey et al. 2018). b Residual gravity contour image derived using Finite Element Method. A–E represents the location of the gravity anomalies discussed in the text. Red dashed lines indicates paleo boundary of the Cuddapah basin (Modified after Fig. 8, Chandrakala et al. 2017)

62

2 Dharwar Craton

depth of around 40 km, and related it to presence of hydrothermal fluids trapped in the mantle. They also delineated low-resistive asthenosphere at a depth of about 160 km beneath EDC. Malleswari et al. (2019), on the other hand, covered entire EDC and EGB, mostly along the erstwhile Kavali-Udupi DSS profile (similar to Naganjaneyulu and Harinarayana 2004; Fig. 2.14) and obtained quite a few moderately conductive features in the crust and underlying uppermost mantle. Importantly however, they delineated a west-verging moderately conductive feature beneath the EGB and interpreted it as the remnant of the Proterozoic collision zone between India and East Antarctica. Lithosphere beneath EDC was found to be thicker (more than 200 km), but beneath EGB it is found to be only 120 km thick.

2.5.3 Regional Gravity Field Numerous studies (e.g. Kaila and Bhatia 1981; Singh and Mishra 2002; Singh et al. 2004; Kumar et al. 2004) have been made to model the gravity field across EDC. These studies were biased on seismic constraints provided by Kaila et al. (1979). Over this terrain, gravity contours trend NW–SE, compared to NE–SW over the granulite terrain down south (Fig. 2.5). Western part of the EDC, which is characterised by the gigantic presence of medium-grade Closepet granites, is associated with relatively positive anomalies (−65 to −70 mGal, H2 in Fig. 2.5), and then, it decreases gradually to −90 mGal over the EDC further east till the southwestern part of the Cuddupah basin, where it again increases to −55 mGal (H1 in Fig. 2.5).

2.5.3.1

Cuddapah Basin and Surrounding Region

Vastly improved gravity map series of India have now become available (GSI-NGRI 2006). This map was re-digitized and re-contoured in order to prepare the Bouguer gravity map as shown in Fig. 2.15a. The residual gravity map prepared following a numerical technique, based on the finite element approach (Mallick and Sharma 1999; Mallick et al. 2012) is reproduced in Fig. 2.15b. These maps have brought out some major geotectonic features quite prominently. The entire Cuddapah basin, except the south western part, is characterised by a series of conspicuous high order gravity lows up to −110 mGal (Fig. 2.15a), which conform well with the presence of thick Proterozoic sediments. The anomalies are particularly negative over the Nallamalai Fold Belt, situated in the eastern half of the basin, where the Bouguer gravity anomaly reaches −110 mGal and residual anomaly, −50 mGal (anomaly A, Fig. 2.15b). In this segment, sediments are thickest at around 4 km (Chandrakala et al. 2013). Gravity patterns would suggest that the Nallamalai formations may extend further into the western part of the Cuddapah basin (anomaly B, Fig. 2.15b). The signatures of this anomaly are much clearer in Free-air gravity (Pandey et al. 2018). The plume infected western part of the basin, which is dominated by volcanosedimentary succession, is associated with much lesser sediment thickness. It is

2.5 Eastern Dharwar Craton

63

characterised by a high order circular positive Bouguer gravity that reaches −55 mGal between Tadipatri and Gandikota (Fig. 2.15a), with residual gravity anomalies reaching to about 30 mGal (anomaly C, Fig. 2.15b). This conspicuous feature has earlier been attributed to various factors like thickening of basaltic layer, crustal upwarp, mafic dyke/sill intrusion (Krishna Brahmam et al. 1986), impact structure (Grant 1983; Krishna Brahmam and Dutt 1985). A mantle plume model (Mall et al. 2008), however, conforms well with the presence of exposed high density basic and ultra basic rocks (sills and intrusives), occurring in an arcuate pattern, paralleling the western margin of Cuddapah basin (Saha and Tripathy 2012). Nellore Schist Belt and East Coast Poterozoic Sedimentary Basin Bouger gravity anomalies jump sharply to −15 mGal over Nellore Schist Belt (NSB), from a low of almost −110 mGal over Nallamalai formation (Fig. 2.15a). The computed residual gravity anomaly is also quite high over the NSB (about 15 mGal, anomaly D, Fig. 2.15b). The low to high gravity transition coincides with the Velikonda thrust zone (Venkatakrishnan and Dotiwala 1987; Saha 2002; Saha et al. 2010), demarcating the Cuddapah basin and NSB; the latter region being associated with high gravity anomalies conforming seismically mapped horst like feature, bounded by deep faults F2 and F6 (Fig. 2.12). A 10–15 km vertical displacement in lower crustal strata, as well as in the underplated magma layer, is observed along the malor fault F2. The residual gravity field is −25 mGal (anomaly E, Fig. 2.15b) over East Coast Sedimentary Basin (ECSB), situated east of the NSB. Quite likely, sediments of this basin may have been metamorphosed by magmatic activity. In essence, gravity field over this region can be attributed to continent-continent collision and suturing. Gravity Modelling Along the Seismic Profile Utilising new seismic and gravity constraints, two-dimensional gravity modelling is attempted along the seismic profile using GM-SYS software (Geosoft Oasis Montaj 2004). In the modelling, assigned densities for various rock types are based on established velocity-density relationships (Barton 1986; Christensen and Mooney 1995). Details of the gravity data and modelling procedure are given in Pandey et al. (2018). The results are shown in Fig. 2.16, where structural and density variations with depth are clearly seen.

2.5.4 Magnetic Studies The CSIR-NGRI conducted aeromagnetic survey over parts of the EDC and Cuddapah basin during 1980–82. In this survey, two E–W profiles were flown at a height of 150 m and 1000 m (NGRI 1985; Babu Rao et al. 1987), close to the existing seismic profiles (Kaila et al. 1979; Kaila and Tewari 1985) in the northern, as well as southern part of the Cuddapah basin and adjacent region lying east of it. Total intensity magnetic anomaly along the Maruru-Kavali profile, with a flight height of 1000 m

64

2 Dharwar Craton

Fig. 2.16 Derived crustal density (in g/cm3 ) model across Cuddapah basin and adjoining areas on its east. CB: Cuddapah basin; NFB: Nallamalai Fold Belt; NSB: Nellore Schist Belt; ECSB: East Coast Sedimentary basin. Black dashed lines represent seismically derived faults (After Fig. 8, Pandey et al. 2018)

and 150 m is shown in Fig. 2.17a and b respectively. To correlate the variations of magnetic anomalies at three different levels, the ground magnetic data acquired along this profile by Kailasam (1976) is included. Observed Bouguer gravity and computed residual gravity anomalies are also shown. The magnetic field correlated fairly well with the geological features, demarcated by seismic and gravity investigations. The prominent ground magnetic high anomaly (−300 to + 400 nT; Fig. 2.17c), observed in western part of the Cuddapah basin, may be correlated with wide-spread occurrences of kimberlite-clan rocks. An ultra basic intrusive body with magnetic susceptibility of 25 × 10−3 appears to be present at a depth of 1.9 km, which has a width of 8.6 km and dips at 16° towards east (Pandey et al. 2018). The seismically derived major faults F1 and F2 which bound this basin on either side, and faults F2 and F6 which bound the Nellore Schist Belt (NSB), are well reflected in magnetics (Fig. 2.17a–c). The Faults F2 and F6 (Fig. 2.12) are characterised by strong fault-type magnetic anomaly. The fault F2 clearly depicts the Velikonda thrust front, as mentioned before. This anomaly lies just east of 79°15 along the profile. Similarly, the F6 demarcates the ECSB from the NSB. Besides, two seismically detected minor faults (faults F3 and F4) can also be seen in ground magnetics. Further, magnetic anomaly shows lots of undulations over the NSB, as it is characterised by several volcanic features (like Kandra and Kanigiri ophiolite complexes)

2.5 Eastern Dharwar Craton

65

Fig. 2.17 Aeromagnetic total intensity anomaly profiles at an elevation of 1 km (a) and 150 m (b) above the ground from Marur in the west, to Kavali in the east. The location of the profile follows closely to the deep seismic sounding traverse (Kaila et al. 1979). Ground total magnetic intensity anomaly profile along the same region is shown in (c). Corresponding Bouguer and residual gravity anomalies are included in (d) and (e) respectively. Locations of the interpreted faults F1– F4 and F6 along the profile, match well correspondingly with those detected by seismic studies. CB: Cuddapah basin, NFB: Nallamalai Fold Belt, NSB: Nellore Schist Belt, ECSB: East Coast Sedimentary basin. Modified after Fig. 9, Pandey et al. (2018)

66

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associated with supra-subduction of east Antarctica below the Dharwar craton during Mesoproterozoic. Over this region, the gravity anomalies are extremely high as shown in Fig. 2.17d, e. Apart from these fault anomalies, magnetic anomalies are rather flat in other areas of the Cuddapah basin (Rajaram and Anand 2014), suggesting absence of magnetic sources, except in southwestern part. Such low magnetic anomalies have also been observed over the sediments of Palnad sub-basin and the Nallamalai Fold Belt, situated in the northeastern sector of the Cuddapah basin (Kailasam 1976; Konda et al. 2015). Further the gravity as well as magnetic anomalies are low over ECSB, east of the fault F6.

2.5.5 Heat Flow and Lithosphere Structure EDC terrain has been covered very well by heat flow measurements as mentioned before (Fig. 2.7). In cratonic part, some 44 measured heat flow values range from 23.0 to 56.0 mW/m2 , with a mean of 38.1 ± 8.1 mW/m2 . In addition, heat flow has been measured at four locations in Cuddapah basin, which show large variation from 26.8 to 75.2 mW/m2 , with a mean of 54.5 ± 18.0 mW/m2 , thereby indicating that EDC, specially the Cuddapah basin, is associated with comparatively much higher heat flow than WDC. Deep temperature-depth regime was estimated for five locations in EDC, which includes, Dharmavaram, Pavagada area of Closepet granites, Kolar, Agnigundala, and Tummalapalli (Fig. 2.8). Considered heat production models for all these locations are given in Table 2.2. Out of 5 locations, Dharmavaram, Pavagada area of Closepet granites and Kolar, located out side of the Cuddapah basin, show lower surface (39– 40 mW/m2 ) and mantle heat flow (25–28 mW/m2 ) and Moho temperatures around 430–465 °C. Consequently, the lithosphere is thicker at about 120–140 km. At the other two locations (Tummalpalli and Agnigundala), in the Cuddapah basin, surface and mantle heat flow is much higher at 51–75 mW/m2 and 33–58 mW/m2 respectively with high Moho temperatures (510–1030 °C) and much thinner lithosphere (50– 100 km). In fact the geothermal parameters estimated at Agnigundala, which is quite close to collisional boundary, compares fairly well with the Proterozoic Jabera basin in the Son valley Vindhyan locality. In that basin, the estimated heat flow is 78 mW/m2 and lithosphere is much thinner at 50 km (Pandey et al. 2014). These values compare with the unusually warm Cambay basin (Pandey et al. 2017). Using receiver functions, Kumar et al. (2007) too estimated LAB at a shallower depth (80 km) below the CUD broadband seismic station (14°29 , 78°46 ). This station in western Cuddapah region confirms that the crust and mantle lithosphere is much warmer below the plume infected region.

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67

2.5.6 Evolution of Cuddpah Basin and Adjacent Continental Terrain In spite of a large number of geological and geophysical studies since many decades, the evolutionary nature of the Cuddapah basin and associated continental terrains on its east remained enigmatic. However, some acceptable consensus has emerged. For example, the regions lying east of Cuddapah basin, was an active continental margin during Paleoproterozoic/Mesoproterozoic, having subjected to multi-stage growth due to sustained oceanic subduction and magmatism, followed by continentcontinent collision and accretion of a part of east Antarctic segment to the earstwhile eastern Dharwar Craton (Henderson et al. 2014; Chandrakala et al. 2015; Saha et al. 2015). Even some questions, like, the extent and sediment thickness of Cuddapah basin, and whether Nallamalai Fold Belt is an alien terrain or the Eastern Ghats Belt occupies the entire east coast, is now reasonably answered (Chandrakala et al. 2013, 2015, 2017; Pandey et al. 2018).

2.5.6.1

Paleo-Extent of the Cuddapah Basin

As per current understanding, formation of the Cuddapah basin was possibly initiated during early part of Paleoproterozoic, which was synchronous to the mafic dyke swarms intrusions into the underlying granite-gneissic basement (Anil Kumar et al. 2015), and it continued even after that (Anil Kumar et al. 1993; Mallikarjuna Rao et al. 1995; Anand et al. 2003). Analysis of newly derived gravity maps indicate that initial extent of the Cuddapah basin was larger, and possibly extended by almost 50–60 km west of Tadipatri (dashed lines in Fig. 2.15a, b) covering the areas at least to 77°30 longitude on its west, Siddanpalli in Raichur Kimberlite field in the north and up to Kadiri in the south. The entire paleo Cuddapah basin is characterised by large negative gravity anomalies. This inference gets support from the remnant of the Cuddapah basinal rocks occurring at Siddanpalli (Dongre et al. 2008; Chalapathi Rao et al. 2010), as well as Vempalle formations at Kadiri (Riding and Sharma 1998) in which stromatolitic microfabrics are still well-preserved. The originally extended Cuddapah basin, was transgressed by marine water during Paleo to Mesoproterozoic, which led to the deposition of Papaghani (Vempalle formation) and Chitravati Group of rocks, followed by younger sequences. Inundation of the basin by sea water probably continued up to 1.1 Ga, till the uplifting of this basin due to a thermal mantle plume. This plume persisted below GandikotaTadipatri-Parnapalle region, as discussed before, which caused widespread kimberlitic intrusive activity. Since such plumes have an excess temperatures of 100–200 °C compared to background, regional uplift coupled with considerable extension, are likely (White and Mckenzie 1989; Saunders et al. 2007), which resulted into large scale removal of the sediments from the entire region. That may be the reason for only 4 km thick sediments in the central part of the Cuddapah basin now (Chandrakala et al. 2013).

68

2.5.6.2

2 Dharwar Craton

SW Cuddapah Mantle Plume

It is suggested that the southwest Cuddapah mantle plume was responsible for the kimberlitic magmatism (Anil Kumar et al. 1993). Signatures of this plume is well mapped in the seismic images below Parnapalle located in the western part of the basin (Fig. 2.12). Below this region, the mid–level crust (Vp: 6.5 km/s) has upwarped close to the surface, with almost total removal of the granitic-gneissic upper crust. Here, disposition of the crustal layers is made up of almost half crystalline crust and rest half (15–20 km), magmatic crust. This would mean that almost half of the crystalline crust from the bottom got assimilated with the underlying buoyant mantle. Usually mantle plumes are often invoked to explain the genesis of kimberlites (Anil Kumar et al. 1993; Heamann et al. 2003). In Cudappah too, the location of the plume, coincides with the Cuddapah-Kurnool diamond fields. The discrete magmatic bodies are indeed present below this region. Association of a very high conductivity (resistivity 43.0

Ultramafic mantle

0.01

3.0

Data source Liu and Zoback (1997), Reddy et al. (2003), Ray et al. (2003), Vijaya Rao and Prasad (2006) Kalpakkam (North block, SGT )

Surface heat flow: 36.0 mW/m2

0–10.0

Charnockites

0.47

2.5

10.0–20.0

Mafic granulites

0.16

2.5

20.0–30.0

Low velocity granulites

0.50

2.5

30.0–43.0

Magmatic crust

0.02

2.6

>43.0

Ultramafic mantle

0.01

3.0

Data source Liu and Zoback (1997), Reddy et al. (2003), Ray et al. (2003), Vijaya Rao and Prasad (2006) (continued)

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2 Dharwar Craton

Table 2.3 (continued) Depth range (km)

Rock type

Heat production (µW/m3 )

Thermal conductivity (W/m°C)

Aundipatti (South block, SGT )

Surface heat flow: 40.0 mW/m2

0–10.0

Charnockites and gneisses

0.60

2.5

10.0–30.0

Mafic granulites

0.16

2.5

30.0–42.0

Magmatic crust

0.02

2.6

>42.0

Ultramafic mantle

0.01

3.0

Data source Liu and Zoback (1997), Ray et al. (2003), Reddy et al. (2003) Sivakasi (South block, SGT )

Surface heat flow: 45.0 mW/m2

0–10.0

Charnockites

0.60

2.5

10.0–30.0

Mafic granulites

0.16

2.5

30.0–42.0

Magmatic crust

0.02

2.6

>42.0

Ultramafic mantle

0.01

3.0

Data source Liu and Zoback (1997), Ray et al. (2003), Reddy et al. (2003) Karadikuttam (South block, SGT )

Surface heat flow: 55.0 mW/m2

0–5.0

Gneisses

2.93

2.5

5.0–10.0

Granulites

0.50

2.5

10.0–30.0

Mafic granulites

0.16

2.5

30.0–42.0

Magmatic crust

0.02

2.6

>42.0

Ultramafic mantle

0.01

3.0

Data source Rao et al. (1976), Liu and Zoback (1997), Ray et al. (2003), Reddy et al. (2003)

(Namagiripettai and Kalpakkam) belong to northern block and the other three localities (Karadikuttam, Aundipatti and Sivakasi), to southern block. Temperature-depth distribution curves are shown in Fig. 2.23. In the northern block, Moho temperature varies from 510–610 °C, mantle heat flow 24–31 mW/m2 and lithospheric thickness from 100 to 135 km. In comparison, Moho temperatures (590–660 °C) and mantle heat flow (31–36 mW/m2 ) are relatively higher in southern block and consequently, lithosphere is thinner at around 85–100 km. A plausible lithospheric section, as deduced from currently available geoscientific data over WDC, EDC and SGT is shown in Fig. 2.24 which indicates considerable upwarping of asthenosphere below SGT.

2.6 Southern Granulite Terrain

79

Fig. 2.23 Estimated temperature-depth distribution beneath chosen geotectonic segments of Southern Granulite Terrain located in southernmost part of south Indian shield, based on crustal heat production models as given in Table 2.3. Locational details are given in Table 9.1, Chap. 9. Depths are in km

80

2 Dharwar Craton

Fig. 2.24 A possible lithospheric cross section beneath WDC, EDC and SGT, as inferred from available geological and geophysical informations. WDC: Western Dharwar craton, EDC: Eastern Dharwar craton, SGT: Southern Granulite Terrain. Modified after Fig. 10, Agrawal and Pandey (2004)

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Ravi Kumar M, Saikia D, Singh A, Srinagesh D, Baidya PR, Dattatrayam RS (2013) Low shear velocities in the sub-lithospheric mantle beneath the Indian shield? J Geophys Res Solid Earth 118:1–14. https://doi.org/10.1002/jgrb.50114 Ray L, Senthil Kumar P, Reddy GK, Roy S, Rao GV, Srinivasan R, Rao RUM (2003) High mantle heat flow in a Precambrian granulite province: evidence from southern India. J Geophys Res 108(B2):2084. https://doi.org/10.1029/2001JB000688 Reddy PR, Chandrakala K, Sridhar AR (2000) Crustal velocity structure of the Dharwar craton, India. J Geol Soc India 55:381–386 Reddy PR, Rajendra Prasad B, Vijaya Rao V, Sain K, Prasada Rao P, Khare P, Reddy MS (2003) Deep seismic reflection and refraction/wide-angle reflection studies along Kuppam–Palani transect in Southern Granulite Terrain of India. In: Ramakrishnan M (ed) Tectonics of Southern Granulite Terrain, Kuppam–Palani Geotransect. Geol Soc India Mem 50:79–106 Reddy PR, Chandrakala K, Prasad ASSSRS, Rao ChR (2004) Lateral and vertical crustal velocity and density variations in the southwestern Cuddapah basin and adjoining eastern Dharwar craton. Curr Sci 87:1607–1614 Riding R, Sharma M (1998) Late Paleoproterozoic (∼1800–1600 Ma) stromatolites, Cuddapah basin, southern India: cyanobacterial or other bacterial microfabrics? Precamb Res 92:21–35 Rogers JJW, Callahan EJ (1987) Radioactivity, heat flow and rifting of the Indian continental crust. J Geol 95:829–836 Rogers JJW, Santosh M (2002) Configuration of Columbia, a Mesoproterozoic supercontinent. Gondwana Res 5:5–22 Roy S, Rao RUM (1999) Geothermal investigations in the 1993 Latur earthquake area, Deccan volcanic province, India. Tectonophysics. 306:237–252 Roy S, Rao RUM (2000) Heat flow in the Indian Shield. J Geophys Res 105(B11):25587–25604 Roy S, Ray L, Senthil Kumar P, Reddy GK, Srinivasan R (2003) Heat flow and heat production in the Precambrian gneiss-granulite province of southern India. Mem Geol Soc India 50:177–191 Roy S, Ray L, Bhattacharya A, Srinivasan R (2007) New heat flow data from deep boreholes in the greenstones-granite-gneiss and gneiss-granulite provinces of south india. DCS-DST News Lett. 17(1):8–11 Roy S, Ray L, Bhattacharya A, Srinivasan R (2008) Heat flow and crustal thermal structure in the late Archean Closepet granite batholith, south India. Int J Earth Sci 97:245–256 Rudnick RL, Fountain DM (1995) Nature and composition of the continental crust: a lower crustal perspective. Rev Geophys 33:267–309 Rudnick RL, Gao S (2003) Composition of the continental crust, treatise on geochemistry, vol 3. Elsevier, New York, pp 1–64 Saha D (2002) Multi stage deformation in the Nallamalai fold belt, Cuddapah basin, south India— implications for Mesoproterozoic tectonism along the southeastern margin of India. Gondwana Res 5:701–719 Saha D (2011) Dismembered ophiolites in Paleoproterozoic nappe complexes of Kandra and Gurramkonda, south India. J Asian Earth Sci 42:158–175 Saha D, Tripathy V (2012) Palaeoproterozoic sedimentation in the Cuddapah basin, south India and regional tectonics–a review. In: Mazumder R, Saha D (eds) Paleoproterozoic of India, vol 365. Geological Society of London Special Publication, London, pp 159–182 Saha D, Chakraborti S, Tripathy V (2010) Intracontinental thrusts and inclined transpression along eastern margin of the East Dharwar craton, India. J Geol Soc India 75:323–337 Saha D, Sain A, Nandi P, Mazumder R, Kar R (2015) Tectonostratigraphic evolution of the Nellore schist belt, southern India, since the Neoarchaean. Geol Soc Lond Mem 43:269–282 Saikia U, Das R, Rai SS (2017) Possible magmatic underplating beneath the west coast of India and adjoining Dharwar craton: imprint from Archean crustal evolution to breakup of India and Madagascar. Earth Planet Sci Lett 462:1–14 Sarkar D, Chandrakala K, Padmavathi Devi P, Sridhar AR, Sain K, Reddy PR (2001) Crustal velocity structure of western Dharwar craton, south India. J Geodynamics 31:227–241

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Vijaya Rao V, Sain K, Rajendra Prasad B (2007) Dipping Moho in the southern part of eastern Dharwar craton, India, as revealed by the coincident seismic reflection and refraction study. Current Sci 93:330–336 Vijaya Rao V, Murty ASN, Sarkar D, Bhaskar Rao YJ, Khare P, Prasad ASSSRS, Sridher V, Raju S, Rao GS, Karuppannan P, Kumar NP, Sen MK (2015) Crustal velocity structure of the Neoarchean convergence zone between the eastern and western blocks of Dharwar craton, India from seismic wide-angle studies. Precambrian Res. 266:282–295 White R, McKenzie D (1989) Magmatism at rift zones: the generation of volcanic continental margins and flood basalts. J Geophys Res 94:7685–7729 Yellappa T, Rao JM (2018) Geochemical characteristics of Proterozoic granite magmatism from Southern Granulite Terrain, India: implications for Gondwana. J Earth Syst Sci 127:22. https:// doi.org/10.1007/s12040-018-0923-6 Yoshida M, Jacobs J, Santosh M, Rajesh HM (2003) Role of Pan-African events in the Circum-East Antarctic Orogen of East Gondwana: a critical overview. In: Yoshida M, Windley BF, Dasgupta S (eds) Proterozoic East Gondwana: upper continent assembly and breakup, vol. 206. Geological Society London Special Publications, London, pp 57–75 Zelt CA (1999) Modeling strategies and model assessment for wide-angle seismic traveltime data. Geophys J Int 139:183–204 Zelt CA, Smith RB (1992) Seismic travel time inversion for 2-D crustal velocity structure. Geophys J Int 108:16–34 Zhao GC, Sun M, Wilde SA, Li S (2004) A Paleo-Mesoproterozoic supercontinent: assembly, growth and breakup. Earth Sci Rev 67:91–123

Chapter 3

Singhbhum and Bastar Cratons

3.1 Singhbhum Craton 3.1.1 Introduction More than four decades ago, Negi et al. (1986, 1987) postulated that the Indian shield does not possess characteristics of a typical stable shield. Singhbhum craton, which is one of the five well recognised cratons of the Indian shield (Naqvi and Rogers 1987; Ramakrishnan and Vaidyanadhan 2008; Sharma 2009), is perhaps the finest example of that. It is surrounded from almost all the sides by active mobile belts, which have been repeatedly reactivated from Mesoproterozoic to present. This craton is located in the eastern part of the Indian shield, between lat. 21º– 23º 15 N and long. 84º 40 –86º 45 E, and covers an area of about 40,000 km2 in the states of Jharkhad and Orissa. It forms one of the most deformed and degenerated segment of the Earth, containing relics of some of the oldest rocks that date back around 4.2 Ga (Acharya et al. 2010a; Chaudhuri et al. 2018). It is also characterized by an unusually thin (65–100 km) and warm lithosphere (Pandey and Agrawal 1999; Mandal 2017a), compared to 250–350 km in similar terrains elsewhere in the world (Chapman and Pollack 1977; Polet and Anderson 1995). Besides, it has been associated with sustained magmatism, which lasted for over two billion years, from Mesoarchaean (3.3 Ga) to Neoproterozoic (~1.0 Ga). Tectono-magmatic processes associated with the evolution of SONATA geo-fracture and Mahanadi graben, together with the burst of Crozet/Kerguelen plumes, seems to have played a major role in restructuring the crust-mantle structure beneath this cratonic block. It has been very well studied geologically and geochemically, but the geophysical structure is not too well understood, therefore deciphering its geodynamic and rheological nature, have always been a challenge.

© Springer Nature Switzerland AG 2020 O. P. Pandey, Geodynamic Evolution of the Indian Shield: Geophysical Aspects, Society of Earth Scientists Series, https://doi.org/10.1007/978-3-030-40597-7_3

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3.1.2 Geological Setting This craton (Fig. 3.1) is bordered by Singhbhum Shear Zone in the north which is in close proximity of Proterozoic Chotanagpur gneissic complex, while in south, it is bounded by the Sukinda Thrust zone and Mahanadi graben and in the east, by Eastern

Fig. 3.1 Simplified geological map of the Singhbhum craton, eastern India (based on Saha 1994; Sengupta et al. 1997; Mondal et al. 2006). OMG: Older Metamorphic Group, OMTG: Older Metamorphic Tonalitic Gneiss, IOG: Iron Ore Group. Inset map shows the location of Singhbhum Craton (SiC). ArC: Aravalli craton, BuC: Bundelkhand craton, BaC: Bastar craton, DhC: Dharwar craton

3.1 Singhbhum Craton

91

Ghats Belt (Mondal et al. 2006). It has been associated with large-scale sedimentation, deformation, sustained episodic magmatism and complex history of thermogeodynamic evolution (Saha 1994; Sengupta et al. 1997; Mazumder et al. 2000; Misra 2006; Bose 2009; Acharya et al. 2010a, b; Chaterjee et al. 2013; Manikyamba et al. 2015; Singh et al. 2017). The main cratonic segment (Fig. 3.1) principally consists of old metamorphic gneisses (OMG), intruded by gneissic tonalities (OMTG), composite granitic batholiths and supracrustal rocks of the IOG greenstone sequence (Saha 1994; Sengupta et al. 1997; Mondal et al. 2006; Acharya et al. 2010a, b; Mazumder et al. 2012), apart from number of mafic/ultramafic volcanic sequences like, Ongarbira, Simlipal, Bonai (Malangtoli), Dalma, Dhanjori, Jagannathpur and Newer Dolerites. These volcanic suites, which are of Paleoarchean to Neoproterozoic in age, cover a major part of this craton (Saha 1994; Mukhopadhyay 2001; Roy et al. 2005; Mondal et al. 2006). Almost all the volcanic exposures are located in areas containing Precambrian metasediments. The OMG rocks, being the oldest supracrustal suite in the region, are largely made up of meta-igneous and meta-sedimentary rocks of amphibolites facies, characterized by zircon U-Pb and 207 Pb/206 Pb ages of 3.5–3.6 Ga (Misra et al. 1999; Mukhopadhyay et al. 2008). These are metamorphosed to the tune of 5.5 kbar and 660–630 °C, which would indicate substantial crustal thickening during Archaean era (Sharma 2009). These rocks are further intruded by relatively younger gneissic tonalities (OMTG rocks) (Blackburn and Srivastava 1994), having a zircon U–Pb age of 3.44 Ga (Misra et al. 1999), which occur as remnants within the Singhbhum granite batholithic complex (Mondal et al. 2006) located interior part of the craton. These batholithic rocks can be sub-divided into two main phases: Type A Granite, which is associated with zircon U–Pb date of 3328 Ma (Misra et al. 1999) and Type B granites, that are relatively younger and dated close to 3100 Ma by Pb–Pb whole rock isochron method (Saha 1994). On the contrary, IOG greenstone sequence by and large contains Banded Iron Formations, surrounding the Singhbhum granitic batholiths (Fig. 3.1), presumably developed after the emplacement of type “A” Singhbhum granites. Granitoid rocks of trondhjemitic and tonalitic compositions have intruded the supracrustals of the IOG greenstone sequence that are characterized by about 3.2 Ga age (Moorebath et al. 1986; Sengupta et al. 1991; Sharma et al. 1994), indicating that IOG belt rocks are older than 3.2 Ga (Sengupta et al. 1997). Based on available isotopic informations, the age of IOG supracrustal sequences lies between the evolutionary stages of Singhbhum granites i.e. between 3.3 and 3.1 Ga (Saha 1994; Mondal et al. 2006). These are mainly present in three sedimentary basins, (i) the Jamda-Koira basin (ii) GurumahisaniBadampahar basin, and (iii) the Tomka-Daitari basin (Mondal et al. 2006) and contain several mafic-ultramafic occurrences. The ultramafics are mostly chromite-bearing and interpreted either to be cumulates or mantle derived rocks, similar to the Alpine type peridoties or ophiolites (Banerjee 1972; Page et al. 1985; Verma 1986; Pal and Mitra 2004; Mondal et al. 2006). Some small patches of rarely occurring deep mantle plume generated Paleoarchean ferropicrites and ferrobasalts, found as remanants, have also recently been discovered from the Iron Ore Group (IOG) greenstone belt,

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3 Singhbhum and Bastar Cratons

situated in the northern part of the craton. Some studies have indicated that these volcanics have subduction related affinity (Raza et al. 1995; Sengupta et al. 1997), similar to those of other volcanics in the region like Jagannathpur (Alvi and Raza 1991; Manikyamba et al. 2015), Dhanjori (Alvi and Raza 1992; Sengupta et al. 1997) and Malagtoli (Singh et al. 2017).

3.1.3 Geophysical Signatures 3.1.3.1

Crustal Seismic Structure

No deep seismic sounding (DSS) studies have yet been undertaken over the Singhbhum craton. However, two DSS profiles are shot in Mahanadi delta from Konark to Mukundpur and from Paradip Port to Kabatabandha (Behera et al. 2004). Depthwise siesmic velocity distribution below Mukundpur and Kabatabandha (location shown in Fig. 3.2), situated in southernmost part of the Singhbhum craton, are given Table 3.1, which reveal presence of pronounced low velocity zone at mid crustal level, besides thick magma underplating possibly due to Crozet and Kerguelen hot spot activity. This craton is, however, very well investigated through receiver function studies, as discussed below. Receiver Function Studies The detailed crust-mantle structure of the Singhbhum craton and its adjoining regions like Chotanagpur Granitic-Gneissic Terrain and some segments of the Eastern Ghats Belt, has recently been studied by Mandal and Biswas (2016) and Mandal (2017a, b), using fifteen 3-component broadband seismic network for about three years. They provided shear velocity structure extending down to a depth of 300 km, obtained through the inversion modelling of P-receiver functions. Crustal velocity models over these regions suggest a Vs of more than 3 km/s near the surface, which reaches as much as 4.4–4.6 km/s at the crust-mantle boundary. This study indicated, presence of mafic, heterogeneous and deformed crustal structure below the Singhbhum craton and Eastern Ghats Belt (average crustal Vs around 4.0 km/s), compared to Chotanagpur Granitic-Gneissic Terrain in the north, which is characterised by a little lower average crustal Vs of 3.9 km/s, reflecting its relatively felsic nature. The receiver function modelling further suggested variation of crustal thickness from 35 km at seismic station CHI (Chaibasa) to 47 km at SAL (Simlipal complex), while the lithospheric thicknesses varied from 78 km at CHI to 100 km below station DEN, as detailed in Table 3.2. The Moho depths beneath this region ranged from 35 to 47 km with an average of 42 ± 4 km and the LAB, from 78 to 100 km with an average of 86 ± 8 km (Mandal 2017a). The thinnest crust (35 km), as well as lithosphere (78 km) is obtained below CHI station, located in Chaibasa region in the northern part of the craton. This station is not far off from the Singhbhum shear zone, which is associated with more than 100 km long Dalma volcanic belt, arguably influenced by 1.6 Ga plume activity. In comparison, the thickest crust of 47 km, overlying a 81 km

3.1 Singhbhum Craton

93

Fig. 3.2 Morpho-structural pattern over Singhbhum craton as deduced from Rantsman et al. (1995), showing correlation with various volcanic suites (shaded regions). Solid dots refer to the locations of Newer dolerite group of rocks in which rare occurrence of native iron is reported through rock magnetic studies (Verma and Prasad 1975). Two DSS profiles (Konark to Mukundpur and Paradip Port to Kabatabandha) which touch the southernmost parts the craton is shown by solid dashed lines. K: Kabatabandha, M: Mukundpur. Solid squares indicate heat flow locations with values in mW/m2 (Gupta and Rao 1970) Table 3.1 Crustal velocity distribution below Mukundpur and Kabatabandha, located close to southern part of Singhbhum craton based on Behera et al. (2004)

Mukundpur Depth (km)

Kabatabandha Vp (km/s)

Depth (km)

Vp (km/s) 6.1

0.0–4.0

6.1

0.0–3.0

4.0–16.5

6.5

3.0–9.0

6.5

16.5–20.5

6.0

9.0–19.0

6.0

20.5–26.5

7.1

19.0–26.5

7.0

26.5–39.0

7.5

26.5–36.0

7.5

>39.0

8.1

>36.0

8.1

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3 Singhbhum and Bastar Cratons

Table 3.2 Estimated Moho, lithosphere and other geophysical parameters below Singhbhum craton, eastern Indian shield as deduced from the study of Mandal (2017a). Station

Location

Crustal thickness (km)

Lithospheric thickness (km)

Avg crustal Vs (km/s)

Estimated crustal Vp (km/s)

Underplated magma layer (Vs > 4.0 km/s)

Singhbhum craton DEN

20.92° N 85.25° E

44

100

3.95

6.91

8

KNJ

21.28° N 86.20° E

42

92

4.00

7.00

12

BLS

21.34° N 86.66° E

44

85

4.00

7.00

14

SAL

21.78° N 86.83° E

47

81

4.10

7.18

21

SRK

22.63°N 86.11° E

41

81

4.10

7.18

17

CHI

22.46° N 85.79° E

35

78

3.90

6.83

5

BAL

22.30° N 86.83° E

41

82

3.95

6.91

13

42 ± 4

Mean

86 ± 8

Chotanagpur granite gneiss terrain LOH

23.37° N 84.65° E

41

83

3.90

6.83

11

RAN

23.43° N 85.43° E

42

100

3.90

6.83

2

HAZ

24.07° N 85.38° E

42

82

3.92

6.86

0

NRS

23.86° N 86.66° E

42

83

3.90

6.83

2

DMK

24.60° N 87.08° E

42

100

3.90

6.83

4

42 ± 1

90 ± 9

Mean

Thickness of underplated magma layer is in km. For estimating Vp , Vp /Vs is taken as 1.75

thick lithosphere, was obtained below the seismic station SAL, located in Simlipal volcanic complex, which is infected by repeated episodes of magma infusion. Average lithospheric thickness 86 ± 8 km below this craton, as obtained by receiver function studies, is comparable with the earlier estimate of 65 km derived from scanty heat flow data (Pandey and Agrawal 1999). Thus, seismically defined LAB below this craton is possibly the lowest recorded anywhere in global shield terrain, only exception may be the North China craton. Chen et al. (2008) studied lateral variation of LAB based on S-receiver function below NE part of that craton,

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Fig. 3.3 A plot between averaged Vs and computed Vp (using Vp /Vs : 1.75) distribution with depth beneath Singhbhum craton, using Vs data from receiver function studies at various locations by Mandal (2017a). For comparison, similar plots have been added for south Indian shield (Pandey et al. 2013), as well as global shields and platform (Christensen and Mooney 1995). Inferred crustal compositional structure is also included

which is also one of the oldest in the world, and found LAB at a depth of 60–70 km in the southeast basin and coastal areas, and not more than 140 km in the northwest mountain ranges and continental interiors. (i) Average crustal seismic structure Figure 3.3, shows a plot between averaged Vs (based on data from all seismic stations) and computed Vp (using Vp /Vs : 1.75) distribution with depth below the Singhbhum craton. For comparison, similar plots are incorporated to this figure for south Indian shield (Pandey et al. 2013) and global shields and platforms (Christensen and Mooney 1995). It is interesting to note that estimated crustal Vs as well as Vp obtained below this craton, is much higher than that recorded in south Indian shield and global shields and platform. It is further observed that upper graniticgneissic crust, as well the mid-level crust, is only about 2 km thick each (Fig. 3.3), implying that a big chunk of the crust has been eroded away in due course of time. The mid-level crust is followed by almost a 40 km thick lower crust, which can be subdivided into three parts. A typical lower crust (Vs : 3.82–3.93 km/s) between 4 and 14 km, followed by a pronounced metasomatised lower crust between 14 and 30 km depth and then magmatically infused crust till 44 km depth. In the metasomatised zone, velocity drops down sharply from 3.93 to 3.76 km/s at about 18–20 km depth. Metasomatism is quite common in the mid to lower crustal rocks due to their interaction with mantle-derived hydrothermal fluids and halogens. It is quite prevalent in magmatically affected areas like Deccan traps (Karmalkar and Rege 2002; Desai et al. 2004; Sen et al. 2009; Tripathi et al. 2012; Pandey et al. 2014). Metasomatism

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3 Singhbhum and Bastar Cratons

can lead to drastic reduction in crustal velocities up to 15% (Pandey et al. 2016). Beneath this terrain, Moho lies at an average depth of about 44 km. (ii) Geodynamic perspective A close look at the Fig. 3.3 indicates that the velocities obtained at the shallow depth of about four kilometres below this craton (below which lower crust is envisaged), is quite similar to velocities obtained at the depth of 25–30 km in global shields and platforms (Christensen and Mooney 1995). This would mean that the crust below this craton is unusually mafic and may have uplifted or exhumed to about 20 km or so. A major chunk of the upper granitic and underlying intermediate crust, appears to have been eroded from a large part of the Singhbhum craton, specially from volcanic infected and nearby regions. Alternate explanation could be that, this craton is infected by heavy dose of infused magma, which appears very likely in view of the thick (on an average 14 km) magma underplating (characterised by Vs > 4.0 km/s) (Fig. 3.3). These inferences are compatible with the occurrence of reasonably high heat flow, thin lithosphere (as discussed in the heat flow section), as well as relatively positive gravity anomaly. Reportedly, this craton has been subducting below the Proterozoic Chotanagpur Granitic-Gneissic Terrain in the north around 1.0–1.6 Ga (Mukhopadhyay 1990; Mahato et al. 2008; Rekha et al. 2011; Mandal 2017a).

3.1.3.2

Gravity Studies

Qureshy et al. (1972) were probably the first to study the gravity field over this region, who reported preliminary results pertaining to some parts of the south Singhbhum craton. This study was further followed by a number of investigations, specially by the researchers of Indian School of Mines, Dhanbad (Verma et al. 1978, 1984, 1996; Das and Agrawal 2001). Recently, a new Gravity Map Series of India (GSI-NGRI 2006) was also launched by CSIR-National Geophysical Research Institute of India and Geological Survey of India at 1:2,000,000 scale with 5 mGal contour interval, but essentially, the gravity features over the Singhbhum craton remain the same, as deciphered earlier. A Bouguer anomaly map of this craton is shown in Fig. 3.4, which is a modified version of Verma et al. (1984, 1996) and Das and Agrawal (2001). A large gravity anomaly variation from +10 mGal to about −60 mGal can be seen over different parts, which corroborates well with the main geological units, like granitic batholiths, volcanic suites, Iron Ore Group lavas, and Proterozoic gabbro-anorthosite- ultramafic complexes. As mentioned before, this region was subjected to a number of orogenic cycles, which has left an imprint on the gravity field in the form of gravity highs. These gravity highs are generally associated with the synclinal structures filled with metasediments and volcanics, such as the Iron Ore Group, Dhanjori and Simlipal basins, which can be seen bordering the Singbhum granite batholiths. These batholiths, located in the central part of the craton and extending laterally for more than 150 km from north

3.1 Singhbhum Craton

97

Fig. 3.4 Bouguer anomaly map (in mGal) of Sighbhum craton. JSR: Jamshedpur, MSB: Moosabani, CBA: Chaibasa, NMD: Noamundi, BDM: Badampahar, KTP: Kaptipada, KJG: Keonjhargarh, PLH: Pallalahara. v: volcanics, x: Iron Ore Group lavas and ultramafics, slanted dashed lines: Proterozoic gabbro-anorthosite suites. Modified after Fig. 2, Verma et al. (1984)

to south, are conspicuously associated with gravity lows. Three prominent gravity trends can be seen in Fig. 3.4; these are: East-west gravity high (0 to −20 mGal) in the northern part of the craton, associated with Dalma volcanics, metamorphic variants, and metasediments. The gravity high reflects presence of high density upper mantle rocks at sub surface depths. (ii) Arcuate north-south trend on the western side of the craton, associated with positive gravity over IOG basin and Ongarbira volcanics. The IOG terrain is known to be underlain by thick volcanics. In contrast, southern part of this anomaly trend coincides with the gravity low associated with the Bonai volcanics and granitic batholith. Over the granitic batholith, gravity anomalies reach as low as −60 mGal, possibly due to extrusion of large thickness of low (i)

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Table 3.3 Density of basic and granitic rocks from Singhbhum craton Rock types

No. of samples

Density range (g/cm3 )

Mean density (g/cm3 )

Dalma metavolcanics

21

2.9–3.2

3.05

Dhanjori lavas

10

3.0–3.07

3.05

Simlipal lavas

23

2.7–3.3

2.95

Gabbro-anorthosites

13

2.88–3.09

3.01

Amphibolites

18

2.7–3.04

3.01

Unclassified

96

2.9–3.05

2.98

Mean of all basic rocks

181

2.7–3.2

3.01

Singhbhum granite basement complex

213

2.59–2.72

2.68

Data source Verma et al. (1978), Sarma (1980)

density (2.68 g/cm3 ) granitic rocks (Table 3.3), which are in close association with high density (3.01 g/cm3 ) volcanic rocks or its equivalents. Low gravity observed over the Bonai volcanics, compared to positive values over other volcanics, may partly be attributed to the felsic nature of the andesites to andesitic lavas, which are reportedly generated under subduction scenario (Singh et al. 2017). (iii) The north-south parallel trend east of the Singhbhum granites. It coincides with the Dhanjori and Simlipal volcanics and with the gabbro-anorthosite suite within the granite terrain down south, where Bouguer gravity anomalies reach as high as 0 mGal, compared to −40 mGal in surrounding areas. Further, between the above mentioned two north-south trending anomalies, the region occupied by large granitic batholith (Singhbhum granites), is associated with a broad gravity low. This granitic batholith region, may be bounded by deep faults as evidenced by steepness of the gravity anomalies on either sides. Usually, positive gravity anomalies are caused by mantle plumes or mantleinduced rift magmatism. These volcanics are associated with high density of 3.01 g/cm3 (Table 3.3), compared to that of Deccan traps density of 2.74 g/cm3 (Vedanti et al. 2018), which would indicate deep-level melting of the magma source, possibly by a mantle plume. This conjecture is supported by the unprecedented presence of a 21 km thick magma layer between 26 and 47 km depth (Vs > 4.0 km/s) below Simlipal volcanic complex (Fig. 3.5). The circular ring-type morphostructures associated with this volcanic region (Fig. 3.1), cannot be created only by subduction scenario. In Fig. 3.6, gravity profiles are plotted along the latitudes 22° and 22° 30 N, which passes through Ongarbira, Dhanjori and Simlipal volcanic terrain and exposures of ultramafics and lavas. Elevation along the latitudes 22° is also plotted in the figure. This figure reveals that over the volcanic terrains, apart from the gravity field, elevations are also quite high. In fact this craton appears elevated by almost 500 m

3.1 Singhbhum Craton

99

Fig. 3.5 S-wave velocity variation with depth below seismic station SAL, located in Simlipal complex, based on receiver function studies. LAB: lithosphere-asthenosphere boundary. Inferred crustal compositional structure below this station is also added. Modified after Fig. 9, Mandal (2017a)

compared to surrounding regions (Fig. 3.6). This provides further credence to the presence of a long lasting mantle plume below this region, which led thermal-induced upflifting. It is quite likely that this craton has exhumed by more than 15 km, if one takes into account the recorded P and T conditions (Sharma 2009) as mentioned earlier. The Fig. 3.3. would confirm this conjecture. No other mechanism, except thermal interaction with mantle, can possibly explain it. The recent study by Singh et al. (2015), based on satellite gravity, geoid anomaly and topography, also reported distinct vertical density heterogeneities throughout the lithosphere. Thus, the Singhbhum craton associated largely with positive gravity bias, can be considered as one of the finest examples of strong crust-mantle thermal interaction in an erstwhile stable shield region due to underlying hot mantle conditions. These characters contrast from that of the Dharwar craton, which is characterized by highly negative gravity anomalies of up to −120 mGal (Fig. 2.5, Chap. 2 ) and a much lower heat flow and mantle temperatures.

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3 Singhbhum and Bastar Cratons

Fig. 3.6 Bouguer gravity anomaly variations along latitude 22° and 22° 30 N, covering some of the prominent volcanic suites of the Singhbhum craton. Elevation along the latitude 22° is also plotted. This figure shows strong association of high gravity anomalies as well as elevation with volcanic suites, indicating regional uplift

3.1.3.3

M. T. Studies

Magnetotelluric (MT) studies have been quite useful in defining Moho and lithosphere-asthenosphere boundary (LAB) in many global areas as detailed by Shalivahan et al. (2014). Over this craton, such studies were attempted initially by Roy et al. (1989), who carried out single site MT measurement covering Archean terrain of Singhbhum granite batholith, along a 100 km long NE-SW profile, from Bangriposhi to Keonjhar in its eastern segment. One-dimensional inversion, revealed depth to the Moho from 23 to 40 km and depth to the LAB between 58 and 76 km. Similar study was further followed by Bhattacharya et al. (2000), Bhattacharya and Shalivahan (2002) and Shalivahan et al. (2014), who made MT investigation over a transect

3.1 Singhbhum Craton

101

across Rairangpur region, falling over Singhbhum granites. This study was part of a larger experiment carried out over the east Indian craton. During this experiment, they acquired remote reference MT data in the frequency range of 320–0.0055 Hz and obtained an electrically homogeneous 38 km thick granitic crust, characterized by very high resistivity of 40,000 -m. This was followed by 8 km thick transitional layer with relatively low resistivity (8500 -m). Moho was detected at 46 km depth. Resistivity of the upper mantle below the Moho, was found to be about 750 -m (Bhattacharya and Shalivahan 2002). Similarly, LAB was detected at 95 km depth (Shalivahan et al. 2014), which is less than half to that obtained under the Archean Slave craton located in the north-western part of the Canadian Shield by similar studies (Jones et al. 2003). Typical values for the resistivity recorded in the electrical asthenosphere is usually less than 100 -m, which can be explained by the presence of a small amount of water, that induces partial melting at shallower depths. Interestingly, depth to the LAB as obtained by MT study, is almost same to that of 86 ± 8 km, inferred from receiver function studies (Mandal 2017a), as mentioned earlier. Thinning of the lithosphere to such an extent has been attributed by these authors to the delamination of the lithospheric roots beneath East India cratonic segments. Recently, a discovery of 65 Ma diamondiferous kimberlites was also made in the Bastar craton, which was synchronous to Deccan flood basalts. Based on this finding, Chalapathi Rao et al. (2011) and Lehmann et al. (2010) indicated that the lithospheric roots of the Indian cratons were intact partially till at least 65 Ma. Based on this argument, Shalivahan et al. (2014) advocated that compression generated by Himalayan orogeny may have played a role in the further delamination of lithospheric roots. This could be one reason, but the role played by early Cretaceous-early Tertiary geodynamic events after the separation of India from east Antarctica in thinning the lithosphere, could be an equally important cause (Pandey and Agrawal 1999, 2000; Kumar et al. 2007; Pandey 2016; Pandey et al. 2017).

3.1.3.4

Heat Flow and Lithosphere Studies

In this geologic terrain, first measurement of terrestrial heat flow was carried out at Mosabani mine (Verma et al. 1966), followed further by two more measurements at Rakha (a copper prospect, situated 14 km northwest of Mosabani) and Narwapahar (a uranium prospect) (Rao and Rao 1974). All the three sites are located in the Singhbhum thrust zone, which separates high-grade metamorphic rocks in the north to relatively lower grade rocks in the south. Measured heat flow varied in a narrow range from 58.9 to 62.7 mW/m2 (Fig. 3.2), with a mean of 60.7 ± 1.6 mW/m2 , suggesting that the surface heat flow is reasonably high and fairly uniform over this region, compared to the south Indian shield. However, it should be remembered that all the three heat flow sites are located over the mineral prospects, thus may not be truly representative of the cratonic terrain. Rao et al. (1976) has further reported a low heat generation of 1.15 µW/m3 pertaining to the Singhbhum granite outcrop

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3 Singhbhum and Bastar Cratons

close to the thrust belt, indicating that they must have exhumed to the surface from deeper depths. Lithosphere Structure If we have a look at the Fig. 3.3, mafic lower crust is quite close to the surface, only at a depth of 4 km, besides enormous magma underplating underneath. Based on the 5-layer crustal model deciphered for this region from seismic studies, and taking into account heat generation of 1.15 µW/m3 , pertaining to the Singhbhum granites (Table 3.4), temperature-depth estimations was made, which is shown in Fig. 3.7. This figure reveals a lithospheric thickness of only about 60 km, similar to that reported earlier by Pandey and Agrawal (1999). Estimated Moho temperatures Table 3.4 Crustal heat production models and adopted geothermal parameters for temperaturedepth estimation beneath Singhbhum, and Bodal and Malajkhand areas of Bastar craton Depth range (km)

Rock type

Heat production (µW/m3 )

Thermal conductivity (W/m °C)

Surface heat flow: 61 mW/m2

Singhbhum craton 0–2.0

Granite-gneiss

1.15

3.0

2.0–4.0

Amphibolite facies rocks

0.78

2.88

4.0–14.0

Mafic granulite

0.20

2.50

14.0–30.0

Metasomatised rocks

0.50

2.50

30.0–44.0

Magmatic crust

0.02

2.6

>44.0

Ultramafic mantle

0.01

3.0

Data source Rao et al. (1976), Roy and Rao (1999), Ray et al. (2003), Pandey et al. (2017), Mandal et al. (2017a) Bodal (Bastar craton)

Surface heat flow: 64 mW/m2

0–7.0

Granite

2.97

3.0

7.0–16.0

Amphibolite-granulite

0.78

2.88

16.0–22.0

Metasomatised rock

0.50

2.5

22.0–44.0

Mafic granulite

0.16

2.5

>44.0

Ultramafic mantle

0.01

3.0

Data source Gupta et al. (1993), Roy and Rao (1999), Ray et al. (2003), Murty et al. (2004), Pandey et al. (2017) Malajkhand (Bastar craton)

Surface heat flow: 52 mW/m2

0–7.0

Granodiorite

2.1

3.0

7.0–16.0

Amphibolite-granulite

0.78

2.88

16.0–22.0

Metasomatised rock

0.50

2.50

22.0–44.0

Mafic granulite

0.16

2.50

>44.0

Ultramafic mantle

0.01

3.0

Data source Gupta et al. (1993), Roy and Rao (1999), Ray et al. (2003), Murty et al. (2004), Pandey et al. (2017)

3.1 Singhbhum Craton

103

Fig. 3.7 Estimated temperature-depth distribution beneath Singhbhum craton and Bodal and Malajkhand areas of Bastar craton, based on crustal heat production models as given in Table 3.4

and heat flow input from the mantle is found to be equally high at about 910 °C and 47 mW/m2 respectively. Comparison With Other Geophysical Studies In a very first lithospheric studies, Negi et al. (1986, 1987) obtained a thickness of 53 km for this craton, which was later updated to 65 km by Pandey and Agrawal (1999). It may be mentioned here that these estimates are based on the data from Singhbhum Thrust Zone, thus need to be treated cautiously. However, based on MT investigations, Roy et al. (1989) too gave a similar lithospheric thickness of 58–76 km. Recent MT estimate (Shalivahan et al. 2014), however put it little deeper at about 95 km below the granitic batholithic region. Receiver function studies by Mandal (2017a), who covered the entire cratonic terrain also came out with similar depths (86 ± 8 km) to the LAB (Table 3.2). Thus the thermally–derived lithospheric thickness of 60 km, is in reasonable agreement with other geophysical studies. The only other estimate that goes beyond 100 km is 130–140 km made by Singh et al. (2015),

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3 Singhbhum and Bastar Cratons

which comes from the analysis of satellite gravity, geoid anomaly and topography data. Nevertheless, fact remains that the cratonic keel beneath this craton, is far less than the 250–350 km obtained in other Archean shield terrains (Chapman and Pollack 1977; Polat and Anderson 1995, Artemieva and Mooney 2001; Kumar et al. 2007). Large Igneous Province (LIP) and Thermo-geodynamic Degeneration There is a recent report that a Large Igneous Province (LIP) probably existed during Neoarchean and covered an area of almost 30,000 km2 of this craton (Anil Kumar et al. 2017). In that case, a deep mantle plume must have been active below this region, as mentioned earlier. There are also suggestions that the Singhbhum volcanic suites have subduction related affinity, based on geochemical investigations (Raza et al. 1995; Sengupta et al. 1997; Manikyamba et al. 2015; Singh et al. 2017). But then, it would be very difficult to explain, how Benioff-zone related magmatism led to episodic volcanism for such a large span of almost 2.0 Ga, that covered the entire craton. There is another interesting aspect that entire craton appears engulfed by surrounding rift valleys, thrust zones and tectonically active mobile belts. It is marked by the circular morphostructure of large dimensions (Rantsman et al. 1995) (Fig. 3.2) and uplifted to the tune of almost half a kilometer than the surrounding region (Fig. 3.6). It suffered due to its close proximity to the Crozet/Kerguelen hotspots, which were active around 115 Ma. Verma and Prasad (1975) reported rare occurrence of native iron in the Newer dolerite group of dyke swarms (shown in Fig. 3.2), source of which can either be extra-terrestrial or deeper mantle. All these characteristics support a strong crust-mantle interaction due to a deep mantle plume, and its transient effects still exist, if the deciphered thermal signatures are taken into account (Fig. 3.7). Possibly this craton has subsequently turned into a mobile belt, after its inception in Paleoarchean or even earlier. Currently, this craton hardly possess any geophysical characteristic of a typical craton.

3.2 Bastar Craton 3.2.1 Geological Settings The Bastar craton, also popularly known as Bastar-Bhandara craton, covers an area of about 130,000 km2 in the eastern part of the Indian shield (Fig. 3.8). It is surrounded by active rifts and Proterozoic mobile belts from all sides. It is bounded by the CITZ (Central India Tectonic Zone), forming part of the Satpura mobile belt in the northwest, Mahanadi graben in the northeast and Godavari graben in the southwest, whereas its SE boundary is demarcated by a shear zone associated with the Eastern Ghats Belt (EGB). This craton is essentially made up of orthogneisses with enclaves of amphibolites, Paleo-Mesoarchean banded TTG gneisses, and low to high-grade metasediments. The gneisses/migmatites and amphibolites are considered

3.2 Bastar Craton

105

Fig. 3.8 Geological set up of the Bastar craton. Location of heat flow sites are shown by 1 (Moggara, heat flow: 51.5 mW/m2 ), 2 (Bodal, heat flow: 63.5 mW/m2 ) and 3 (Malanjkhand, heat flow: 52 mW/m2 ). Modified after Fig. 1, Gupta et al. (1993)

early crustal components of this craton. Gneisses of the Bastar terrain are of various types and ages, varying from Mesoarchean to Paleoproterozoic (e.g. Sharma 2009). The tonalite gneisses located to the east and west of Kotri-Dongargarh linear belt, are dated 3562–3580 Ma (Ghosh 2004), which are comparable to one of the oldest rocks found in other Indian cratons. Ancient supracrustal of this region contain quartzitecarbonate pelite, banded iron formation and minor portion of mafic-ultramafic rocks, including felsic and mafic volcanics, and Neoarchean-Paleoproterozoic intrusive granites. Further, the basement gneiss complex in the craton contains granulite facies rocks. Two such prominent belts include (i) Bhopalpatnam granulite belt, lying at the boundary demarcating Bastar craton and Godavari graben (Ramakrishna and Vaidyanadhan 2008), dated at around 1.6–1.9 Ga (Santosh et al. 2004), and (ii) 2.6 Ga Kondagaon granulite belt, occurring in the middle part of the craton. Such occurrences at surface level would indicate massive uplift and erosion in the Bastar cratonic terrain, that brought mafic crust close to the surface, which are well reflected in heat flow and gravity field, as discussed later. Supracrustals of this region are further overlain by many unmetamorphosed Proterozoic basins like Neoproterozoic Chattisgarh basin, which is a prominent feature of this craton. Mafic dyke swarms are also reported from

106

3 Singhbhum and Bastar Cratons

this region (Srivastava et al. 1996), including late Cretaceous kimberlites (Lehmann et al. 2010) and almost similar aged mafic dyke intrusion in Chhattisgarh Basin (Chalapathi Rao et al. 2011). Geology of this region has been discussed in great detail by Naqui and Rogers (1987), Ramakrishnan and Vidyanadhan (2008) and Sharma (2009).

3.2.2 Geophysical Studies Similar to Singhbhum region in the north, this craton too appear to be squeezed from all sides either by active rift valleys or mobile belts, thereby making it extremely inhomogeneous. Geophysically however, it remains least studied, compared to all the other cratons of the Indian shield. There are few works on gravity studies, but no DSS or MT study is attempted yet. Whatever limited information now available, is discussed below.

3.2.2.1

Crustal Seismic Structure

Deep Seismic Sounding (DSS) Studies No DSS study has been directly undertaken over this region, However, the Hirapur Mandla profile (discussed in Chap. 5), which runs across the Son valley Vindhyans, passes through the Jabalpur-Mandla segment of the Bastar craton (Murty et al. 2004). Crust beneath this segment is made up of mainly four layers, below Deccan volcanics. The granitic upper crust (Vp : 5.9 km/s) beneath Deccan traps extends down to 7 km depth, which is underlain by 15 km thick middle crust extending down to 22 km depth. This second layer can be divided into two parts, 9 km thick typical middle crust (Vp : 6.5 km/s), followed by 6 km thick prominent zone of low velocity (6.35 km/s), which may be metasomatised. Below this metasomatised layer, lies a 22 km thick typical lower crust above the Moho. This layer is characterised by velocity of 6.8 km/s. Murty et al. (2008) reanalysed this seismic section and found a little higher velocity (6.96 km/s) for this layer. The Moho is found at the depth of 44 km. The Receiver function study, however, reported the Moho discontinuity at 52.5–65.0 km depth. No magma underplating is reported from this region, which is surprising, as all other crustal sections in Vindhyan basin as well as Singhbhum craton, are associated with thick magma underplating above the Moho. The crustal structure derived by the DSS studies (wide angle reflection data) may be more reliable than the receiver function studies, as the broadband stations were located close to the boundary of the craton, rather than in the central part of the region. Receiver Function Studies Teleseismic data of eight broadband sites located mostly close to the periphery of this craton, were used for receiver function studies; no broadband station was located

3.2 Bastar Craton

107

Fig. 3.9 A plot between averaged Vs and computed Vp (using Vp /Vs : 1.75) distribution with depth beneath Bastar craton, using Vs data from receiver function studies at various locations by Julia et al. (2009). For comparison, similar plots have been added for south Indian shield (Pandey et al. 2013), as well as global shields and platform (Christensen and Mooney 1995). Inferred crustal compositional structure is also added

in the central part of the craton (Julia et al. 2009). Figure 3.9 shows average Vs and computed Vp (using Vp /Vs : 1.75) distribution with depth. For comparison, similar plots representing the south Indian shield (Pandey et al. 2013) as well as global shields and platform (Christensen and Mooney 1995) are also shown. The estimated crustal Vs as well as Vp below Bastar craton, are considerably lower compared to that below Singhbhum craton and south Indian shield. The results, however, match well with the average global shields and platforms. This indicates that the upper and middle crust beneath the Bastar craton, may be highly felsic compared to mafic nature in Sighbhum craton as well as Dharwar craton. This is surprising, as lots of intrusive bodies, as well as some granulitic exposures are also present over here. The derived crustal structure indicates that the upper crust is about 12.5 km thick, followed by 5.0 km thick middle crust, which is underlain by a much thicker (35.0 km) lower crust. The bottom-most 20 km thick lower crust resembles to underplated magma. No mid-crustal metasomatic zone is found here, suggesting that the underplated magma was devoid of mantle volatiles and halogens.

3.2.2.2

Gravity Studies

Gravimetrically, this cratonic segment has not been studied well. Subrahmanyam and Verma (1986), while studying the Eastern Ghats Belt in detail, covered a part of Bastar craton between Lat. 18–22° N and Long. 80–84° E, that included the Chattisgarh basin, granitic terrain, green stone belt and Dongargarh group of rocks (Fig. 3.10). They found a clear distinction between the Bastar cratonic segment

108

3 Singhbhum and Bastar Cratons

Fig. 3.10 Bouguer gravity anomaly map (in mGal) of part of the Bastar craton (BC) and adjoining Eastern Ghats Belt region (EGB). GG: Godavari graben. Modified after Fig. 4, Subrahmanyam and Verma (1986)

and the Eastern Ghats Belt, the boundary between the two was demarcated by high gradient in gravity field, suggesting it to be a faulted zone. The Eastern ghats Belts (EGB) lying east of the Bastar craton, may be a block-uplifted one, representing deeper level crust. Consequently, the gravity field is positive, as it is made up of high grade metamorphic rocks like pyroxene granulites, gneisses, charnockites and gabbro-anorthosite masses. Several high gravity closers can be seen over this terrain, which are associated mainly with basic charnockites. In contrast, gravity anomalies are negative reaching as low as −90 mGal over the Bastar craton. The Chattisgarh basin, which is one of the prominent Proterozoic basins of the country, is associated with Bouguer gravity anomalies from −30 to −70 mGal.

3.2.2.3

Magnetic Studies

Anand and Rajaram (2003) made an attempt to study part of the Eastern Ghats Belt (EGB) and Bastar craton, using aeromagnetic data. They found that the Bastar block is dominated by NS trending and some NW-SE trending structural features, which terminate near the boundary of the EGB. A much detailed investigation was carried out by Murthy and Babu (2009), who studied aeromagnetic anomaly map over this craton and the adjacent Parnhita Godavari basin on its south. The anomaly map, contoured at an interval of 20 gammas, is reproduced in Fig. 3.11, which

3.2 Bastar Craton

109

Fig. 3.11 Aeromagnetic anomalies (in gammas at an attitude of 1500m) over Bastar craton and a part of Pranhita-Godavari basin. Modified after Fig. 2, Murthy and Babu (2009)

indicates a large variation from −200 to +200 gammas, although anomalies of larger magnitude have also been found at some isolated places. In general, this craton is occupied by positive magnetic anomalies, except a long patch of E–W trending negative anomaly in the northern part of the craton, at the latitude 18° 45 N. Positive magnetic anomalies may be correlated with the NW–SE striking mafic intrusives, located at various depths, which apparently intruded in a tensile regime created by the collison of Bastar and Dharwar craton during Paleoproterozoic. Two prominent anomalies over the Bastar craton are modelled by Murthy and Babu (2009). The first one (anomaly A in Fig. 3.11) situated southeast of Jagarkonda, has a length of about 25 km and a magnitude of +160 gammas. This anomaly is attributed to a 16 km wide dyke, located at 4.0 km depth. The other anomaly (anomaly B in Fig. 3.11), which is 70 km in length and has a magnitude of about +120 gammas, is situated southeast of Dantewara in the northeastern part of the craton. It is also interpreted as 28 km wide dyke, but located at deeper depth 11.3 km.

3.2.2.4

Heat Flow and Lithosphere Studies

Gupta et al. (1993) reported surface heat flow measurements from seven boreholes, from three different geological areas of the central Bastar craton, which include (i) Bodal located over the Bijli rhyolite and Pitapani volcanics, (ii) Malanjkhand, in Paleoproterozoic granitic terrain (Sikka 1989), and (iii) Mogarra in unclassified granites and gneisses. Heat flows from these locations vary from 51.5 to 63.5 mW/m2

110

3 Singhbhum and Bastar Cratons

(Fig. 3.8) with a mean of 55.7 ± 5.5 mW/m2 . The temperature gradients varied from 13.5 to 23.8 °C/km (Gupta et al. 1993), reflecting the changes in in situ lithology. Surface radioactive heat generation was measured at 2.97 µW/m3 for the Dongargarh granites of Bodal, while a heat generation of 2.1 µW/m3 was estimated for Malajkhand granodiorite (Gupta et al. 1993). Lithospheric Thickness Based on four-layer crustal heat production model deciphered from deep crustal seismic studies, and taking into account the measured heat production and heat flow (Table 3.4), temperature-depth calculations are made for two locations, Bodal and Malajkhand (Fig. 3.7). Estimated Moho temperatures and heat flow input from the mantle are found to be 630 °C and 30 mW/m2 for Bodal and 530 °C and 24 mW/m2 for Malajkhand. For these locations, lithospheric thickness is found to be about 100 and 135 km respectively. The lithosphere below Bastar craton appear much thicker than that of the adjacent Singhbhum craton.

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Shalivahan, Bhattacharya BB, Chalapathi Rao NV, Maurya VP (2014) Thin lithosphere–asthenosphere boundary beneath Eastern Indian craton. Tectonophysics 612–613:128–133 Sharma M, Basu AR, Ray SL (1994) Sm-Nd isotopic and geochemical study of the Archaean tonalite—amphibolite association from the eastern Indian craton. Contrib Mineral Petrol 117:45– 55 Sharma RS (2009) Cratons and fold belts of India. Springer, Berlin, 304 pp Sikka D (1989) Malanjkhand proterozoic porphyry copper deposit, M.P., India. J Geol Soc India 34:487–504 Singh AP, Kumar N, Zeyen H (2015) Three-dimensional lithospheric mapping of the eastern Indian shield: a multi-parametric inversion approach. Tectonophysics 665:164–176 Singh MR, Manikyamba C, Ganguly S, Ray J, Santosh M, Singh TD, Kumar BC (2017) Paleoproterozoic arc basalt-boninite-high magnesian andesite-Nb enriched basalt association from the Malangtoli volcanic suite, Singhbhum Craton, eastern India: geochemical record for subduction initiation to arc maturation continuum. J Asian Earth Sci 134:191–206 Srivastava RK, Hall RP, Verma R, Singh RK (1996) Contrasting Precambrian mafic dykes of Bastar craton, central india: Petrological and geochemical characteristics. J Geol Soc India 48:357–546 Subrahmanyam C, Verma RK (1986) Gravity field, structure and tectonics of the Eastern Ghats. Tectonophysics 126:195–212 Tripathi P, Parthasarathy G, Ahmad SM, Pandey OP (2012) Mantle derived fluids in the basement of the Deccan traps: evidence from stable carbon and oxygen isotopes of carbonates from the Killari borehole basement, Maharashtra, India. Int J Earth Sci 101:1385–1395 Vedanti N, Malkoti A, Pandey OP, Shrivastava JP (2018) Ultrasonic P- and S- wave attenuation and petrophysical properties of Deccan flood basalts, India as revealed by borehole studies. Pure Appl Geophys 175:2905–2930. https://doi.org/10.1007/s00024-018-1817-x Varma OP (1986) Some aspects of ultramafic and ultrabasic rocks and related chromite metallogenesis with examples from eastern region of India. In: Proceedings of the seventy-third session Indian Science Congress Association, Delhi, pp 1–72 Verma RK, Prasad SN (1975) Probable existence of native iron in newer dolerites from Singhbhum, Bihar, India. J Geophys Res 80:3755–3756 Verma RK, Mukhopadhyay M, Roy SK, Sinha RPP (1978) An analysis of gravity field over North Singhbhum. Tectonophysics 44:41–63 Verma RK, Sarma AUS, Mukhopadhyay M (1984) Gravity field over Singhbhum, its relationship to geology and tectonic history. Tectonophysics 106:87–107 Verma RK, Subba Rao PBV, Mukhopadhyay M (1996) Analysis of gravity field over Singhbhumregion and evolution of Singhbhum craton. Recent Res Geol 16:123–133 Verma RK, Rao RUM, Gupta ML (1966) Terrestrial heat flow in Mosabani mine, Singhbhum district, Bihar, India. J Geophys Res 71:4943–4948

Chapter 4

Aravalli and Bundelkhand Cratons

4.1 Aravalli Craton 4.1.1 Introduction Aravalli craton of Mesoarchean-Paleoproterozoic age, situated in northwestern part of the Indian shield (Fig. 4.1), is one of the complex geotectonic segments of the Indian land mass, apart from the earlier discussed Singhbhum craton. It strikes in NNE-SSW and extends from north of the Cambay graben to Delhi and covers an area of almost 100,000 km2 in the states of Rajasthan, Gujarat, Delhi, Haryana and Madhya Pradesh (Naqui et al. 1974; Naqui and Rogers 1987; Sinha-Roy et al. 1995; Ramakrishnan and Vaidhyanadhan 2008; Sharma 2009). This craton is bounded on its eastern side by the Chambal Valley Vindhyans and Great Boundary Fault, which separates Aravalli-Delhi Fold Belt (ADFB) from the adjacent Bundelkhand craton. Down south, it extends up to Son-Narmada-Tapti rift zone and in the north, up to the Indo-Gangatic plain. This region was largely formed at the close of the Dharwar era, but got rejuvenated several times since then. It witnessed four major tectonomagmatic and metamorphic events: (i) The Bhilwara gneissic complex ~3.0 Ga, (ii) Aravalli orogeny ~1800 Ma, (iii) The Delhi orogeny ~1100 Ma and (iv) The post Delhi magmatic event ~873–800 Ma (Ramakrishnan and Vaidyanadhan 2008; Sharma 2009). The evolutionary history of this region has always remain controversial. Based on recent geophysical evidences, it has been, however, suggested that the collision between the Bundelkhand and Mewar cratons was primarily responsible for the development of the Aravalli mountain ranges (Mishra and Ravi Kumar 2014).

4.1.2 Regional Geology Major part of the Aravalli craton can be sub-divided into five geotectonic units: (i) Archean Mewar gneisses, granitoids and supracrustals, (ii) Meso-Neoarchean © Springer Nature Switzerland AG 2020 O. P. Pandey, Geodynamic Evolution of the Indian Shield: Geophysical Aspects, Society of Earth Scientists Series, https://doi.org/10.1007/978-3-030-40597-7_4

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Fig. 4.1 Geological map of the Aravalli–Delhi Fold Belt region along with the locations of NagaurKunjer geotransect (with shot points) and Kekri-Kota MT profile. After Fig. 1, Vijaya Rao et al. (2000)

Bhilwara Supergroup, (iii) Paleoproterozoic Aravalli Fold Belt, (iv) Mesoproterozoic Delhi Fold Belt, and (v) Neoproterozoic basins and Malani Igneous Suite. The oldest unit, Archean Mewar TTG gneisses and granitoids, consists mainly enclaves of supracrustal rocks like, amphibolite, quartzite, banded iron formation and mica schist, while the 3.3 Ga Bhilwara Supergroup complex, forms the basement of most of the Proterozoic successions (Sinha-Roy et al. 1995). It comprises three tectono-stratigraphic units, Hindoli Group, Mangalwar Complex and Sandmata Complex. Out of these, the Hindoli group consists of green schist to amphibolite facies greywacke and phyllites with minor quartzite, dolomite and limestone and some volcanic sequences, while Mangalwar Complex contains metavolcanic and metasedimentary rocks, which are considered as an older granite-greenstone

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belt, characterised by amphibolite facies metamorphism (Sugden et al. 1990; SinhaRoy et al. 1995). The Sandmata complex, consists of granulite facies rocks (largely charnockites), which occur as isolated bodies, exhumed from a much deeper level. They are understood to have been formed at a high temperature of 650–850 °C and 8–11 kb pressure (Tewari et al. 2018). U-Pb zircon dating of these granulites yielded an age of 1730 Ma (Sharma 1988). Apart from granulites, amphibolites, amphibolite facies migmatitic gneisses, granitoids are also found in this belt. Towards the west, this complex is in thrust contact with Delhi Fold Belt. The Paleoproterozoic Aravalli Fold Belt, a prominent litho-tectonic unit of this craton, consists primarily of quartzites, volcanics, conglomerate, greywacke, phyllites, dolomites and schists. These rocks have undergone amphibolite to greenschist facies metamorphism. On the other hand, the Delhi fold belt, which is a 450 km long linear belt, exposes the rocks of Delhi super group of 1.7–1.5 Ga (Mckenzie et al. 2013). This fold belt can be divided into two parts, the north Delhi Fold Belt and the South Delhi Fold Belt. North Delhi Fold Belt comprises conglomerates, quartzites, basic volcanics, phyllites and granitic bodies, apart from amphibolites and carbonites, while the South Delhi Fold Belt is made up of metabasalts, schist, metapalites, quartzite, ophiolites and granitic intrusions. Neoproterozoic basins mainly contain deformed and metamorphosed sediments, while the Malani igneous suite, which is a collective term for bimodal volcanic and plutonic rocks, is dated between 873–800 Ma (Meert and Pandit 2015). This igneous suite represents one of the largest felsic volcanic eruption in the world, covering an area of almost 51,000 km2 . The volcanics are of mainly rhyolite, perlite and trachyte type, which are intimately associated with volcano-clastics. Some granitic plutons and dyke swarms are also present in this suite. Detailed geological history of this craton can be found in Roy (1988), Sinha-Roy and Gupta (1995), Gupta et al. (1997), Sinha-Roy et al. (1998), Paliwal (1999), Kataria (1999), Roy and Jakhar (2002), Ramakrishnan and Vaidyanadhan (2008) and Sharma (2009).

4.1.3 Geophysical Studies As mentioned earlier, this craton forms one the most complicated geologic segment of the Indian peninsular shield, having undergone paleo-subduction, continentcontinent collision, multiple folding and active deformation. Considerable amount of geophysical work has thus been carried out over this terrain, to understand its deeper geology, which are summarized by many workers like, Gupta et al. (1967), Reddy and Ramakrishna (1988), Sinha-Roy et al. (1998), Tewari et al. (1998, 2018), Tewari and Vijaya Rao (2003), Roy and Rao (2000) and Mandal et al. (2018). Geophysically, this region is characterised by (i) a high order positive gravity anomaly (unlike most of India, marked by negative gravity field), cause of which are still not clearly understood, (ii) relatively higher heat flow values than the cratonic shields, specially in the Delhi system, and (iii) significant seismotectonic activity (Khattri et al. 1984).

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Crustal Seismic Structure

Deep Seismic Sounding Studies Multi-fold deep seismic reflection data was acquired by CSIR-National Geophysical Research Institute (Hyderabad), along a 436 km–long profile, from Nagaur to Kunjer (Fig. 4.1) (Tewari et al. 1995; Reddy et al. 1995; Rajendra Prasad et al. 1998). This profile passed through all the major geological units of the Aravalli-Delhi Fold Belt region that included (i) Marwar basin situated in the western part of the craton, (ii) Delhi fold belt, (iii) Sandmata granulitic complex, (iv) Mangalwar complex, (v) Hindoli Group of rocks, and (vi) Great Boundary Fault. This profile also covered Bundi-Kota and Kethun-Azadpura-Kunjer sectors in the Chambal Vindhyan basin. The acquired data were initially interpreted by Reddy et al. (1995) and Tewari et al. (1995) and further followed by many workers like, Tewari et al. (1997a), Satyavani et al. (2001), Rajendra Prasad et al. (1998), Krishna and Vijaya Rao (2011), Vijaya Rao and Krishna (2013) and Mandal et al. (2014, 2018), but no coherent model emerged for the entire length of the profile, except a few velocity-depth models. (i) Nagaur-Masuda sector: SP 0-SP 1700 This 150 km long section that covers Marwar basin and South Delhi Fold Belt, has been studied in detail by Krisna and Vijaya Rao (2011) and Vijaya Rao and Krishna (2013). They delineated a six layer crust below the Marwar Terrain (Fig. 4.2). The first couple of kilometre thick layer above the crystalline basement, consisted Tertiary sediments, Marwar sediments and Malani volcanics, characterised by Vp 2.1 km/s, 4.4 km/s and 4.8 km/s respectively. These formations rest directly over a 8 km thick granitic-gneissic upper crust (Vp 5.9–6.1 km/s). This layer is underlain by a 10 km thick middle crust (Vp 6.1–6.3 km/s), followed by a 4 km thick metasomatised low velocity layer (Vp 5.9 km/s) between 20 and 24 km depth. Below this layer lies a 12 km thick typical granulitic lower crust (Vp 6.7 km/s) that continues till 36 km depth. This layer is further underlain by a 4 km thick magma layer (Vp 7.3 km/s) above the Moho, which is at the depth of 40 km. The mantle is characterised by a typical ultramafic velocity of 8.1 km/s. The crust is made up of mainly four layers below the South Delhi Fold belt falling between Alniyawas (near Rian) and Masuda. A 12 km thick upper crust (Vp 5.9 km/s) is followed by a 6 km thick middle crust (Vp 6.2 km/s), which is further underlain by the lower crust down to 29 km depth. The bottom of the lower crust is stacked by an unprecedently thick (about 18 km) magma layer above the Moho at 47 km depth. This underplated magmatic layer is characterised by velocity 6.9–7.5 km/s (mean velocity: 7.3 km/s), which would mean that the Moho below this region is deeper by 7 km compared to that found below the Marwar terrain. A prominent steeply southeast dipping discontinuous bands of isolated reflections (shown by arrows in Fig. 4.2) are also observed below Marwar basin, which coincide with the western margin of the Delhi fold belt. These reflection bands are much wider than the Jahazpur thrust observed further east on the geotransact. They even appear

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Fig. 4.2 2-D seismic structure of the crust and the underlying uppermost mantle below NagaurMasuda segment, modeled by the wide angle recordings (Krishna and Vijaya Rao 2011; Vijaya Rao and Krishna 2013). Disposition of different crustal layers are also shown. Velocities are in km/s. LVZ refers to low velocity zone. SDFB denotes the location of South Delhi Fold Belt, which can be interpreted as a volcanic front, related to paleo-subduction. Arrows indicate areas of isolated southeast steeply dipping reflection bands. SP refers to shot points

to extend below the Moho. Such reflectors were also observed in the migrated nearoffset reflection images studied by Krishna and Vijaya Rao (2011). These reflectors may be correlated with the remnants of erstwhile subducting Marwar craton, that took place during Mesoproterozoic collision episode related to Delhi orogeny. This was the time, when Rodinia supercontinents were getting evolved. Inspite of intense deformation around that time, signature of collisional suture are found below Rian (Fig. 4.3). This may have been formed due to collision by two subducting slabs, Marwar craton from the west and Mewar craton from the east, which resulted into formation of Delhi Fold Belt. Similar inference are drawn by Sychanthavong and Desai (1977) and Deb and Sarkar (1990) also. (ii) Masuda-Nandsi-Deoli-Jahajpur-Bundi Sector: (SP1700-SP 3615) This 140 km long stretch consists of Sandmata granulitic and Mangalwar Complexes besides Hindoli Group of rocks. All belong to Bhilwara Super Group Complex. The region is bounded by the Great Boundary Fault to the east. The Sandmata complex is poorly reflective seismically, in comparison to Mangalwar and Hindoli groups, which are in contrast, highly reflective (Tewari et al. 1997b). The Great Boundary Thrust (GBT) zone is well identified below Hindoli region (Mandal et al. 2018), and the Jahazpur Thrust Zone (JT), below Mangalwar complex (Fig. 4.3). This would mean that all these three sectors have continuously getting evolved throughout the geological period, as they have undergone multiple geotectonic upheavals. Satyavani et al. (2004) provided velocity–depth section for the Sandmata granulitic complex, together with Marwar basin and Delhi Fold Belt. Below Sandmata complex, they reported presence of a six layer crust, with Vp 4.6 km/s (surface

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4 Aravalli and Bundelkhand Cratons

Fig. 4.3 Geodynamic and geologic interpretation of the 2-D crustal seismic structure along NagaurKunjer geotransect, covering various segments of the Aravalli-Delhi Fold Belt and adjacent Chambal Valley Vindhyan region based on Tewari et al. (1995) and Reddy et al. (1995). Velocities (in km/s) are from Krishna and Vijaya Rao (2011) for Marwar basin and South Delhi Fold Belt and Satyavani et al. (2004) for Sandmata complex. Shaded area indicates underplated magmatic region. GBF: Grest Boundary Fault, JT: Jahazpur Thrust Zone, GBT: Grest Boundary Thrust, and CT: Chambal Thrust

to 2 km), 5.5 km/s (2–12 km), 6.2 km/s (12–20 km), 6.5 km/s (20–28 km) and 6.8 km/s (28–35 km) (Fig. 4.3). These layers are underlain by a 10 km thick magma layer (7.3 km/s) and the Moho between 42 and 45 km. Mandal et al. (2014) on the other hand identified the Moho at a much deeper depth ~50 km below the Sandmata complex. Deeper Moho below this region was also contemplated by gravity studies (Rajendra Prasad et al. 1998; Vijaya Rao and Krishna 2013). Further, a conspicuous domal-shaped structure found by gravity studies below Delhi Fold Belt, appears to continue below Sandmata complex also (Fig. 4.3). In the Mangalwar-Hingoli sector, which is a highly reflective zone, a spectacular crustal scale thrust named Jahazpur Thrust (JT) as mentioned above, is seismically mapped; it dips at about 30° towards west (Mandal et al. 2014; Vijaya Rao et al. 2000). This thrust zone begins from below SP 2800 near Jahazpur and continues to SP 1700 below the Sandmata Complex. This thrust zone extends into the deeper parts of the lower crust and possibly beyond. Its high reflective nature could be related to its composite fault zone structure. Further, this thrust zone is about 80 km in length and about 25 km in width, characterised by parallel dipping refletions. It may thus represent the remnants of a paleo-subduction. The NW dipping lower crustal reflections observed between 12.0 and 16.0 s twt below Sandmata terrain, can be considered further extension of the JT. This thrust zone may have played an important role in the development of Mangalwar and Sandmata complexes. Presence of such structures are in conformity with the geologically complex nature of the Aravalli craton. A similar northwest dipping thrust zone, named Great Boundary Thrust (GBT) as mentioned before, is delineated below the Hindoli Group of rocks (Mandal et al. 2018). This thrust zone is characterised by wide reflection bands. This structure is about 20 km thick and extends laterally for about 35 km, below which Moho is at ~43 km depth. The top of this crustal-scale thrust fault lies close to surface just east

4.1 Aravalli Craton

121

of Bundi and at about 15 km depth below the middle part of the Hindoli Group terrain. Adjacent to the GBT one more thrust zone named Chambal thrust (CT) is mapped below the Vindhyan sediments (Fig. 4.3) in the Bundi-Kota segment of the geotransect. All the three thrust zones (JT, GBT and CT) (Fig. 4.3), may have originated due to compressive forces related to subduction and collision episodes between Bundelkhand and Mewar cratons during Paleoproterozoic. Another important finding of the seismic studies has been the detection of lower crustal magma underplating at an average depth of about 35–45 km, below all the three complexes, Sandmata, Mangalwar and Hindoli (Fig. 4.3.). These complexes exhibit all the characteristics of layered and laminated zones. Immediately east of the JT, lies the Great Boundary Fault that separates the Aravallis from the Chambal Valley Vindhyan basin. The nature of the crustal seismic profile representing Chambal valley Vindhyans sector, has been dealt in Chap. 5 on ‘Vindhyan basins’ Receiver Function Studies In comparison to southern parts of the Indian shield, northern parts, specially the Aravalli and Bundelkhand cratons, are not well studied by broadband seismics. In Aravalli craton, receiver function studies are carried out at three stations, but the reliable data are available only for two stations, NDI and GRG. The station details are given in Table 4.1 and S- and P-wave velocity variation with depth is shown in Fig. 4.4. In northern part of the craton, Moho is found at much deeper depth (52.5– 55 km), compared to those reported by seismic reflection studies as discussed in the above section. Deeper Moho depth is attributed to thick (15–17.5 km) magma underplating below the lower crust. The crust-mantle thermal interaction and subcrustal erosion is caused by shallowing of the LAB to about 90–100 km (Kumar et al. 2013). Further, P- and S-wave velocities in the mid to lower crust, are much lower below the NDI seismic station compared to the GRG station (Fig. 4.4a, b). Table 4.1 Crustal seismic structure of Aravalli and Bundelkhand cratonic blocks based on receiver function studies by Julia et al. (2009) Station

Lat. Long. (in degrees)

Upper crust (km)

Middle crust (km)

Lower crust (km)

Underplated magmatic lower crust (km)

Depth to Moho (km)

Aravalli block GRG

28.31, 76.94

10.0

10.0

15.0

17.5

52.5

NDI

28.68, 77.22

0.0

35.0

5.0

15.0

55.0

Bundelkhand block JHN

25.51, 78.54

12.5

7.5

17.5

–

37.5

BRT

27.08, 77.39

12.5

5.0

17.5

2.5

37.5

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4 Aravalli and Bundelkhand Cratons

Fig. 4.4 a, b Vs and computed Vp distribution with depth beneath GRG and NDI seismic stations (Table 4.1) located in North Aravalli craton based on Julia et al. (2009). For comparison, similar plot has been added for global shields and platform (Christensen and Mooney 1995). c, d Averaged Vs and computed Vp distribution below Bundelkhand craton, based on studies by Julia et al. (2009) for seismic stations JHN and BRT (Table 4.1). Inferred crustal compositional structure below this craton is also shown. For computing Vp, Vp/Vs ratio is taken as 1.75

This indicates possibility of large scale metasomatic alteration in the mafic crust. The estimated velocities are in general on lower side than observed in other parts of Indian shields as well as global shields and platform, specially below NDI station (Fig. 4.4b). Over the Aravalli region, both the filtered geoid as well as filtered free air anomalies, are reasonably higher at about 5 m and +40 to +60 mGal respectively, compared to much lower values in adjacent regions (Kumar et al. 2013). This would conform to shallowing of the LAB, but the observed low crustal velocities by receiver function is unusual. Usually crustal velocities are expected to be higher in high gravity areas.

4.1 Aravalli Craton

4.1.3.2

123

Gravity Field

Regional inhomogeneous nature of this region, is well reflected in the gravity map, which has been studied in detail by various workers (e.g. Verma et al. 1986; Reddi and Ramakrishna 1988; Rajendra Prasad et al. 1998; Mishra et al. 1995, 2000; Vijaya Rao et al. 2000; Mishra and Ravi Kumar 2014). The regional gravity high all along the strike of Aravalli in the Bouguer, Free-air and isostatic gravity anomaly maps, may be marked as the one of the most anomalous regions in the peninsular shield (Verma et al. 1986). The Bouguer gravity anomaly map of the Aravalli region and an interpreted crustal section along the Nagaur-Kunjer profile is reproduced in Figs. 4.5 and 4.6 respectively. Figure 4.5, which is mainly based on the studies of Reddi and Ramakrishna (1988) and Mishra et al. (1995), reveals Bouguer anomaly of −40 mGal near SP1, drops down to −60 mGal near SP 400 and thereafter increases steeply to about zero mGal near SP 1029 and continues till SP 1900. The gravity anomaly drops down further to about −70 mGal near Jahazpur thrust zone, and −40 to −60 mGal over the Chambal valley Vindhyans. All along the 700-km long Aravalli Delhi Fold Belt, the NE-SW gravity anomaly is high (about 0 to +10 mGal). It has a width around 100 km, and spreads between SP 1029 and SP 2200 of the Nagaur-Kunjer seismic profile, coinciding with the Delhi Fold Belt and Sandmata complex. The anomaly is demarcated by ‘H’ in Fig. 4.5. Since, this gravity high feature is of considerable importance in understanding the geodynamic evolution of the Aravalli-Delhi Fold Belt, Bansal and Dimri (1999) carried out forward gravity modelling over the mid crustal domal structure, detected by the seismics. They found that this feature is 25 km in height and has a density of 3.04 g/cm3 . It is located at a depth of about 14 km down to the Moho boundary at 39 km. Reddi and Ramakrishna (1988), on the other hand, related this gravity high to a horst structure with bounding faults, extending well into the mantle. Apart from the horst structure, it is felt that this gravity high can also be caused partly by the upwarping of high-grade mafic granulites at shallower depths due to massive plateau uplift beneath the Delhi fold belt and parts of Bhilwara gneissic region. It can further be attributed to thick magma underplating and subcrustal erosion underneath, consequent to rise of isotherms and shallowing of the LAB. This belt of gravity high is flanked on either side by high order gravity lows (L), across which, gravity gradients are quite sharp indicating faulted margin on either side.

4.1.3.3

Total Magnetic Intensity

Total magnetic intensity profile along the Nagaur-Kunjer section (Mishra et al. 1995), is shown in Fig. 4.7, which matches fairly well with the major faults and geologic boundaries. The magnetic intensity is high over the Delhi fold belt-Sandmata complex and adjacent regions. For example, contact between Sandmata and Mangalwar complex is well marked by a deep seated fault anomaly. Fault zone associated with the Jahazpur mega-thrust zone can also be clearly seen. Similarly, Delhi fold belt that falls between Alniyawas and Masuda, appears to be bounded on either sides by

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4 Aravalli and Bundelkhand Cratons

Fig. 4.5 Bouguer gravity anomaly map of the NW Indian Shield. Location of Nagaur-Kunjer seismic reflection profile is also shown. After Fig. 1, Mishra et al. (1995)

two major faults. Magnetic intensity is much lower over the Marwar basin on the western side of the profile and over the Vindhyan basin east of Great boundary fault.

4.1.3.4

Geoelectric Section

Gokarn et al. (1995) acquired broadband magnetotelluric (MT) data in the frequency range 0.01 ± 100 Hz and at a station spacing of 10 km, along a 120 km profile

4.1 Aravalli Craton

125

Fig. 4.6 Observed and computed Bouguer gravity anomalies together with two-dimensional crustal density model along Nagaur-Kunjer seismic profile. JT refers to Jahazpur Thrust. Modified after Fig. 3, Rajendra Prasad et al. (1998)

Fig. 4.7 Total magnetic intensity profile from Nagaur to Jhalawar (Kunjer). After Fig. 3, Mishra et al. (1995)

between Kekri and Kota (location shown in Fig. 4.1). It also cuts across the Great Bóundary Fault (GBF) and Jahazpur Thust (JT) on its western side. The data were collected at 19 stations. This study indicated presence of about 3–5 km thick midcrustal conductor below all the MT stations at the depth range of 7–12 km. They also identified a northwest dipping conductor in the middle part of the crust, having a resistivity of about 50 m. However, it is disrupted near Jahazpur, where it is limited to shallower depth of 3 km, showing steep upward trend. The structural trend of the delineated conductor conforms well with the seismically imaged JT.

126

4.1.3.5

4 Aravalli and Bundelkhand Cratons

Magnetometer Array Study

A magnetometer array study was undertaken in 1979 and 1984, which led to discovery of a major conductivity structure striking across the western part of Ganga basin and extending into the foothills of Himalaya (Lilley et al. 1981; Arora et al. 1982; Srivastava et al. 1984; Vozoff 1984; Arora and Mahashabde 1987; Chamalaun et al. 1987). This conductor appears to be the northward continuation of the Sandmata– Mangalwar geotectonic blocks. The top of the conductor is slated to be between a few kilometers to 100 km; some other investigations, however, estimate it at only 15 km (Arora and Mahashabde 1987) and 32 km (Chamalaun et al. 1987). Although the origin of this conductor was reported to be due to upward intrusion of magma from the asthenosphere (Arora and Mahashabde 1987), or pressure-released partial melting during the collision of India with Asia (Chamalaun et al. 1987), but that does not look feasible, as the depth to the conductive body is too shallow to support melting conditions. It may be possible that the conductivity anomaly depicted by magnetometer array, reflects upwarped molten asthenosphere beneath the west coast high heat flow anomaly zone down south (Pandey and Agrawal 2000), as discussed in Chap. 6. Quite likely, the current induced in adjoining areas, may be channelled through the zone of high heat flow, which provide a passage to the anomalous flow of induced currents in the Aravalli conductor. This may be feasible since the lineaments in the southern part of Aravalli region are related to the formation of Cambay graben and Narmada-Son lineament (Gupta and Bharktya 1982).

4.1.3.6

Heat Flow and Lithosphere Structure

Regional inhomogeneity beneath Aravallis as well as adjoining regions of north western India, is well reflected in heat flow distribution too. Altogether, nine heat flow values are reported from the areas covered by Aravalli-Delhi Mobile Belt; the details are given in Table 9.1 (Chap. 9) and locations in Fig. 4.8. The heat flow values range widely from 46 to 96 mW/m2 . Heat flow is little lower in southern part (46.0–67.0 mW/m2 ; mean: 59.0 ± 7.4 mW/m2 ) but reasonably high in the north (59– 96 mW/m2 ; mean: 74.8 ± 13.5 mW/m2 . Highest heat flow of 96 mW/m2 is recorded in Tusham granitic region, located within the Trans-Aravalli igneous suite (Sunder et al. 1990). This is the highest heat flow value so far recorded from any Meso- to Neoproterozoic terrains of India. It is almost twice to that recorded in the adjacent Bundelkhand craton. The Tusham granites are young dated around 750 Ma, nearly similar to the age of Malani rhyolites. The granites of this region are also highly radioactive. Even then, mantle heat flow is expected to be much higher below the Aravalli region. In comparison to Aravalli region, heat flow is much lower, averaging only around 32 and 38 mW/m2 in western and eastern Dharwar cratons of the south

4.1 Aravalli Craton

127

Fig. 4.8 Heat flow (in mW/m2 ) distribution in north-central India covering Aravalli craton, Vindhyan Basin and areas located close by NSL with location of DSS profiles (Fig. 5.2, Chap. 5). Data source Annual Report (1969–1970), Gupta and Rao (1970), Verma and Gupta (1975), Rao et al. (1978), Rao and Rao (1980, 1983), Gupta (1981), Gupta and Gaur (1984), Gupta et al. (1988, 1993), Sunder et al. (1990), Roy and Rao (2000), Nagaraju et al. (2012, 2017) and Pandey et al. (2014). ADMB: Aravalli-Delhi Mobile Belt, CG: Cambay graben, DG: Damodar graben, NSL: Narmada-Son-Lineament. Revised after Fig. 15, Pandey et al. (2014)

Indian shield respectively, as discussed in Chap. 2. This would point out to the presence of a thermally anomalous upper mantle below the Aravallis. Lithospheric Thickness (a) Khetri (north Aravalli) In this part of the Aravalli craton, reliable heat flow-heat generation data are available only for Khetri, which is located over the Mesoproterozoic Delhi Super Group (fold belt) terrain, overlying Archean banded gneissic complex. It is surrounded in the east by Sandmata Complex, composed primarily of high-grade granulitic rocks. Deep seismic reflection profile across the Delhi Fold Belt (Krishna and Vijaya Rao 2011; Vijaya Rao and Krishna 2013) has revealed 12 km thick upper crust (Vp: 5.9 km/s), 6 km thick middle crust (Vp: 6.2 km/s), 11 km thick lower crust (Vp: 6.7 km/s), 18 km thick underplated magmatic crust (Vp: 7.3 km/s) and Moho at 47 km depth. Heat flow at Khetri is estimated at 74 mW/m2 (Rao et al. 1976) and upper crustal

128

4 Aravalli and Bundelkhand Cratons

heat generation of the basement phyllite is 2.19 µW/m3 , based on measurements 11 samples of phyllites (Rao et al. 1976). (b) Rajpur Dariba (south Aravalli) This area is situated just east of Cambay graben. The heat flow site is located over a metal prospect belonging to Paleoproterozoic Aravalli super Group. A deep seismic reflection profile across the Sandmata Complex in the north, suggests 12 km thick upper crust (Vp: 4.6–5.5 km/s), 16 km thick mid to lower crust (Vp: 6.2–6.5 km/s) (out of which first 8 km thick layer appear metasomatised and rest 8 km, normal granulitic crust), 7 km thick mafic granulites (Vp; 6.8 km/s) and then a 10 to 11 km thick underplated magmatic crust (Vp: 7.3 km/s) (Satyavani et al. 2004). Measured heat flow is 64 mW/m2 and the upper crustal heat generation is 2.40 µW/m3 based on 23 samples of quartz mica schists (Roy and Rao 2000). Details of crustal heat production models used for calculating deep-seated temperatures are given in Table 4.2 and estimated temperature-depth profiles are shown in Fig. 4.9. Below Khetri, the Moho temperature and mantle heat flow are quite high at about 850 °C and 41 mW/m2 respectively. Consistent with this estimate, the lithosphere below this region is only about 70 km thick, which is unusual but somewhat similar to 60 km thickness obtained below the Singhbhum craton. The Moho temperature and Mantle heat flow, on the other hand, are recorded to be lower at 630 °C and 28 mW/m2 respectively for the Rajpur Dariba area, located in the southern part of Aravalli belt. Here, the lithosphere is about 105 km thick. It can thus be inferred that the northern part of the Aravalli, where measured heat flow is higher, is much warmer compared to the southern part of the craton.

4.1.3.7

Seismicity

Seismic activity is quite common in highly deformed orogenic belts which are associated with active deformation. Thus, being located close to the Himalayan collision zone, Aravalli belt remains seismically active. This region has been repeatedly rejuvenated and neotectonism is fairly common (Fig. 1.21, Chap. 1). The seismicity is largely associated with northeast-trending faults in the basement (Khattri et al. 1984), which intersect the Himalayan mountain ranges in the north. This may be a plausible reason for high seismic activity in the northern part of the belt located around Delhi (Khattri et al. 1984; Ramalingeswara Rao 2000, Koulakov et al. 2018). This region witnessed at least three major earthquake events in the past like, the 1907 Mallani (M 5.0), 1956 Kurja (M 6.0) and 1991 Jaisalmer (M 5.6). High strain rate of about 1.29–1.85 × 10−9 per year is also reported for the region (Ramalingeswara Rao 2000).

4.2 Bundelkhand Craton

129

Table 4.2 Crustal heat production models and adopted geothermal parameters for temperaturedepth estimation beneath Aravalli and Bundelkhand cratons Depth range (km)

Rock type

Heat production (µW/m3 )

Thermal conductivity (W/m °C)

Khetri (Aravalli craton) Surface heat flow: 74.0 mW/m2 0–12.0

Phyllites

2.19

3.0

12.0–18.0

Amphibolite facies rocks

0.78

2.5

18.0–29.0

Mafic granulite

0.16

2.5

29.0–47.0

Magmatic crust

0.02

2.6

>47.0

Ultramafic mantle

0.01

3.0

Data source Rao et al. (1976), Ray et al. (2003), Krishna and Vijaya Rao (2011), Vijaya Rao and Krishna (2013) and Pandey et al. (2017) Rajpur Dariba (Aravalli craton) Surface heat flow: 64 mW/m2 0-12.0

Quartz mica schist

2.40

3.0

12.0-20.0

Metasomatised granulite

0.50

2.5

20-28.0

Granulite

0.20

2.5

28.0-35.0

Mafic granulite

0.16

2.5

35.0-46.0

Magmatic crust

0.02

2.6

> 46.0

Ultramafic mantle

0.01

3.0

Data source Roy and Rao (2000), Ray et al. (2003), Satyavani et al. (2004), Mandal et al. (2014) and Pandey et al. (2017) Bundelkhand craton Surface heat flow: 37 mW/m2 0–7.5

TTG gneisses and metavolcanics

0.95

3.0

7.5–15.0

Felsic granulite

0.50

2.5

15.0–20.0

Metasomatised LVZ rocks

0.30

2.5

20.0–41.0

Mafic granulite

0.16

2.50

>41.0

Ultramafic mantle

0.01

3.0

Data source Nagaraju et al. (2017), Ray et al. (2003), Murty et al. (2004, 2008) and Pandey et al. (2017)

4.2 Bundelkhand Craton The Bundelkhand craton in north-central India, is separated from the Aravalli craton by the Great boundary fault (GBF) and vast tract of sedimentary rocks belonging to Vindhyan Supergroup. The Bundelkhand Granite Massif (also known as Bundelkhand Granitoid Complex) is a prominent geotectonic unit of this craton. Many researchers consider this unit as a part of Aravalli craton (e.g. Sharma 2009). Physiographically, this cratonic block resembles as a semi-circular outcrop, covering an

130

4 Aravalli and Bundelkhand Cratons

Fig. 4.9 Estimated temperature-depth distribution beneath Khetri and Rajpura-Dariba regions of Aravalli craton and average Bundelkhand craton, based on crustal heat production models given in Table 4.2

approximate area of 26,000 km2 or even more (around 35,000 km2 ), if the area concealed below Indo-Gangetic alluvium is taken into account (Basu 1986; Ramakrishnan and Vaidyanadhan 2008). It primarily contains TTG gneisses, metamorphosed volcano-sedimentary sequences and syn- to post-tectonic granitoids aged between 3.5 and 2.4 Ga. These formations are further intruded by 1.7 Ga dolerite dyke swarms, related to the development of rift basins associated with Bijawar (or Mahakoshal Group) and Gwalior formations (Sharma 2009). A number of basic dykes of tholeiitic affinity also occur here. The dyke activity is similar to those recorded in the Aravalli craton. This cratonic massif also contains younger granitoids and felsic volcanics, emplaced around 2.5 Ga (Mondal 2003). Some intrusion of giant quartz reefs of large dimensions, which followed after granitic magmatism, are associated with extensive hydrothermal activity in the Bundelkhand massif region. These reefs are considered as a unique feature of this massif, occurring mostly along brittle-ductile shear zones (Sharma 1998). As per Mondal et al. (2002), they appear related to granitic magmatism that took place around 2.3–1.9 Ga. Further, the granitoids and gneisses are also of calc-alkaline nature, characterised by highly fractionated REE patterns depleted in HREE, thereby conforming with Archean TTG suites and resembling remnants of early crustal components (Sharma 2009). This massif is bounded on its west, east and south by Meso- to Neoproterozoic Vindhyan sediments, while on its north, it is covered by young Tertiary sediments. Both the cratons, Aravalli as well as Bundelkhand, have similar lithology and have been affected by similar deformational events (Naqui and rogers 1987). Both of them seems to have cratonised around 2.5 Ga, compared to Singhbhum craton,

4.2 Bundelkhand Craton

131

which stabilised much earlier at around 3.1 Ga (Mishra et al. 1998). Further, in both the cratons, oldest granitic crustal component has a similar age of about 3.3 Ga (Sharma 2009). Also the gneisses from both the terrains have common compositional characteristics (Gopalan et al. 1990; Mondal 2003). Therefore it is felt that both of these cratonic provinces may have evolved as single large protocontinent that stabilised around 2.5 billion years ago (Mondal 2003; Yedekar et al. 1990), an initial hypothesis floated by Naqui et al. (1974). A Detailed reviews on regional geology can be found in GSI (2004), Basu (2007) and Singh et al. (2007). A simplified geologic map of the region is reproduced in Fig. 4.10a.

4.2.1 Geophysical Characteristics 4.2.1.1

Crustal Seismic Structure

No DSS study is undertaken over this craton, although the surrounding Vindhyan sediments, are well studied as discussed in Chap. 5. The only seismic profile which touches the southern part of the Bundelkhand massif is the Hirapur-Mandla seismic profile. The interpreted crustal section is shown as profile AB in Fig. 5.2 (Chap. 5). This section reveals that close to the massif terrain, between Hirapur and Narsinghar, the granitic-gneissic upper crust is about 7–8 km thick, underlain by about 12 km thick middle crust and more than 20 km thick lower crust. The Moho is detected at 41–43 km depth (Murty et al. 2004, 2008). Narsighar in the south (located on profile AB in Fig. 5.2), can be considered as the southern boundary of the massif. This would mean absence of Bundelkhand cratonic terrain below the Son valley Vindhyan sediments. However, some researchers suggests that this craton extends below Son valley Vindhyans. No underplating is reported from this region. The results obtained by seismic studies are largely compatible with those derived from receiver function studies. Receiver Function Studies Both broadband seismic stations, JHN and BRT (Table 4.1), reveal almost similar S-wave velocity structure (Julia et al. 2009). In comparison to 7–8 km thick granitic gneissic upper crust in the southern part of the craton as revealed by DSS studies, it is little thicker (~12.5 km) below the northern parts (Fig. 4.4c, d), where these two stations are located. Here, the middle and lower part of the mafic crust may be about 5 km and 17.5 km thick respectively. Underplated magma layer above the Moho may either be very thin (~2–3 km) or could be altogether absent. The Moho is thinner by about 5 km in the northern parts of the craton, where a consistent thickness of about 37.5 km is found below both the seismic stations. Julia et al. (2009), based on results obtained at the SGR station, close to the outcropping Bundelkhand block, suggested that this block is not extending under the Vindhyan sediments, as also reported by DSS studies (Murty et al. 2004, 2008). Magnetotelluric (MT) studies carried out by Gokarn et al. (2013) over the Vindhyan

132

4 Aravalli and Bundelkhand Cratons

Fig. 4.10 a Geological map of Bundelkhand region together with Magnetotelluric sites, and b Geoelectric section of Bundelkhand craton as obtained from the magnetotelluric studies. A-F’s are major resistivity features revealed from MT study, as discussed in the text. Modified after Figs. 1 and 3, Gokarn et al. (2013)

4.2 Bundelkhand Craton

133

Group of rocks further supported this conjecture. They did not find highly resistive Bundelkhand cratonic rocks underneath the Vindhyan sediments. The receiver function studies at the seismic station BRT, on the other hand, suggest that the Bundelkhand cratonic block may be extending by several hundreds of kilometres below Indo-Gangetic plain (Julia et al. 2009).

4.2.1.2

Gravity Studies

Several attempts were made in the past (e.g. Qureshy and Warsi 1975; Agrawal et al. 1993; Mishra and Rajasekhar 2008; Gokarn et al. 2013; Mishra and Ravi Kumar 2014) to study the gravity field over this region using the then available gravity data. A vastly improved new Gravity Map Series of India on 1:2,000,000 scale and with 5 mGal contour interval (GSI-NGRI 2006) is now also available. This map series, which is published by CSIR-NGRI and GSI, included large number of gravity data acquired by various Indian agencies. The updated Bouguer gravity anomaly map pertaining to Bundelkhand and its surrounding areas (Vasanthi and Singh 2019) is shown in Fig. 4.11. Gravity field over the Bundelkhand massif region, varies from −30 to −70 mGal and divides the massif into two parts. The northern half of the Bundelkhand craton above the latitude 25° N, is associated with a prominent regional gravity high (−30 to −45 mGal) that passes through Jhansi. Immediately north of this latitude, lies the Bundelkhand Tectonic Zone (BTZ), containing mantle rocks peridotites, pyroxenites and relics of the oceanic crust like, the pillow basalts and ophiolitic mélanges (Malviya et al. 2006; Pati et al. 2007; Gokarn et al. 2013), apart from ultramafic rocks further north (Sharma and Rahman 2000). This region is also known to contain Gwalior and Bijawar metasedimentary basins (characterised by high density), possibly overlying the middle crust (Srivastava et al. 2009; Pandey et al. 2014), apart from biotite granites, grey granites, metavolcanics and a number of gneissic complexes, all of which are mafic in nature. The Bijawar group of rocks, consisting metasediments with major basic/ultrabasic intrusives, may have been formed in a rift environment during Paleo-Mesoproterozoic (Mishra 2011; Mishra and Ravi Kumar 2014). The DSS studies suggest, that this region may have been associated with a paleo volcanic front, caused by the subduction of oceanic part of the Bastar cratonic lithosphere, below the northern part of Bundelkhand terrain (Mandal et al. 2013). In these areas, vertical block movements, uplift and exhumations are quite common. There are also reports of felsic volcanism in the central part of the massif. The regions lying south of 25° N latitude are, however, characterised by a broad regional gravity low (−50 to −70 mGal) associated with the southern part of the Bundelkhand, consisting mainly potassic granites. Such granites may have thicker roots with deeper Moho at about 43 km depth, as mentioned earlier. This gravity low appear sandwiched between two gravity highs (Fig. 4.11).

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4 Aravalli and Bundelkhand Cratons

Fig. 4.11 Bouguer gravity contour image of the Bundelkhand craton, Vindhyan basin and its environs, using new gravity map series (GSI-NGRI 2006). J: Jabera basin, NSL: Narmada-Son Lineament. After Fig. 2, Vasanthi and Singh (2019)

4.2.1.3

Magnetotelluric Studies

A detailed MT study over this craton has been carried out by Gokarn et al. (2013). They acquired broadband MT data at 25 sites, in the frequency range of 320–0.001 Hz, along a 250 km long Jabera–Jhansi profile (Fig. 4.10), which is virtually a northward extension of the earlier profile (Gokarn et al. 2001). This survey covered the entire exposed terrain of the Bundelkhand granitic massif. The interpreted geoelectric section is shown in Fig. 4.10b; it reveals presence of high resistive (5000 -m) block in the northern part of the craton (Marked as F in Fig. 4.10b) that extends from surface to 60 km depth. They delineated a three-layer resistivity structure (marked by C, D and E) overlying a fourth conductive layer at a depth of about 60 km in

4.2 Bundelkhand Craton

135

the southern part of the craton. This area lies between exposed Bijawar sediments, and the Bundelkhand tectonic zone (BTZ) containing ultramafic rocks, as mentioned earlier in the gravity section. The Vindhyan basin and Bijawar sediments exposed further south are characterised by low resistivity. Broadly, the conductivity distribution suggests a northward dipping trend. It is inferred that the Bundelkhand craton, characterized by high resistivity crustal rocks, does not extend to the south beneath the low resistive Vindhyan basin crust (marked as A, B in Fig. 4.10b), containing about 4–5 km thick layer of the upper and lower Vindhyan rocks between Jabera and Hirapur, overlying the Bijawar sediments (Gokarn et al. 2001). Presence of partial melt at subcrustal depths, as deciphered below Jabera (Pandey et al. 2014), may have released hydrothermal fluids and volatiles, that caused high crustal conductivity, as also indicated by Gokarn et al. (2013).

4.2.1.4

Heat Flow and Lithosphere Structure

Recently, Nagaraju et al. (2017) have reported heat flow values for ten new sites out of which 7 sites are located over the Bundelkhand massif (Fig. 4.8). Most of these sites are close to the fringes of the massif, except the Dinara (Jhansi), which lies in its northwestern part. Heat flow over this craton varies in a narrow range from 34 to 42 mW/m2 with a mean of about 37 mW/m2 . This excludes the values obtained over Panna region, which lies over the Vindhyan outcrops. This study also carried out radioactive heat generation measurements over 243 samples of drill cores, which resulted into a high average heat production of 4.0 µW/m3 for NeoarcheanPaleoproterozoic granites and 2.0 µW/m3 for Mesoarchean TTG gniesses. Using crustal structure provided by Murty et al. (2004, 2008) and heat flow-heat generation by Nagaraju et al. (2017) and Ray et al. (2003) and taking into account other relevant thermal parameters from Pandey et al. (2014) (as detailed in Table 4.2), temperature-depth distribution has been calculated for this craton, which is shown in Fig. 4.9. It reveals Moho temperature about 435 °C, mantle heat flow ~21 mW/m2 and lithosphere thickness about 170 km, which is one of the highest recorded over the Indian terrain.

4.3 A Composite Geodynamic Model A generalised crustal seismic structure of the Aravalli craton and adjacent Chambal valley Vindhyan region, is shown in Fig. 4.3, which provides some interesting inferences about the evolution of the north Indian cratons. It is now well known that during the Paleoproterozoic (2.5–1.6 Ga), a supercontinent Columbia existed. Subsequent breakup, and further accretion and collision of the segments of this supercontinent, resulted into the formation of another supercontinent Rodinia. This was

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a global event and it appears that during the Columbia period (~1900 Ma; Mandal et al. 2018), Bundelkhand cratonic micro-plate started subducting below the Mesoarchean Mewar terrain of the Aravalli craton, along the Great Boundary Fault (GBF) (Fig. 4.3). In the present scenario, Hindoli Group and Mangalwar Complex would represent erstwhile forearc, while the Aravalli Fold Belt-Sandmata complex and the Delhi Fold Belt, as the volcanic front. Presence of parallelly dipping Jahazpur thrust (JT), Great boundary thrust (GBT) and Chambal thrusts(CT) below the Mangalwar and Hindoli complexes and partly below Vindhyans (Mandal et al. 2014, 2018; Vijaya Rao et al. 2000), and their generation due to compressive forces related to subduction and subsequent collision episodes, bears the testimony to such an event. This was also the period when the development of the Aravalli Fold Belt might have taken place due to multiple extensional and orogenic events that took place between 2.0 and 1.5 Ga on a weak lithosphere. This was further followed by the development of Delhi Fold Belt during 1.5 and 0.75 Ga (Sinha-Roy 2008) due to subduction related rifting and volcanic arc formation. An unprecedented 18 km thick magma underplating (Vp ~ 7.3 km/s) is observed below this region in the form of domal structure, as demarcated by prominent seismic reflectors. The lower crust, all along east of the Marwar basin, is heavily underplated by 10–18 km thick mafic magma, consequent to sustained magmatism and crust-mantle thermal interaction, associated with the Aravalli terrain. It further appears that the western part of the Aravalli craton, is associated with another subduction phase, as revealed by steeply dipping discontinuous bands of isolated reflections (Krishna and Vijaya Rao 2011; Vijaya Rao and Krishna 2013). These discontinuous bands are considered remanents of earstwhile subducting Marwar craton, during Mesoproterozoic. Signature of a collisional suture formed by two subducting slabs, Marwar craton from the western side and the Bundelkhand-Mewar craton from the east, can clearly be seen below Rian around SP 891 (Fig. 4.3), which ultimately led to deformation of the Delhi Fold Belt.

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Chapter 5

Vindhyan Basin: Anomalous Crust-Mantle Structure

5.1 Introduction In the Earth’s evolutionary history, Paleo- to Mesoproterozoic period is characterised by the development of several intra-cratonic basins in stable continental shields and platform areas world-wide (Windley 1977; Condie 1989). Indian subcontinent was no exception. It contains four such major sedimentary basins; Vindhyan basin, Cuddapah basin, Pranhita–Godavari basin, and Chhattisgarh basin (Fig. 1.15, Chap. 1). All these basins came into existence due to sustained rifting and subsidence of the central part of the Indian peninsular shield. Out of these, the sickle shaped Vindhyan basin (Fig. 5.1), containing thick sequence of Vindhyan Supergroup of rocks, is considered one of the largest Proterozoic basins in the world that seemingly contains clues to the many paleobiological problems. This basin is situated north of the Narmada-Son Lineament (NSL) in the northcentral part of India (Chakraborty 2006; Ramakrishnan and Vaidyanadhan 2008). It spreads in east-west direction from Sasaram in Bihar to Chittorgarh in Rajasthan, thereby covering vast areas of Rajasthan, Bundelkhand and Son valley sectors. It more or less encircle the Bundelkhand granitic-gneissic massif (Fig. 5.1). It is believed that the depositional and evolutionary history of the Vindhyan basin, is intimately connected to the evolution of NSL, as no trace of these rocks are found south of this mega lineament. The Vindhyan sedimentary succession covers an area of about 1,20,000 km2 in central India, while about 80,000 km2 is understood to be concealed below Deccan volcanics. Geophysical and deep drilling data (Narain and Kaila 1982; Srivastava et al. 1983) further suggest that around 10,000 km2 area is further concealed below the Indo-Gangetic plain (Mathur 1987). Several productive diamond mines are known to be located over this region. The southern part of the basin contains 5–6 km thick sediments in which presence of hydrocarbons is expected (Das et al. 1999). In view of this, Oil and Natural Gas Corporation (ONGC) drilled three exploratory wells, Jabera-1, Damoh-1 and Kharkhari-1. A detailed gravity survey, with 1500 new gravity observations, is made © Springer Nature Switzerland AG 2020 O. P. Pandey, Geodynamic Evolution of the Indian Shield: Geophysical Aspects, Society of Earth Scientists Series, https://doi.org/10.1007/978-3-030-40597-7_5

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Fig. 5.1 Simplified geological map of the Vindhyan basin (Azmi et al. 2007; Bengtson et al. 2009). Solid rectangles refer to location of broadband seismic stations. GBF is great boundary fault. AB, CD, EF are the respective locations of three DSS profiles, Hirapur–Mandla, Khajuriakalan– Rahatgaon and Ujjain–Mahan. Location of deep seismic reflection profile shot across Chambal valley Vindhyans between Chandli and Kunjer, is also shown. Irregular rectangle in southern part of Son valley Vindhyans (shown by broken lines), indicate the area covered by detailed gravity studies by Srivastava (2006). Modified after Fig. 1, Pandey et al. (2014b)

in and around Jabera-Damoh-Katangi areas of this region, utilising a newly developed fractal based gridding approach (Srivastava 2006; Srivastava et al. 2007, 2009). Apart from the gravity, other geophysical methods are also used to decipher the basement configuration and deeper structures. Three DSS and one seismic reflection profile, cover the northern, western, and southern part of the Vindhyan basin (Fig. 5.1). The updated seismic sections indicate an anomalous crustal structure (Tewari et al. 2002; Murty et al. 2004), which is also supported by subsequent MT studies (Gokarn et al. 2001). Presence of a large thickness of Vindhyan strata (about 7.5 km) is recorded below the Chambal Valley Vindhyan region (Mandal et al. 2018). Receiver function studies are carried out in five areas, out of which two are directly located over the thick Vindhyan strata (Julia et al. 2009; Vijay Kumar et al. 2012). Heat flow studies are carried out at only two heat flow sites (Fig. 4.8, Chap. 4); one in Shivpuri in upper Vindhyans with heat flow 45–61 mW/m2 (Nagaraju et al. 2012) and the other in Panna, close to Bundelkhand massif, with heat flow 36– 37 mW/m2 (Nagaraju et al. 2017). Due to thick sequences of Deccan volcanics and Indo-Gangatic alluvium, evolution of this basin is not much understood; thus more geophysical studies are needed.

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5.2 Geotectonic and Geologic Features The Vindhyan basin apparently evolved as a foreland basin on a rifted Indian crust, caused by continued crustal extension (Jokhan Ram et al. 1996). Its northwest margin is marked by a NE-SW trending Great Boundary Fault (GBF), which virtually separates this basin from the Aravalli-Delhi Mobile Belt (ADMB), as has been discussed in Chap. 4. Southern boundary of the basin is demarcated by the Proterozoic NSL and eastern boundary by the folded Satpura Massif. The NSL region (also referred as the SONATA rift), associated with complex thermo-tectonic history, divides the Indian peninsular shield virtually into two major geotectonic segments, that exhibit distinct geological, geothermal and geophysical signatures on its either sides (Pandey and Agrawal 1999). Sediment depositional history of this basin appear intimately connected to the evolution of this rift structure as mentioned before. Neoarchean Bundelkhand granitic-gneissic massif, situated in north-central part of the Bundelkhand craton (Fig. 5.1), divides this basin into two major segments; Chambal valley Vindhyans to the west and Son valley Vindhyans to the east. Deep seismic sounding studies reveal possible existence of a basement ridge between these two distinct Vindhyan terrains (Rajendra Prasad and Vijaya Rao 2006).

5.2.1 Geochronology and Stratigraphy Several geological investigations are carried out to study the geochronologic sequence of Vindhyan rocks (Kumar and Sharma 2010). Based on recent radiometric dates of different horizons of the Lower Vindhyans, the onset of the Vindhyan sedimentation is suggested to be around 1600 ± 50 Ma (Rasmussen et al. 2002; Ray et al. 2002, 2003a; Sarangi et al. 2004). Ray (2006) reviewed the available ages of the Vindhyan Supergroup and supported above findings. It is suggested that sedimentation in the Son Valley basin, may have started some time prior to 1721 Ma and continued possibly till the end of Proterozoic, thus covering a span of nearly 1000 Ma. This region is also known for kimberlitic magmatism at number of places. One of such kimberlite pipes, which intrude the Kaimur sandstone in Panna area, is dated 1073 ± 13 Ma by 40 Ar/39 Ar method (Gregory et al. 2006), which correlates with the kimberlite-clan rocks found in the vicinity of Cuddapah basin (Anil Kumar et al. 1993). Stratigraphically, the Vindhyan rocks can be classified into four individual stratigraphic groups, Semri, Kaimur, Rewa and Bhander (Chakraborty 2006). The Semri group containing lower Vindhyan rocks, is the oldest and the thickest, and forms the bottom of Vindhyan sediments that rests directly over the variety of crystalline rocks like Archean Bundelkhand granite-gneisses, Bijawar Group of metamorphics and Mahakoshal group of rocks, belonging to Paleoproterozoic age. Lithostratigraphically, the Vindhyan Supergroup comprises mainly un-metamorphosed and mildly

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deformed sandstone, shale and limestones, with a few volcanoclastic and conglomerate beds (Chakraborty and Bhattacharyya 1996). These were deposited in a shallow marine environment with large sedimentary thicknesses (5–6 km) occurring in its southern part near Jabera (Srivastava et al. 2007, 2009; Kaila et al. 1989; Valdiya et al. 1980; Bhattacharya and Morad 1993), as well as in Kota-Kunjer section of the Chambal valley Vindhyans (Mandal et al. 2018). Some investigations have however revealed that the Vindhyan sediments may represent deeper part of the shelf as well, spread by fluvial processes (Chakraborty and Bhattacharya 1996).

5.2.2 Jabera-Damoh Region This region, shown as zig-zag rectangle in Fig. 5.1, borders Narmada-Son lineament in the south and contains large thickness of Vindhyan sediments. The Hirapur-Mandla DSS profile (profile 1) passes through this region. This part of the Vindhyan basin, is largely disturbed, associated with folds, wrenches, boundary faults and contains a prominent erosional feature (Jabera dome) located at about 40 km NW of Jabalpur that extends in ENE-WSW direction (Jokhan Ram et al. 1996; Srivastava 2006). It occupies an area of about 320 km2 , wherein lower horizons of the upper Vindhyan rocks are exposed, which are further surrounded by upper Vindyans, dipping in all directions. It contains sandstone, shale and limestone with some basaltic intrusive. On the other hand, the Damoh structure, situated north of Jabera, is associated with lower Vindhyans only. It is bounded by two major faults trending NW–SE. Detailed geological and geotectonic information of this region can be found in Auden (1933), Valdiya et al. (1980), Bhattacharya (1996), Chakraborty (2006), Ray and Chakraborty (2006), Srivastava (2006), Srivastava et al. (2007, 2009) and Kumar and Sharma (2010).

5.3 Geophysical Characteristics This region with a number of tectono-thermal upheavals in the geological past, played an important role in rifting, deforming and reshaping the underlying crust as well as mantle lithosphere in entire central India. Geophysical information, other than the gravity and deep crustal seismic studies, are largely lacking.

5.3.1 Crustal Seismic Structure from DSS This basin has been seismically imaged using three DSS profiles which run through the southern part of the basin comprising Son Valley Vindhyan formations (Fig. 5.1). The fourth one, the crustal seismic reflection profile, runs through the Chambal

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valley Vindhyans in its northwest segment. Son-valley Vindhyan profiles include, (i) Hirapur-Mandla profile passing through the eastern part of the basin, while (ii) Khajuriakalan-Pulgaon and (iii) Ujjain-Mahan profiles, pass through the western part of the Son Valley Vindhyans, which are concealed below Deccan volcanics. Along all the three profiles, shallow and deep seismic refraction and wide-angle reflection data are acquired earlier in analog form. These data were later digitized to produce composite record sections, which provided first hand crustal structure and the Moho configuration (Kaila and Krishna 1992). The updated seismic crustal sections (Murty et al. 2004; Tewari et al. 2002) along all the three profiles (AB, CD, EF) are shown in Fig. 5.2, which indicate highly anomalous nature of the crust below these areas.

5.3.1.1

Hirapur-Mandla (Profile 1)

Detailed crustal seismic section along this 235 km long profile (Murty et al. 2004, 2008) which runs from Hirapur near the southern border of Bundelkhand granitic massif to Mandla, located in Deccan Traps terrain, is shown as profile AB in Fig. 5.2. This profile covers almost entire stretch of the exposed Vindhyan rocks north of the NSL. Two major faults detected at Narsinghar and Katangi, conform to graben structure, filled with thick pile of upper and lower Vindhyan sediments. The region further south, bounded between the Katangi and Jabalpur faults, corresponds to a horst feature coinciding with the NSL rift zone. A Five-layer crust delineated between Damoh and Katangi regions indicate that the Vindhyan sediments are as much as about 5 km thick and sit directly over an extremely thin granitic–gneissic upper crust (5.9 km/s), which may even be altogether absent in many areas. These inferences conform very well with the gravity and MT studies (Gokarn et al. 2001), as discussed in subsequent sections. Further, the sediments are characterize by two distinct P-wave velocities; Upper Vindhyans with a Vp 4.5–4.7 km/s, compared to a much higher Vp 5.3–5.4 km/s for the lower Vindhyans (Murty et al. 2004). It means that the lower Vindhyans are denser and much more compact, compared to the upper Vindhyan sediments. Presence of high velocity crust (Vp 6.5 km/s or even more) is observed immediately below the thin wafer of granitic-gneissic layer (5.9 km/s), at the depths between 6 and 20 km. This layer is characterized by Vp 6.5 km/s at the top section, followed by a Vp 6.35 km/s at the bottom. Magnitude of the velocity drop in this layer situated at the mid to lower crustal transition, may represent metasomatic alteration of the in situ rocks due to biotitisation, saussuritization, iron enrichment etc. (Pandey et al. 2016). This metasomatically altered layer is underlain by a 20–22 km thick lower crust with Vp 6.8 km/s above the Moho; the Moho is delineated at depth 41–44 km. No magmatic layer is reported from this region, which has been conspicuously found along other two profiles.

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Fig. 5.2 2-D crustal velocity model along Hirapur–Mandla (Modified after Fig. 8, Murty et al. 2004), and Khajuriakalan–Rahatgaon and Ujjain–Mahan seismic profiles (Modified after Figs. 7 and 8, Tewari et al. 2002). SP refers to shot points. Layer velocities are in Km/s

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149

Khajuriakalan-Rahatgaon (Profile 2)

Only a couple of hundred meters thick Proterozoic Vindhyan and Bijawar sediments are present along this profile (Profile CD in Fig. 5. 2), which rest directly over a 5– 7 km thick upper (granitic-gneissic) crust (Vp 6.0–6.1 km/s). This layer is underlain by a 13–14 km thick high velocity (Vp 6.5–6.6 km/s) crust, bottom half of which appear metasomatised and thus has a much lower velocity (Vp 6.3 km/s), similar to that found below the Hirapur-Mandla profile. Below this region, lower crust is about 20 km thick, having a typical lower crustal Vp of 6.7 km/s in top 10 km, followed by a more than 10 km thick high velocity (Vp 7.2 km/s) magmatic layer just below it. No such high velocity magmatic layer is reported from the earlier discussed Hirapur-Mandla profile. Below this region, on an average, the crust is about 41 km thick.

5.3.1.3

Ujjain-Mahan (Profile 3)

Along this traverse (Profile EF in Fig. 5.2), the seismic crustal structure is somewhat similar to that found below the Khajuriakalan-Rahatgaon profile. In the Vindhyan segment of this profile, cumulative thickness of the Deccan volcanics, Lameta beds, Vindhyans and Bijawar group of rocks are only about 600 m thick, above the crystalline basement, (Kaila et al. 1985). Granitic-gneissic layer (Vp 6.1 km/s) is again about 7–8 km thick, which rests over 16 km thick mid-level crust, first half of which is associated with a high Vp of 6.6 km/s, while the metasomatised bottom layer, has a Vp 6.3 km/s. The lower crust is about 17 km thick, out of which magmatic crust (Vp 7.2 km/s) at the bottom is 10 km thick. Moho is detected at about 41 km depth.

5.3.1.4

Metasomatised Low Velocity Layer at Mid to Lower Crustal Transition

All the three DSS profiles (Fig. 5.2) revealed presence of about 5 to 12 km thick low velocity zone, in which velocities vary from 6.3 to 6.4 km/s, compared to 6.5–6.7 km/s at the top and 6.7–6.8 km/s at the bottom. Such velocity drop (~6%) may indicate mantle-fluid driven metasomatic transformation (or alteration) of the crust, caused by thermal and magmatic influx from below (Pandey et al. 2016). Such processes usually takes place during retrograde metamorphic conditions, wherein prevailing P&T conditions shifts to lower grade due to exhumation of underlying rocks. Such processes are quite common below the Deccan volcanic Province (Desai et al. 2004; Sen et al. 2009; Tripathi et al. 2012a, b; Pandey et al. 2014a, 2016). Laboratory measurements on KLR-1 borehole samples from Killari earthquake region, indicate that such velocity drop in mid crustal rocks due to metasomatism, could be as much as 15% (Pandey et al. 2016). This would further indicate that southern part of the Vindhyan basin was subjected to much higher in situ temperature, than usually recorded

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in Proterozoic terrains elsewhere. Possibly no other explanation can justify presence of such low velocity zones at mid to lower crustal levels.

5.3.1.5

Bundi-Kota–Kunjer Seismic Profile

A 60-fold deep seismic reflection profile was shot along a 165 km transect between Chandli and Kunjer (Fig. 5.1), which after cutting across the Great Boundary Fault, traversed through almost entire sequence of the Meso- to Neoproterozoic Chambal valley Vindhyans. This section consists of Bhander, Rewa and Kaimur Group of rocks, belonging to Upper Vindhyan sequence, followed underneath by Semri Group, representing the lower Vindhyans. The seismic data along this profile has earlier been processed by Rajendra Prasad and Vijaya Rao (2006) and Reddy and Vijaya Rao (2013), using CMP technique. Recently, Mandal et al. (2018) also analysed the crustal seismic reflection data along this profile, but used common reflection surface stack procedure, which has better resolution. They found conspicuous presence of a major thrust zone (named Chambal Thrust, CT) below–the Bundi-Kota section of the Chambal valley Vindhyans. It lies at a depth between 9 and 18 km below Kota and between 20 and 30 km depth below Bundi (Fig. 4.3, Chap. 4), which may correspond to part of the relic associated with Paleoproterozoic subduction of Bundelkhand craton below Aravallis. Their depth-migrated shallow seismic section (Fig. 5.3) further revealed gently southeast dipping three highly reflective bands at 3, 6 and 7.5 km depth below SP 4101 (situated southeast of Kota) and 4, 7 and 9 km depth below Kunjer (SP 4747). This would suggest that beneath Chambal Valley sequence, sediments are extremely thick at about 7.5 km, out of which Upper Vindhyans are about 4 km thick and lower

Fig. 5.3 Crustal seismic section obtained along deep seismic reflection profile between Kota and Kunjer segment of Chambal valley Vindhyans, together with estimated upper and lower Vindhyan sediment horizons, based on Mandal et al. (2018). SP refers to shot points

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Vindhyans, 3.5 km. These sediments are underlain by 1.5 km thick volcanic layer above the granitic-gneissic basement, detected at the depth of 9 km. This would support the observation made by Raza et al. (2009), who has earlier indicated occurrence of basalts close to western and southern margin of the Vindhyan basin. The Moho depth ranges from 40 km in western side, to little deeper at around 44 km near Kunjer.

5.3.2 Broadband Seismic Studies 5.3.2.1

S-Wave Velocity Structure

The S-wave velocity structure of the Vindhyan region was investigated at 5 sites (SGR, RWA, BTL, MHR and BPL; Fig. 5.1) by Kumar et al. (2007), Julia et al. (2009) and Vijay Kumar et al. (2012). They used receiver function technique to obtain S-velocity variations, Moho depth and Vp/Vs ratios. All these seismic stations are located in southern part of the basin, consisting Son valley Vindhyans. The S-wave velocity structures below each station, are included in Fig. 5.4, while their averages are shown in Fig. 5.5. These Figures indicate Moho depth variation in a narrow range from 38 to 44 km, with a mean of around 42 km. However, Julia et al. (2009) found deep Moho (52.5 km) below the BPL station, due to presence of a thick high velocity-high density underplated magma layer. In this layer, Vs gradually increases from 4.03 to 4.43 km/s at the depths between 35 and 52.5 km, indicating its differentiated nature. Quite likely, crustal magma layer below BPL is only about 3 km thick above the Moho, while the remaining may be a mantle magma layer. That would indicate strong thermal interaction between the lower parts of the crust and the underlying mantle. Kumar et al. (2007) too reported only a 38 km thick crust below this station, not 52.5 km as mentioned above. In Fig. 5.5, an inferred crustal cross section beneath southern part of the Vindhyan basin is also shown, which is based on averaged S-wave velocity distribution. It suggests that on an average, Vindhyan sediments are about 5 km thick in the southern part (similar to that obtained by DSS studies), which is followed by 7.5 km thick upper (granitic-gneissic) crust, where velocities gradually increase from 3.35 to 3.57 km/s with depth. This layer is followed by 2.5 km thick middle crust characterised by Vs 3.65 to 3.71 km/s. Gradual increase in velocities in both of these layers indicate its fractionated nature, suggesting assimilation and fabric alteration during exhumation of the lower crust to shallow depth (15 km). The lower crust beneath this region is, however, much thicker (~27 km), which can further be subdivided into two distinct parts; upper one between 15 and 35 km depth corresponding to the lower crust, while the lowermost part between 35 and 42 km depth, represents stacked magma in which shear velocities gradually increase from 4.01 to 4.31 km/s. Such situations often arise in rift zones, which are subjected to strong thermal interaction between the lower crust and the hot, enriched and buoyant asthenosphere, that induces subcrustal erosion consequent to rise in mantle solidus. They often cause magma underplating,

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Fig. 5.4 S-wave velocity variations with depth below five broadband seismic stations (Julia et al. 2009; Vijay Kumar et al. 2012), located in the southern part of the Vindhyan basin. After Fig. 8, Pandey et al. (2014b)

a process responsible for growth and thickening of the lower crust and consequently, deepening of the Moho. Crustal Exhumation Figure 5.5 may draw another inference that the crust beneath this region may have exhumed by as much as 10–13 km, with its top located at about 12 km depth, thereby indicating erosion of a thick chunk of the upper crust from this region due to exhumation. A velocity increase (~0.2 km/s) is seen between the depths 12 and 25 km, in comparison to global shields and platforms (Christensen and Mooney 1995). Below this depth, lower crustal velocities are lower, but conforms with the global velocities. The average shear wave velocity at depths between 15 and 35 km is

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Fig. 5.5 An inferred crustal seismic section below southern part of the Vindhyan basin based on averaged S-wave velocity distribution for seismic stations shown in Fig. 5.4. For comparison, Swave velocity distribution beneath global shields and platforms (shown as solid dots and obtained by converting Vp values from Christensen and Mooney 1995 into Vs, using Vp/Vs = 1.75), is also included. Significantly higher velocities can be seen between 13 and 25 km depth, suggesting massive exhumation of the mafic crust (Pandey et al. 2014b)

relatively low ~3.87 km/s (normal being 3.80–4.00 km/s); this indicates possibility of metasomatic alteration, as seen in DSS derived seismic sections too (Fig. 5.2). Restructuring of the ancient continental crust under the southern part of the Vindhyan basin, is thus well supported by geophysical data. Geodynamic processes associated with the evolution of the NSL rift highly influenced its crustal structure; the restructured Moho is about 42 km deep. Lithosphere-Asthenosphere Boundary Based on the receiver function studies, a lithosphere thickness of about 95 km, has been reported below the BPL broadband seismic station by Kumar et al. (2007). The BPL station is located on the south-western flank of the Vindhyan basin (Fig. 5.1),

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concealed below the Deccan volcanics. This estimate is quite close to that obtained for other parts the Indian peninsular shield (Polet and Anderson 1995; Pandey and Agrawal 1999; Pandey et al. 2017).

5.3.3 Gravity Investigations 5.3.3.1

Regional Gravity Field

Regional gravity field distribution over the Bundelkhand craton and its adjacent regions has earlier been shown in Fig. 4.11 (Chap. 4), which also encompasses the regions occupied by the Vindhyan sediments, exposed west and east of the Bundelkhand massif. Over the Chambal valley Vindhyans, occurring west of the massif, Bouguer gravity anomalies vary from −55 to −65 mGal in northern part and −35 to −45 mGal towards south, indicating small thickness of sediments in the vicinity of the Bundelkhand massif, or presence of volcanics/Bijawar group of rocks underneath. Mishra et al. (2000) carried out gravity and magnetic investigations across the Nagaur-Jhalawar (Kunjer) geotransect, as discussed earlier in Chap. 4. It passed through a section of Chambal Valley Vindhyans, as shown in Figs. 4.5 and 4.6 (Chap. 4). As can be seen, the observed gravity anomalies over this sedimentary sequence, range from −40 to −65 mGal. Such order of gravity anomalies would not match with the occurrence of about 7.5 km thick Vindhyan sediments, which on an average, have a low bulk density of 2.56 g/cm3 (Mishra et al. 2000). Therefore, infusion of magmatic material at subsurface depths is a real possibility. Mandal et al. (2018) did find occurrence of 1.5 km thick volcanic layer just above the basement. Relatively higher gravity can be ascribed to exhumed and upthrusted lower crustal rocks, as indicated by conducting nature of the basement. Over this region, total magnetic intensity is also much higher (exceeding 46,000 nT on eastern side) than found over the adjacent Aravalli cratonic segments comprising Mangalwar and Hingoli group of rocks. This would further support presence of high density material at subsurface depths. As for as Son Valley Vindhyans are concerned, over a major part, Bouguer gravity anomalies are of the order of −35 to −55 mGal, which continue over Deccan Traps (Fig. 4.11, Chap. 4). This observation suggests extension of Vindhyan sediments underneath; the sequence, however, many not be thick. In contrast, in the southern part of the Vindhyan basin, gravity anomalies reach around −90 mGal in the Jabera basin, which lies close to the NSL and contain 4–6 km thick Vindhyan sequence.

5.3 Geophysical Characteristics

5.3.3.2

155

Detailed Gravity Survey Over Southern Part of Vindhyan Basin

The southern part of the Vindhyan basin, which adjoins the NSL rift zone in the south, is associated with multiple geologic and geodynamic upheavals, like massive uplift, crustal rifting and volcanic extrusions. Srivastava (2006) and Srivastava et al. (2009) undertook a detailed gravity survey. They designed a close spaced gravity network, based on optimum fractal dimension, to delineate detailed subsurface structures below 112 × 100 km area (Fig. 5. 1). This survey covered parts of Katni, Damoh, and Panna districts of Madhya Pradesh in Central India. In this experiment, as many as 1500 new gravity stations were covered and 40 new gravity bases were established (Srivastava 2006; Srivastava et al. 2009) at about 10–15 km spacing. Measurements were carried out along all the accessible roads and tracks at a varying interval of 0.5–2.0 km, using Lacoste-Romberg G-type gravimeters (LRG), having a precision of 0.01 mGal. Bouguer gravity anomaly map, prepared using fractal based gridding method, is shown in Fig. 5.6. Although, the study area is small (about 1° × 1°), it shows significant variations in Bouguer gravity anomaly from −25 to

Fig. 5.6 Bouguer gravity anomaly map of Jabera–Damoh and adjacent areas of the Vindhyan basin in central India, obtained using fractal method (Srivastava 2006). Locations of the modeled gravity Profile AA1 (as shown in Fig. 5.7), as well as deep boreholes W1, W2 drilled by ONGC, are also shown. H1, H2 and L1, L2 represents delineated gravity highs and lows in the study area. A prominent boundary feature (red line) divides H2 and H1. Modified after Fig. 6.1, Srivastava (2006)

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−74 mGal and reflects underlying subsurface features quite clearly. It shows two prominent gravity highs H1 and H2 and two gravity low regions L1 and L2, which were not reflected in earlier regional maps. Gravity anomalies are particularly low in the southern part of the Vindhyan basin near Jabera and further south. High gravity anomaly (H2) located towards NE of Damoh can be related to the presence of high density Mahakoshal group of rocks, which are exposed near Panna region (~24° 43 N, 80° 12 E). The incomplete pattern of this high Bouguer gravity anomaly (Fig. 5.6), indicates its further extension in NW direction. H1 represents the extent of this high gravity anomaly in the east. It is likely that 1.1 Ga kimberlitic event that affected Panna, Majhgaon (both proven diamondiferous regions) and its surroundings, caused uplift of the basement comprising of Mahakoshal group of rocks. Besides this, numerous kimberlitic and lamproitic intrusions can also cause high gravity anomaly. Some kimberlite pipes and associated ultramafic rocks, may be hidden in a widespread area below the Mahakoshal group of rocks, similar to Panna and Majhgaon. This study further revealed presence of an anomalous rifted structure, which is bounded by parallel faults on either side in the Jabera region. This rifted structure is marked as L2 in Fig. 5.6. Another gravity low (L1) which is of moderate nature near Damoh, possibly characterizes shallow sequences of lower Vindhyans. An inverse modelling of gravity data was attempted along profile AA1 passing through the gravity low L2 in south (Fig. 5.6), which cuts across major geotectonic features of the central part. Subsurface lithostratigraphy in the region is well constrained through well log data from two exploration wells (W1, W2 shown in Fig. 5.6) drilled by the ONGC near Jabera and Damoh. Initial gravity modelling constraints were taken from seismic information as well as well log data over the study area (Kaila et al. 1989; Das et al. 1999; Murty et al. 2004; Srivastava et al. 2007), besides measured density on the borehole core samples of different formations (Srivastava 2006). Profile AA1 The DSS study indicated maximum sediment thickness of 5 km in this area (Kaila et al. 1989; Murty et al. 2004). The gravity anomaly shows a typical anomaly pattern of a sedimentary basin faulted on either sides (Fig. 5.7). The preliminary depth estimates of the major interfaces were obtained from the 2D scaling spectral analysis of the gravity data. Modeling results indicate a deep faulted basin; the upper layer with 2.46 g/cm3 density corresponds to the upper Vindhyan, which is underlain by a thick layer (from 1.0 to 6.5 km depth) of lower Vindhyan sediments, with a gentle slope from NW to SE direction. 6.5 km thick Vindhyan sediments are underlain by high density (2.80 g/cm3 ) Bijawar/Mahakoshal group of rocks. In a major part of the Jabera basin, the Bijawar/Mahakoshal groups of rocks got deposited directly over high velocity (6.5 km/s) and high density (2.93 g/cm3 ) exhumed mid-crustal segment. This conjecture is supported by the seismic studies (Murty et al. 2004). It is even contemplated that the Vindhyans may have directly deposited over the lower crust in many segments (Gokarn et al. 2001). This model would further suggest that the granitic-gneissic layer is missing in the right margin of the faulted basin, while only a thin patch remains to its left side. Since this area lies in the vicinity of Jabera, it is termed as Jabera basin gravity low.

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Fig. 5.7 Shallow crustal structure as obtained from the inversion of gravity data along the profile AA1 (Fig. 5.6). ρ refers to density in g/cm3 . Modified after Fig. 6.3a, Srivastava (2006)

Extension of Vindhyan Sediments South of NSL Only a part of the rifted Jabera basin has been mapped gravimetrically (Fig. 5.6); in this map an incomplete closure of the Jabera basin gravity low is seen in the southern part, which extends for a much larger distance further south, as also revealed by Fig. 4.11 (Chap. 4). This gravity low is well reflected in the residual gravity field also. This would strongly suggest lateral extent of the Son valley Vindhyans extending further south below the Deccan volcanics, beyond the NSL rift zone, although there is no such reports yet and the deep borehole information is lacking to support this conjecture.

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5.3.4 Magnetotelluric (MT) Studies 5.3.4.1

Son Valley Vindhyans

Gokarn et al. (2001) carried out MT studies between Damoh and Anjaneya, which cuts across the NSL rift zone and covers the Damoh-Jabalpur section of the HirapurMandla DSS profile (Fig. 5. 2, profile 1). The results are compatible with the findings of the seismic studies. Presence of a 5 km thick conductive (6–30 m) Vindhyan sediments between Damoh and Katangi region, resting directly over the lower crust (Vp 6.5 km/s) characterized by 200 m resistivity, is reported. This study also indicated near absence of the upper crust from this region due to uplift and erosion. The MT study undertaken by Gokarn et al. (2013) over the Bundelkhand craton, is virtually a northward extension of this profile (Gokarn et al. 2001). Obtained results along this profile has already been discussed in Chap. 4. A conductive feature (conductivity 42.0

Ultramafic mantle

0.01

3.0

Data source Ray et al. (2003b), Pandey et al. (2014b, 2017), Nagaraju et al. (2012), Murty et al. (2004)

Fig. 5.9 Estimated temperature–depth profile for Jabera region of the Vindhyan basin, using crustal heat production model given in Table 5.1. Below this region, melting conditions are expected at an extremely shallow depth of about 50 km

the peridotite incipient melting point curve (Gass et al. 1978), a lithospheric thickness of about 50 km has been obtained for the Jabera rifted basin. Since radioactively rich granitic-gneissic layer is almost absent, a very high contribution of heat from the mantle (~56 mW/m2 ) is envisaged, apart from extremely high temperature of 1030 °C at the Moho level. Thermal regime of this part of the Vindhyan basin is comparable to that of the north Cambay basin of Gujarat (Pandey et al. 2017). Usually in Proterozoic terrains, such anomalous thermal regime is seldom observed.

5.4 Geodynamic Evolution of Vindhyan Basin Crustal extension and rifting over this terrain seems to have started at the beginning of Mesoproterozoic and lasted till the end of Neoproterozoic, spanning for over 900 Ma (Verma 1996; Venkatachala et al. 1996). During this period, almost 5–6 km thick upper and lower Vindhyan sediments deposited directly over the MahakoshalBijawar group of rocks in the rifted Jabera basin and as much as 7.5 km in the

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Chambal valley (Fig. 5.3). At some places, such deposition took place directly over a couple of km thick veneer of granitic-gneissic rocks. Sediment thickness is however much smaller (around a few hundred meters only), in the southwestern part of the Son valley Vindhyans. Considering high Moho temperatures and shallow depth to the melting at 50 km below Jabera basin, it looks plausible that a long sustained thermal anomaly persisted below the Vindhyan basin. This induced magmatism, upwarping and erosion, followed by crustal extension, rifting and sedimentation. Even before the Vindhyan deposition took place, almost entire felsic upper crust seems to have been eroded specially from Damoh-Jabera-Katangi-Jabalpur regions, which were in close vicinity of the NSL.

5.4.1 1.1 Ga Super Plume Interaction Association of higher crustal temperatures consequent to rise in mantle solidus and confirmed signatures of sustained uplift, exhumation, crustal rifting and magmatic extrusions, suggest that this region was under the sustained influence of a super mantle plume. It acted in a widespread area of the peninsular shield around 1.1 Ga (Anil Kumar et al. 1993), part of which has been mapped by Mall et al. (2008). The evidences of positive gravity (Mishra and Rajasekhar 2008) and magnetic anomalies, specially in areas of basic kimberlitic and lamproitic intrusions in Panna and Majhgaon, and dyke swarms in Bundelkhand massif and other areas (Rao et al. 2005), support above interpretation. Not surprisingly, in regions which surround the Vindhyan basin, like Cambay graben, Aravalli-Delhi mobile Belt, Damodar graben and NSL rift zone, the heat flow is high (Fig. 4.8, Chap. 4), and lithosphere is thin. Thus these areas became prone to multiple tectonic reactivations during the Gondwana and Deccan volcanic eruptive periods (Jokhan Ram et al. 1996). Further, since the crustal seismic structure below the Bundelkhand massif, differs considerably from the adjoining Vindhyan basin, Bundelkhand block may not be extending below the Vindhyan basin terrain (Vijay Kumar et al. 2012), as understood earlier. MT studies (Gokarn et al. 2001) further confirm this conjecture.

References Anil Kumar, Padma Kumari VM, Dayal AM, Murthy DSN, Gopalan K (1993) Rb–Sr ages of Proterozoic kimberlites of India: evidence for contemporaneous emplacement. Precambrian Res 62:227–237 Arora BR, Waghmare SY, Mahashabde MV (1995) Geomagnetic depth sounding along the Hirapur– Mandla–Bhandara profile, central India. Mem Geol Soc India 31:519–535 Auden JB (1933) Vindhyan sedimentation in Son valley, Mirzapur district. Mem Geol Surv India 62:141–250

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Azmi RJ, Joshi D, Tewari BN, Joshi MN, Mohan K, Srivastava SS (2007) Age of the Vindhyan super group of central India: an exposition of biochronology vs radiochronology. In: Sinha DK (ed) Micropaleontology: application in stratigraphy and paleo oceanography. Narosa publishing house, New Delhi, pp 29–62 Bengtson S, Belivanova V, Rasmussen B, Whitehouse M (2009) The controversial “Cambrian” fossils of the Vindhyan are real but more than a billion years older. Proc National Acad Sci United states of America 106:7729–7734 Bhattacharya A (1996) Recent advances in Vindhyan geology. Mem Geol Soc India 36:331 Bhattacharya A, Morad S (1993) Proterozoic braided ephemeral fluvial deposits: an example from the Dhandraul sandstone formations of the Kaimur Group, Son valley, Central India. Sed Geol 84:101–114 Blackwell D, Richards M, Stepp P (2010) Final report, Texas geothermal assessment for the 135 corridor east, for Texas state energy conservation office contract CM709. SMU geothermal laboratory, Dallas, Texas, USA, p 78 Chakraborty C (2006) Proterozoic intracontinental basin: the Vindhyan example. J Earth Syst Sci 115:3–22 Chakraborty C, Bhattacharya A (1996) The Vindhyan basin: an overview in the light of current perspectives. Mem Geol Soc India 36:301–312 Christensen NI, Mooney WD (1995) Seismic velocity structure and composition of the continental crust; a global view. J Geophys Res 100:9761–9788 Condie KC (1989) Plate tectonics and crustal evolution. Pergamon, New York, p 476 Das AK, Baruah RM, Bisht SS, Agrawal B (1999) An integrated analysis of late Proterozoic lower Vindhyan sediments for hydrocarbon exploration in western part of Son Valley, Central India. J Geol Soc India 53:239–253 Desai AG, Markwick A, Vaselli O, Downess H (2004) Granulite and pyroxenite xenoliths from the Deccan trap: insight into the nature and composition of the lower lithosphere beneath cratonic India. Lithos 78:263–290 Gass IG, Chapman DS, Pollack HN, Thorpe RS (1978) Geological and geophysical parameters of mid-plate volcanism. Philos Trans R Soc Lond A 288:581–597 Gokarn SG, Rao CK, Singh BP (1995) Crustal structure in southeast Rajasthan using magnetotelluric techniques. Mem Geol Soc India 31:373–381 Gokarn SG, Rao CK, Gupta G, Singh BP, Yamashita M (2001) Deep crustal structure in central India using magnetotelluric studies. Geophys J Int 144:685–694 Gokarn SG, Rao CK, Selvaraj C, Gupta G, Singh BP (2013) Crustal evolution and tectonics of the Archean Bundelkhand craton, Central India. J Geol Soc India 82:455–460 Gregory LC, Meert JG, Pradhan R, Pandit MK, Endale T, Malone SJ (2006) A paleomagnetic and geochronologic study of the Majhgawan kimberlite, India: implications for age of the upper Vindhyan Supergroup. Precamb Res 149:65–75 Jokhan Ram, Shukla SN, Pramanik AG, Varma BK, Chandra G, Murthy MSN (1996) Recent investigations in the Vindhyan basin: implications for the basin tectonics. In: Bhattacharya A (ed) Recent advances in Vindhyan Geology. Mem Geol Soc India, vol 36, 267–286 Julia J, Jagadeesh S, Rai SS, Owens TJ (2009) Deep crustal structure of the Indian shield from joint inversion of P wave receiver functions and Rayleigh wave group velocities: implications for Precambrian crustal evolution. J Geophys Res 114:B10313. https://doi.org/10.1029/ 2008JB006261 Kaila KL, Krishna VG (1992) Deep seismic sounding studies in India and major discoveries. Current Sci 62:117–153 Kaila KL, Reddy PR, Dixit MM, Koteswara Rao P (1985) Crustal structure across Narmada-Son lineament, central India from deep seismic soundings. J Geol Soc India 26:465–480 Kaila KL, Murthy PRK, Mall DM, Dixit MM (1989) The evolution of the Vindhyan basin vis a vis the Narmada- Son lineament, Central India, from deep seismic soundings. Tectonophysics 162:277–289

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Polet J, Anderson DL (1995) Depth extent of cratons as inferred from tomographic studies. Geology 23:205–208 Rajendra Prasad B, Vijaya Rao V (2006) Deep seismic reflection study over the Vindhyans of Rajasthan: implications for geophysical setting of the basin. J Earth Syst Sci 115(1):135–147 Ramakrishnan M, Vaidyanadhan R (2008) Geology of India. Geol Soc India, Bangalore 1:556 Rao JM, Poornachandra Rao GVS, Widdowson M, Kelley SP (2005) Evolution of Proterozoic mafic dyke swarms of the Bundelkhand granite massif, central India. Curr Sci 88(3):502–506 Rasmussen B, Bose PK, Sarkar S, Banerjee S, Fletcher IR, McNaughton NJ (2002) 1.6 Ga U-Pb zircon age for the Chorhat sandstone, lower Vindhyan, India: possible implications for early evolution of animals. Geology 30:103–106 Ray JS (2006) Age of the Vindhyan Super group: a review of recent findings. J Earth Syst Sci 115:140–160 Ray JS, Chakraborty C (2006) Preface. In: Vindhyan Geology-status and perspective. J Earth System Sci 115(1): 1–2 Ray JS, Martin MW, Veizer J, Bowering SA (2002) U-Pb zircon dating and Sr isotope systematics of the Vindhyan Supergroup, India. Geology 30:131–134 Ray JS, Veizer J, Davis WJ (2003a) C, O, Sr and Pb isotope systematics of carbonate sequences of the Vindhyan Super group, India: age, diagenesis, correlations and implications for global events. Precamb Res 121:103–140 Ray L, Senthil Kumar P, Reddy GK, Roy S, Rao GV, Srinivasan R, Rao RUM (2003b) High mantle heat flow in a Precambrian granulite province: evidence from southern India. J Geophys Res 108(B2):2084. https://doi.org/10.1029/2001JB000688 Raza M, Khan A, Khan MS (2009) Origin of late Proterozoic Great Vindhyan basin of north Indian shield: geochemical evidence from mafic volcanic rocks. J Asian Earth Sci 34(6):716–730 Reddy PR, Vijaya Rao V (2013) Seismic images of the continental Moho of the Indian shield. In: Artemieva I, Brown L, Kennett BLN, Thybo H (eds) Moho: 100 years after Andrija Mohorovicic, Special volume on Moho. Tectonophysics, vol 609, pp 217–233 Sarangi S, Gopalan K, Kumar S (2004) Pb-Pb of earliest megascopic, eukaryotic alga bearing Rohtas Formation, Vindhyan Supergroup, India: implications for Precambrian atmospheric oxygen evolution. Precamb Res 132:107–121 Sen G, Bizimis M, Das R, Paul DK, Ray A, Biswas B (2009) Deccan plume, lithosphere rifting, and volcanism in Kutch, India. Earth Planet Sci Lett 277:101–111 Sinha-Roy S, Kirmani IR, Reddy BVR, Sahu RL, Patel SN (1986) Fold pattern of the Vindhyan sequence in relation to the Great Boundary Fault: example from Chittorgarh area, Rajasthan. Quat J Geol Mine Met Soc India 58(4):244–251 Srivastava RP (2006) Design of potential field network and interpretation technique for fractal sources and its application to Jabera-Damoh area of Vindhyan basin, Central India. Ph.D. Thesis, Osmania University, India, p 133 Srivastava BN, Rana MS, Verma NK (1983) Geology and hydrocarbon prospects of the Vindhyan basin. Petroleum Asia J 6:179–189 Srivastava RP, Vedanti N, Dimri VP (2007) Optimum design of a gravity survey network and its application to delineate the Jabera—Damoh structure in the Vindhyan basin, Central India. Pure Appl Geophys 164:1–14. https://doi.org/10.1007/s00024-007-0252-1 Srivastava RP, Vedanti N, Pandey OP, Dimri VP (2009) Detailed gravity studies over Jabera-Damoh Region of the Vindhyan basin (Central India) and crustal evolution. J Geol Soc India 73:715–723 Tewari HC, Krishna VG, Dixit MM, Mall DM, Murty ASN, Kumar P, Sridhar AR, Rao GSP (2002) Seismic structure of the crust and sub-crustal lithosphere across the Narmada zone, project completion report, NGRI-2002-LITHOS-352, p. 1–30 Tripathi P, Parthasarathy G, Ahmad SM, Pandey OP (2012a) Mantle derived fluids in the basement of the Deccan traps: evidence from stable carbon and oxygen isotopes of carbonates from the Killari borehole basement, Maharashtra, India. Int J Earth Sci 101:1385–1395

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Tripathi P, Pandey OP, Rao MVMS, Koti Reddy G (2012b) Elastic properties of amphibolite and granulite facies mid-crustal basement rocks of the Deccan volcanic covered 1993 Latur-Killari earthquake region, Maharashtra (India) and mantle metasomatism. Tectonophysics 554–557:159– 168 Valdiya KS, Bhatia SB, Gaur VK (1980) Geology of Vindhyachal. Hindustan Publishing Corp, New Delhi, p 231 Venkatachala BS, Sharma M, Shujla M (1996) Age and life of the Vindhyans-Facts and conjectures. Mem Geol Soc India 36:137–165 Verma PK (1996) Evolution and age of the great boundary fault of Rajasthan. In: Bhattacharya A (ed) Recent advances in Vindhyan geology. Mem Geol Soc India, vol 36, pp 197–212 Vijay Kumar T, Jagadeesh S, Rai SS (2012) Crustal structure beneath the Archean-Proterozoic terrain of north India from receiver function modeling. J Asian Earth Sci 58:108–118 Windley BF (1977) Evolving continents. Wiley, London, p 385

Chapter 6

Western Continental Margin and Adjacent Oceanic Regions

6.1 Introduction The passive continental margin of western India and adjoining offshore region, is associated with a transitional-type thin crust having semi-continental character. Water column in these areas rarely exceeds 300 m. This margin extends for over 1000 km in length (Fig. 6.1) and assumes special significance in the realm of geodynamic evolution of Indian subcontinent. It has a broad shelf region containing a number of promising hydrocarbon prospects, including Bombay high (currently also known as Mumbai high), which may well be the continuation of Cambay oil/gas field situated further north. This rifted margin witnessed a number of catastrophic and geodynamic events in the last 130 Ma, and can be considered distinctly different from the eastern coast, as well as rest of India. Major geotectonic features of this margin include, Cambay graben, Narmada-Son Lineament, Bombay high offshore region, Khambat Junction, West Coast Boundary Fault and Panvel Flexure Zone. It is dotted by high temperature gradients, high heat flow, number of hot springs and high seismicity including several damaging ones. This region is well marked by intense Reservoir Triggered Seismicity (RTS), which are seen at Bhatsa, Warna and Koyna etc. (Rastogi et al. 1997). All these signatures indicate stressed and degenerated nature of the entire terrain; the crust beneath this region is largely transitional as mentioned before.

6.2 Super Mobility and Geodynamic Events India’s eastern margin lay close to Enderby Land of east Antarctica within the Gondwanaland assembly till early Cretaceous (Lawver and Scotese 1987; Yoshida et al. 2003; Collins and Pisarevsky 2005; Pandey and Agrawal 2008; Chetty 2014). The period which followed thereafter, was an extremely dynamic phase in the Earth’s

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Fig. 6.1 Major geotectonic features of western continental margin of India and adjoining Arabian sea. Bathymetry is in km. Modified after Fig. 1, Pandey et al. (1996)

history, specially around Indian sub continent. Extensive global plate reorganisation, increase in spreading rates and formation of the Indian Ocean took place due to progressive disintegration of the Gondwanaland, around 130 Ma. Subsequent to this, the Indian subcontinent traversed over various mantle hotspots, like Reunion, Kerguelen, Crozet, and Marion in quick succession that arguably led to substantive erosion of almost 150 km of the lithosphere from its bottom (Pandey and Negi 1987a). At around 90 Ma, Madagascar got detached from the India’s west coast due to upwelling of Marion plume (Agrawal et al. 1992; Raval and Veeraswamy 2003). These two events, Indian lithosphere thinning and Madagascar breakup, induced extraordinary mobility (20 cm/year) to the Indian continent, which lasted till 53 Ma when it collided with Eurasia. Prior to India-Madagascar separation, both of them moved slowly (3.5–4.5 cm/year) together. During the post breakup period, several other geodynamic events also took place close to the western margin, like K-T impact,

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Deccan volcanism, Seychelles breakup to name a few (Negi et al. 1992, 1993; Pandey et al. 1995). Even today, many geotectonic segments close to western margin, are experiencing repeated uplifts.

6.2.1 Madagascar Breakup The morphological kinship of Madagascar, which is the largest island in the Indian Ocean, has always been debated till nineteen eighties. In almost every paleo reconstruction, it is shown sandwiched between India and east coast of Africa, however prior to Gondwana breakup, its precise location has remained controversial. Based on magnetic satellite and gravity data, besides geophysical and geotectonic similarities, Agrawal et al. (1992) proposed that in the pre-breakup Gondwana period, Madagascar was a continental fragment of the paleo-super Dharwar craton of India, which was supported by several studies that followed later (Raval and Veeraswamy 2003; Pandey and Agrawal 2008; Tucker et al. 2011; Gibbons et al. 2012; Chetty 2017). Madagascar breakup from the western margin of India, which was a pre-existing mobile belt, took place due to thermal upwelling of the Marion mantle plume, as mentioned before, location of which was quite close to the SE margin of Madagascar (Curray and Munasinghe 1991; Storey et al. 1995; Raval and Veeraswamy 2003). This region in Madagascar, is characterised by several mega shear zones and mobile belts as well, that further extend into Indian subcontinent. Geophysically, this finding is well supported. For example the elliptical shaped splitted nucleus of the Dharwar craton (Radhakrishna and Naqvi 1986), located close to western margin, is brought out very clearly by the long wavelength gravity features (Fig. 6.2). Distribution of gravity anomalies pattern, show an incomplete gravity closure of unusually low magnitude (−50 to −100 mGal) over the Dharwar craton. The incomplete closure with high gradient contours paralleling the wastern coast on Indian, correlate well with the long wavelength contours on the Madagascar side also. Such high gradient contours usually represent rifted continental blocks, possibly uplifted on either side. Similar conclusions can be drawn from the distribution of radially polarized MAGSAT anomalies also (Agrawal et al. 1992). The breakup between these two terrains was accompanied by a large scale emplacement of trachyte, felsite and tuff in the late Cretaceous sediments off Cochin and felsic volcanism at St. Mary island, located close to Malpe, on the southern part of the west coast.

6.2.2 K-T Impact and Deccan Volcanism As per recent studies (Schoene et al. 2015; Renne et al. 2015), the major fraction of the Deccan magma erupted in a quick succession within half a million years across the Cretaceous–Paleocene (K/Pb) boundary (also known as K-T boundary),

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Fig. 6.2 1° × 1° Bouguer gravity anomaly (in mGal) over the south Indian shield (Qureshy and Warsi 1977) and Madagascar (Woollard 1972). Modified after Fig. 3, Agrawal et al. (1992)

or even less (only 10,000 years), as suggested by Courtillot (1990). Palaeomagnetic data firmly put the Deccan volcanic formations in chron 29 R (Schoene et al. 2015). Interestingly, the timing of Deccan volcanism coincided with the K-T boundary asteroidal impacts (like Chicxulub crater, Yucatan Peninsula, Gulf of Mexico and Boltysh crater, Ukraine) and a major biological mass extinction that killed dinosaurs (Pandey and Negi 1987b). Among the mass extinction episodes, the K-T extinction is considered by far one of the most severe, as it marks the end of an era dominated by reptiles. Two possible explanations were initially given for this biological catastrophe (i) asteroidal impact on the surface of the earth (Alvarez et al. 1980), and (ii) intense global volcanism which has a periodicity of 33 Ma (Pandey and Negi 1987b; Rampino and Stothers 1988). Both the explanations have been supported subsequently by different evidences, but debate continues. If one believes that the duration of Deccan volcanic eruption was indeed short, then it would require a powerful triggering for rapid discharge of magma from sub-crustal depths. Since lots of dinosaur remains

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have been found within the Deccan trap itself, there have been also suggestions that the K-T impact site may be hidden below the Deccan Traps (Alvarez et al. 1982; Courtillot et al. 1986; Rampino 1987; Alt et al. 1988; Clube and Napier 1990). However, the exact location of the impact remained elusive as the Deccan basalt flows are quite thick (about 2–3 km) near western margin.

6.2.2.1

Anomalous Gravity High Over Bombay and Detection of Magma Body

Several gravity highs can be noticed in the vicinity of western margin (Fig. 1.18, Chap. 1), including an oval shaped gravity high (about 100 mGal) located between 18° 40 –19° 30 N and 72° 20 –73° 20 E, covering an area of about 10,000 km2 near Bombay and adjacent offshore region (Fig. 6.3) (Negi et al. 1992). Three-dimensional crustal structure derived from forward modelling and inversion of gravity data, along

Fig. 6.3 Residual gravity anomaly (in mGal) over Bombay and surrounding areas. Modified after Fig. 4, Takin (1966)

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Fig. 6.4 2-D model of fossil volcanic conduit over the Moho boundary beneath offshore near Bombay along profile G-G (shown in Fig. 6.3). Modified after Fig. 6, Negi et al. (1992)

profile G-G , which cuts across this gravity anomaly, indicated presence of a fossil conduit structure 12 km in height from the base of a shallowed Moho (~18 km) (Fig. 6.4). Its top is located at the depth of 6 km below surface and it has a diameter of 35 km at its base. The derived model matched well with deep seismic results obtained by Kaila et al. (1981a), who also obtained a Moho depth of about 18 km near Billimora. The derived feature can be interpreted as a massive intrusion, a mantle plume conduit or structure caused by asteroidal impact. Intrusions of such size are rare and have so far not been reported around western margin. Similarly, deep signatures of the Deccan mantle plume is found below NW India (Kennett and Widiyantoro 1999) but not close to western margin. Thus, Negi et al. (1992) suggested that this structure may have been caused by a large bolide impact, which teared the Earth’s crust, and triggered the eruption of Deccan Volcanics on the western margin of the fast moving Indian plate. It also led to detachment of Seychelles block from western margin and reorganized plate boundaries in the Indian ocean (Fig. 6.5), including the initiation of the Laxmi ridge in the eastern basin of the Arabian sea. The derived magmatic feature (Fig. 6.4), was later on named as ‘NAPSI’ structure by Chatterjee and Rudra (1996) based on the proposers last names. This postulation was later on supported by detailed studies carried out by Chatterjee (1997), Chatterjee and Rudra (1996) and Chatterjee et al. (2003, 2006).

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Fig. 6.5 Initiation of rifting between Seychells and western margin of India and reorganisation of plate boundaries in the Arabian sea, due to K-T asteroidal impact offshore near Bombay. Extent of Deccan Traps is also shown by dotted envelope. CG: Cambay graben, AFB: Aravalli Fold Belt, NSL: Narmada Son Lineament, KO: Koyna rift, KU: Kurduwadi rift, Star: Khambat junction

6.2.2.2

Shiva Impact Crater

The study carried out by Chatterjee and his co-workers, as referred above, defined the boundaries of the crater, based on several geoscientific studies and termed this structure as Shiva crater. They indicated that the crater area affected by the asteroidal impact, is much larger, having a diameter of about 400 km, due to obliqueness and downrange direction of the impact trajectory (Fig. 6.6). The impacted region is surrounded by intersecting set of boundary faults (Chatterjee et al. 2006; Vasudevan et al. 2012). Northern border is bounded by the offshore

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Fig. 6.6 Present day location of two halves of the shiva crater with reference to India and Seychelles, based on Chatterjee and Rudra (1996). Seychelles was an integral part of India before the impact led initiation of the Carlsberg Ridge and Deccan volcanism

Narmada Fault (Chatterjee and Rudra 2008), western border by the Kori Arch, southern boundary by Ratnagiri fault, while the eastern rim of this crater is represented by Panvel Flexure Zone (Chatterjee et al. 2006), which is marked by a line of hot springs, dikes and deep crustal faults (Kaila et al. 1981b). The area covered by this crater includes, Bombay structural high, Mukta and Bassein highs and Panna-Heera blocks in the offshore region, apart from some parts of Surat, Saurashtra, Murad and shelf margin basins. Tertiary basins lying western side of the NAPSI structure contain thick-strata of post-impact Tertiary sediments. All of these basins are rich in hydrocarbon occurrences, out of which Bombay high is considered a giant oil/gas field. Structure of Shiva crater is well reflected in the gravity field (Srivastava 1996;

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Bhattacharyya et al. 2009). Southern half of the impact structure now lies near the Amirante basin, close to Seychelles island (Fig. 6.6).

6.2.2.3

Occurrence of Ir Anomaly and High Pressure-Temperature Fullerene

The volcano-sedimentary K-T boundary section at Anjar (Gujarat), located at the western periphery of the Deccan Volcanic Province, is one of the most thoroughly studied section in India that consists nine lava flows and about four intertrappean beds (Bhandari et al. 1995). 6 m thick intertrappean bed, between Deccan flows 3 and 4, contains several cometry signatures like iridium and osmium (Bhandari et al. 1996; Courtillot et al. 2000). Besides, it also contains natural, toluene-soluble C60 and the toluene-insoluble high-pressure and temperature (HPT) phase of fullerene C60 in the carbonaceous matter, extracted from the iridium-rich layers of the intertrappean sediments (Parthasarathy et al. 2002, 2008). Location of Ir anomalies and Fullerene occurrences (Br-1, Br-2, and Br-3) in the K-T boundary intertrappean beds at Anjar (Gujarat) are shown in Fig. 6.7. The presence of HPT fullerene in the K-T boundary layer at Anjar, has been linked to this impact event. From paleontologic evidences, Br-1 coincides with the KT boundary (Chatterjee et al. 2006) (Fig. 6.7). Three closely spaced iridium and fullerene spikes may favour multiple impacts around this boundary.

6.3 Thermal Structure The asteroidal impact caused large scale structural uplift and severely affected the region lying between 17° and 21° N and 69°–74° E (Fig. 6.1). It also modified the thermal and crustal structure underneath. In the impact affected region, Moho is extremely shallow at around 16–30 km and geothermal gradients as well as heat flow are quite high. Measured temperature gradients (Fig. 6.8a) fall in the range of 29–78 °C/km in the Mumbai offshore (Pande et al. 1984; Pandey and Agrawal 2000, 2001). Similarly at Ganeshpuri in the Konkan region, measured heat flow is very high at 97 mW/m2 (Fig. 6.8b). There is no direct measurement of heat flow in the offshore Bombay region, but based on temperature gradients and assumed conductivity, a high heat flow of about 83 mW/m2 is estimated for the northern and eastern part of the Bombay offshore region, while lower value of 56 mW/m2 in its southwestern part (Fig. 6.8b). Further, below this region, measured subsurface temperature reach to about 110–260 °C at 3 km depth (Pandey and Agrawal 2000).

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Fig. 6.7 K-T boundary section at Anjar, showing occurrences of iridium-rich layers (Br-1, Br-2, Br-3) and fullerenes, related to K-T impact. KTB: Reported location of K-T boundary. Modified after Fig. 1, Parthasarathy et al. (2002)

6.3.1 Cambay Basin It is a narrow intra-cratonic hydrocarbon bearing graben structure, located in alluvial plains of Gujarat (Fig. 6.1). This basin, formed by discontinuous normal faults, extends between latitude 21 °N–25 °N and longitude 71° 30 E–73° 45 E and runs for about 425 km, covering an approximate area of 56,000 km2 (Mohan et al. 2008). It trends in NNW-SSE direction before taking a swing towards NNE-SSW direction near the Gulf of Cambay. In all likelihood, it appears to join Bombay offshore basins, further south. Narmada rift zone, which borders this basin in the south, is seemingly related to the development of this intra-cratonic rift structure, which contains as much as 6 km thick Tertiary sediments that overly Deccan volcanic basement. The geology and tectonics of this graben has been discussed by many workers like Sen Gupta (1967), Mathur et al. (1968) and Biswas (1987). Inspite of thick sedimentation in the basin, the gravity anomalies are unexpectedly positive compared to those in surrounding areas.

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Fig. 6.8 a Geothermal gradient distribution (°C/km) in the northern part of western continental margin and adjacent areas (Gupta 1981; Gupta et al. 1988; Pande et al. 1984). I: north Cambay graben, II: south Cambay graben, III: Konkan thermal province, IV: northern and eastern Mumbai offshore, V: southwestern Mumbai offshore. b Distribution of heat flow (in mW/m2 ) over the same regions (Gupta 1981; Gupta et al. 1988; Negi et al. 1992; Vedanti et al. 2011). I-V refer to same segments as in a. VI: Bhuj earthquake region. Modified after Figs. 3 and 4, Pandey and Agrawal (2001)

The measured heat flow (Fig. 6.8b) in the basin ranges between 55 and 93 mW/m2 (Gupta 1981), with an averages of about 83 and 61 mW/m2 in the northern and southern parts respectively. Similarly, the temperature gradients are equally high (Fig. 6.8a) reaching above 50 °C/km in several areas (Gupta and Rao 1970; Gupta 1981). Measured in situ temperatures in some of the wells till about 1200 m depth are shown in Fig. 6.9. In the northern part, subsurface temperatures are expected to reach as high as 150–200 °C at 3 km depth. An important aspect of this basin has been that the original crystalline crust, which remains now, is only about 17.5 km thick, due

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Fig. 6.9 Temperature-depth distribution in some of the boreholes in Cambay graben (Gupta and Rao 1970) 1: Nawagam, 2: Cambay, 3: Kathna, 4: Ankleshwar, 5: Kalol. After Fig. 6, Pandey et al. (2017)

to shallow sub-crustal melting and exhumation-led denudation. Consequently, 8 km thick crustal magma underplating is seen in its northern part, while an unprecedented 22 km thick in its southern part (Kaila et al. 1981a), which may be one of the highest recorded anywhere in the global graben structures. Using appropriate crustal heat production model, as given in Table 6.1 and discussed in detail in Pandey et al. (2017), temperature-depth variation was estimated (Fig. 6.10), which resulted into an extremely high mantle heat flow (65 mW/m2 ) and Moho temperatures (860 °C), with a thin lithosphere of only about 45 km in the northern part, indicating occurrence of melting conditions just 12 km below Moho. In comparison, mantle heat flow and Moho temperatures are relatively lower at 42 mW/m2 and 675 °C respectively in the southern part, where lithosphere could be about 70 km thick. There is a high possibility that the observed thermal anomaly in the Cambay basin may continue further south in the oil/gas bearing Mumbai offshore region, as both of them are geotectonically similar and have a comparable thermal history (Pandey and Agrawal 2000, 2001).

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Table 6.1 Crustal heat production models and adopted geothermal parameters for temperaturedepth estimation beneath northern and southern part of the Cambay graben (Gujarat), India Depth range (km)

Rock type

Heat production (µW/m3 )

Thermal conductivity (W/m °C)

North Cambay graben

Surface heat flow: 83 mW/m2

0–5.0

Sediments

1.5

3.5

5.0–6.5

Deccan volcanics

0.02

1.7

6.5–10.0

Granite-gneiss

1.82

3.0

10.0–12.0

Low velocity mafic rocks

0.78

2.5

12.0–24.0

Granulites

0.16

2.5

24.0–32.0

Magmatic crust

0.02

2.6

>32.0

Ultramafic mantle

0.01

3.0

Data source Gupta (1981), Kaila et al. (1989, 1990), Tewari et al. (1991), Liu and Zoback (1997), Roy and Rao (1999), Ray et al. (2003) and Pandey et al. (2017) South Cambay graben

Surface heat flow: 61 mW/m2

0–2.0

Sediments

1.5

3.5

2.0–3.0

Deccan volcanics

0.02

1.7

3.0–5.0

Mesozoic sediments

1.5

3.5

5.0–9.0

Granite-gneiss

1.82

3.0

9.0–16.0

Amphibolite-granulite

0.78

2.5

16.0–38.0

Magmatic crust

0.02

2.6

>38.0

Ultramafic mantle

0.01

3.0

Data source Gupta (1981), Kaila et al. (1981a), Liu and Zoback (1997), Roy and Rao (1999), Ray et al. (2003) and Pandey et al. (2017)

Recently, Ganguli et al. (2018) analysed the bottom hole temperatures recorded at differen depths in the Ankleshwar oil field, and found higher temperature gradients (~53 °C/km) at deeper levels. Based on this, a heat flow of 75.2 mW/m2 was estimated at Ankleshwar, together with high Moho temperatures of about 880 °C and a shallow depth to subcrustal melting (~50 km). Earlier reported heat flow at the Ankleshwar was 67 mW/m2 . This inference may suggest that the entire hydrocarbon bearing Cambay basin (including southern part), along with the Mumbai offshore region, is associated with high heat flow and shallow melting conditions.

6.3.2 Geothermal Springs Some 60 thermal springs with temperature ranging from 34 to 71 °C, clustered around 18 different places, are located in the Konkan area (Saxena and Gupta 1987). In fact chain of such springs spread in a narrow belt all along the western coast between

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Fig. 6.10 Estimated temperature–depth distribution below northern and southern part of the Cambay graben based on crustal heat generation model, shown in Table 6.1. This figure indicates extremely shallow depth to the melting in northern part of the Cambay basin

16° and 24° latitude (Fig. 6.11). A large number of them are located between the Western Ghats and the west coast, which are mostly associated with local tectonic movements. Many of them occur close to the borders of the dolerite dykes (Gupta et al. 1975). Thermal manifestations are strongest at Unhavre (Khed) area, where the spring temperature reaches 71 °C. All the tectonic trends in the region as well as a line of hot springs are oriented north-south.

6.3.3 West Coast Thermal Anomaly Zone As discussed earlier, geothermal gradients and heat flow values are much high in (i) the Northern part of the Cambay graben, (ii) Konkan geothermal province, and (iii) in northern and eastern parts of the Bombay offshore region. Pandey and Negi (1995) and Pandey and Agrawal (2000, 2001) broadly named this region as “West Coast Thermal Anomaly Zone” (Fig. 6.11). Beneath this region, the estimated temperatures are anomalously high and asthenosphere is located at a shallower depth of 40–50 km (Pandey and Agrawal 2001; Pandey et al. 2017), indicating presence of melting conditions, at a few kilometre below the Moho.

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Fig. 6.11 Heat flow (q, in mW/m2 ), geothermal gradients (G, in °C/km) and lithospheric thickness (LT, in km) beneath geotectonic segments of West Coast Thermal Anomaly Zone. Location of thermal springs are shown by solid dots (Ravi Shankar et al. 1991). Location of DSS profiles TD (Tharad-Degam) and MB (Mehmadabad-Billimora) are also shown

6.4 Gravity Field, MT Studies and Seismicity Entire continental margin, infused by Deccan magma, is characterized by linear positive Bouguer anomalies paralleling the coast (Fig. 6.12), reflecting crustal thinning and asthenospheric upwarping underneath (Qureshy 1981). This is specially true for the areas located north of 17° latitude. The positive gravity trend seems to continue further north along the Cambay graben, which is also associated with high heat flow anomaly and thin lithosphere, as mentioned before. Lateral gradients seen in

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Fig. 6.12 Bouguer gravity anomaly map of western India (in mGal, 1:5000,000 scale) adopted from 1975 edition of the Bouguer gravity anomaly map, published by CSIR-NGRI, Hyderabad (India)

gravity anomalies conforms well with the extensively stretched, fractured, rifted and faulted nature of the western continental margin, dotted by numerous geothermal springs, which show close association with tectonic movements. Telluric and magnetotelluric studies carried out in the northern part of the Konkan region indicated presence of two distinct conductors (Harinarayana and Sarma 1991), one at a depth of 400 m (20 m resistivity) and the other, at 1500 m (4 m resistivity), implying

6.4 Gravity Field, MT Studies and Seismicity

183

Fig. 6.13 Distribution of earthquake epicentres (1700–1997) near western continental margin of India prepared from Ramlingeswara Rao (2000). Modified after Fig. 7, Pandey and Agrawal (2001)

presence of hot water zone. In Ganeshpuri, which is associated with high heat flow of 97 mW/m2 , the temperature of hot water discharge is 42–59 °C. Therefore it is not surprising that the entire region is seismically active (Fig. 6.13), having witnessed several earthquakes of large magnitude in the past. High temperature regime have also played an important role in enhancing the thermal maturation process in oil and gas-rich Tertiary sediments. It is felt that the underlying mantle below this region is hot, less viscous, fertile and buoyant, which caused rise of isotherms and consequent lithospheric mantle deformation, Moho shortening, magmatic underplating, exhumation of high velocity mafic crust besides almost total erosion of granitic-gneissic upper crust. Infact, in Heera block of Mumbai offshore region, granulitic crust is exposed at the basement level itself.

6.5 Deep Crustal Seismic Studies 6.5.1 DSS Studies In order to study detailed crustal structure of the Cambay basin, DSS studies were carried out along a 220-km long N-S profile from Mehmadabad (22° 50 N 72° 46 E) to Billimora (20° 46 N 72° 58 E) and both refraction as well as narrow and wide angle reflection travel time data were obtained (Kaila et al. 1981a). In the region

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Fig. 6.14 Simplified crustal seismic section along Mehmadabad-Billimora traverse, derived from deep seismic sounding studies carried out by Kaila et al. (1981a) and Kaila (1986). Layer velocities are in km/s. Modified after Fig. 7, Negi et al. (1992)

located south of this graben, Moho as well as the granitic basement was recorded at much shallower depths. The study revealed that near Billimora (between Surat and Bombay), mafic crust lies at a depth of 6 km, and Moho at an extremely shallow depth of 18 km (Kaila 1986) (Fig. 6.14) which conforms with the spectral estimates made by Negi et al. (1992). Upwarping of the Moho coincides with the linear gravity anomalies. Further, a high Vp (7.26 km/s) is also detected at a shallow depth of about 10 km below the Ambica river close to Billimora, which may represent the top of the thickly underplated magma layer underneath. The entire region lying all along the profile is underplated by 15–20 km thick magma in which Vp exceed 7.0 km/s. This inference is consistent with the impact-led magma generation and its consequent stacking below the lower crust, further imploying strong crust–mantle interaction below this area.

6.5.2 Receiver Function Studies Tiwari et al. (2006) deployed three broadband and five short period stations in the northern part of the Western Ghats (Fig. 6.15). Using receiver function studies, average crustal thickness below two broadband stations, MULG and VARE, was found to be about 36 and 35 km respectively with Poissons ratio around 0.26 for both the locations. Almost similar results were obtained for the third broadband station MPAD. However, the short period stations (HATT and JAYD), located over the northernmost part of the Western Ghats, revealed 3–4 km thinner crust to around 30 km.

6.6 Impact Induced Rifting and Deccan Volcanism

185

Fig. 6.15 Physiographic map of Western Ghats region, showing locations of seismic stations, with red and black symbols denoting broadband and short period stations respectively. KOYN: Koyna seismic zone. Modified after Fig. 2, Tiwari et al. (2006)

6.6 Impact Induced Rifting and Deccan Volcanism The area covering anomalous Bombay offshore region and its neighbourhood, characterized by anomalous thermal field, crustal seismic structure, positive gravity and thinner lithosphere, is quite small, compared to the area normally covered by deep plume source (White and McKenzie 1989; Campbell and Griffiths 1990). Thus, the mushrooming plume head beneath this region could not have been supported by the crust, or even lithosphere. A deeper mantle plume should manifest in a far greater area, in terms of thermal anomalies and gravity field. Further, the highly differentiated nature of the Deccan basalts, erupted from crust-mantle boundary (Pandey and Negi 1987c; Sen 1988; Sen and Chandrasekharam 2011) would not support it. So, there is a fair possibility that magma erupted through a shallow secondary plume, signatures of which are present in the Bombay offshore region (Fig. 6.4), which would conform with the highly differentiated nature of the Deccan magma also. It is now understood that the breakup of Gondwanaland and super mobility of India during 90 and 53 Ma, led to the loss of almost 150 km of the lithosphere from its bottom (Pandey and Negi 1987a; Kumar et al. 2007, 2013; Pandey 2016; Pandey et al. 2017). Such process should have generated substantial melt at the base of the lithosphere due to frictional heating and cause rise of mantle solidus, subcrustal

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erosion and secondary lithospheric melting. If we accept that the magma eruptive phase was indeed short, then triggering of the rapid eruption of the such accumulated magma by the asteroidal impact, could be a distinct possibility. This was the time when the Earth was passing through a dynamic phase, with extensive global plate reorientation, abrupt increase in spreading rates and micro-continental rifting like Seychelles breakup, which caused considerable stress build-up and weakening of the crust and lithosphere. The western margin of India was part and parcel of this scenario. The K-T impact shattered the entire western Indian region and rejuvenated pre-existing zone of weaknesses within a supposedly rigid block, like Koyna and Kurduvadi rifts (Krishna Brahmam and Negi 1973), which are presently associated with moderate seismic activity (Pandey 2009). Since the area shattered by the K-T impact was quite large, at around 400 km in diameter, it is envisaged that the impacted asteroid probably had a diameter of 10 km or more, with a depth of penetration exceeding 20 km, thereby facilitating decompression melting and triggering a rapid uplift and seepage of magma to the surface through a conduit pipe. The hydrodynamic modeling indicates that the total melting volume increases substantially due to decreasing pressure (i.e., decompression melting) beneath a large impact site (Jones et al. 2002). In fact, it could lead to 10–30% additional melting of mafic/ultramafic rocks, thus triggering large-scale flood basaltic volcanism. In case of Deccan basalts, the uplift and mantle upwelling, accompanying the collapse of the offshore Mumbai, might have resulted into pressure-release melting, thereby creating a large igneous province, like Deccan traps. As per Pandey and Negi (1987b) and Sobolev et al. (2011), large impacts do accelerate volcanism and cause geomagnetic reversals too (Muller and Morris 1986). The elliptical nature of the impact site would suggest that the bolide hit the Bombay offshore region at an oblique angle, causing large-scale crustal deformation and reactivation of deep-seated faults, fractures, rift structures, etc. (Chandrasekharam 1985), as indicated by the thermal springs and seismically active zones around the continental margin. These weak zones appear to have facilitated rapid magma movement, resulting in the build-up of a huge pile of tholeiitic basalts over western India in a short duration. A schematic diagram showing lithospheric structure across the western margin that passes through the proposed impact structure, is shown in Fig. 6.16

6.7 Seychelles Breakup and Initiation of Laxmi and Calsberg Ridges Apart from triggering the Deccan volcanic eruption, K-T impact also led to the initiation of Carlsberg and Laxmi ridges in the Indian Ocean. An impact in this locality can easily explain the sudden detachment of the arcuate Seychelles block from the west coast of India, as well as the large-scale reorganisation of plate boundaries in the Indian Ocean, as mentioned before. Pandey et al. (1995) postulated that horizontal compressive regime from either side (rifting associated with Deccan volcanism on

6.7 Seychelles Breakup and Initiation of Laxmi and Calsberg Ridges

187

Fig. 6.16 Crustal and thermal structure of the lithosphere along profile AB (Fig. 6.5), running across the K-T impact site near offshore Mumbai

the northeast and rapid spreading across the Carlsberg ridge in the southwest), facilitated buckling up of the lithosphere, leading to the formation of the Laxmi Ridge. There are also suggestions, based on sea floor magnetic lineations (Naini 1980) that the ridge would be slightly younger than the Deccan volcanic event. Collier et al. (2008) mentioned that the severing of the Seychelles possibly took place around 62 Ma, while the initiation of seafloor spreading at the Carlsberg Ridge little earlier at around 63.4 Ma, based on seafloor magnetic anomaly modeling and seismic results.

6.7.1 Laxmi Ridge Evolution As mentioned in earlier sections, the Arabian Sea floor, located between the Arabian abyssal plain and India’s western continental margin, contains a number of prominent elongated features like Chagos-Laccadive, Prathap, and Laxmi ridges and uplifted features like Bombay and Kori highs (Fig. 6.1). All these features developed between 80 and 53 Ma, after the breakup of Madagascar from the western coast of India (Agrawal et al. 1992). This was also the period, when many plate tectonic changes and plate reorganizations were taking place, including the ridge jump of greater than

188

6 Western Continental Margin and Adjacent Oceanic Regions

500 km in the northern Mascarene Basin, Deccan flood volcanism on western margin, and initiation of the northern Carlsberg ridge as well as Laxmi Ridge (Hartnady 1986). Out of these structural features, the geophysically unusual Laxmi Ridge (named by Naini and Talwani 1977), located in the eastern basin of the north Arabian Sea, remains an enigmatic feature, which is associated with a prominent elongated negative gravity anomaly. This submarine ridge, between 14° and 19° N latitude and 64°–69 °E longitude, is about 150 km wide and trends NW-SE. Naini (1980), possibly first indicated that the Laxmi ridge is underlain by a thick crustal layer above the Moho, with an anomalous Vp 7.2 km/s, which is rather unusual for a normal ocean basin crust. It has a basement relief of up to 2 km and Moho is at the depth of about 21 km, somewhat similar to some aseismic ridges, ocean rises, and plateaus, and mantle plume generated areas, like Iceland and Hawaii, but they are known to be intrinsically associated with positive gravity anomalies (Malahoff 1969; Babu Rao 1970). In contrast, the gravity anomalies are negative (~30–50 mGal) over this ridge (Fig. 6.17), which is also characterized by a negative Airy (T = 30 km) isostatic anomaly of about 30 mGal (Naini and Talwani 1982). Later on, Pandey et al. (1995)

Fig. 6.17 Long-wavelength Free-air gravity anomaly (in mGal) over the Laxmi Ridge and surrounding regions (Naini and Talwani 1982). AB is the profile along which the lithospheric model is derived, as shown in Fig. 6.18. Area covered by spectral analysis is shown by dashed lines. Modified after Fig. 3, Pandey et al. (1995)

6.7 Seychelles Breakup and Initiation of Laxmi and Calsberg Ridges

189

analysed the gravity field and discussed its evolutionary history in the light of the late Cretaceous geodynamic history of the region.

6.7.1.1

Gravity Field Over Laxmi Ridge

The free-air gravity anomaly map of the Laxmi Ridge and its surrounding region (Naini and Talwani 1982), is shown in Fig. 6.17. It is largely negative over the entire Arabian Sea (Naini 1980), with some high and low anomalies superimposed on it. This ridge and its surroundings areas are characterized by long-wavelength linear anomalies (+25 to −50 mGal) oriented in NW-SE direction. The anomalies to the west of the Laxmi Ridge are, however, of diffused nature, characterising normal oceanic crust, devoid of any appreciable features. Calculated radial spectrum of the gravity anomaly over this ridge, provided source depths at 11 and 34 km below the sea surface, which correspond to Moho and base of the lithosphere respectively. The crust is of normal oceanic type, which is underlain by an upwarped asthenosphere at 35 km depth. The residual gravity field by subtracting the regional field and the gravity effect of different layers from the observed gravity field, reflects the undulations in the sub-crustal transitional layer. Pandey et al. (1995) further carried out detailed crustal structure modelling, across two profiles, approximately 400 km in length, one profile AB, that cuts across the central part of the ridge axis, is shown in Fig. 6.18, which reveals presence of a normal oceanic lithosphere (with no transitional layer) to the west of the Laxmi Ridge and about 100–150 km NE of it. They hypothesized a thick, underplated low-density layer (2.95 g/cm3 ) underlies the Moho at depth of 11 km and extends down to 19 km depth. This low density layer could be the probable cause of the gravity low over the Laxmi ridge. Naini (1980), however, reported that the crustal thickness below this ridge is about twice that of the normal oceanic crust. The negative gravity anomaly over the Laxmi Ridge and Vp 7.2 km/s velocity (which is characteristic of underplated magma layer in continental regions) are rather unusual for a typical ocean basin crust. Usually, velocities in oceanic lower crustal layers vary between 6.5 and 6.9 km/s. Although, in some active volcanic regions, such anomalous velocities can be found, for example, in Iceland and Hawaii (Flovenz and Gunnarsson 1991), but, they are associated with hotter upper mantle. That is, however, not the case with the Laxmi ridge. Although the asthenosphere below this ridge is shallow, as mentioned earlier, it possibly did not act as a spreading centre. In the absence of the active rifting, the negative gravity anomaly may be possible only with emplacement of some low-density material below the crust, which could be either in the form of serpentinized olivine basalt/peridotite or low-density fractionated magma, related in some way or other to the Deccan volcanic episode. This conjecture was supported by Miles et al. (1998) based on various studies. Occurrences of such magmas are not uncommon. Such a low density underplated material, related to Deccan volcanism, is known to be present below crustal and subcrustal depths in various geotectonic segments of the western India, including offshore regions (Kaila et al. 1990; Krishna et al. 1991; Singh et al. 1991; Pandey et al. 1996; Vasanthi and

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Fig. 6.18 Lithospheric model (with density variations in the g/cm3 ) along profile AB, location of which is shown in Fig. 6.17. Dotted area represents low density (2.95 g/cm3 ) subcrustal underplated magma layer. Modified after Fig. 6, Pandey et al. (1995)

Satish Kumar 2016). The Deccan magmas are indeed highly differentiated and have low Mg ratios (Pandey and Negi 1987c). Miles et al. (1998) reported some evidences to support the postulation made by Pandey et al. (1995) as mentioned before. Similarly, Singh (1999) and Rajaram et al. (2011) supported its oceanic affinity. Krishna et al. (2006), however, based on seismics, gravity and magnetic data, indicated that both, the Laxmi and Panikkar ridges, are fragments of stretched continental crust developed under extension and riftrelated regimes, which was further supported by Kolla and Coumes (1990), Miles and Roest (1993) and Arora et al (2012). Nair et al. (2015) supported magma underplating below the Laxmi ridge, but advocated in favour of the continental crust.

6.7 Seychelles Breakup and Initiation of Laxmi and Calsberg Ridges

191

6.7.2 Crust-Mantle Structure in Arabian Sea and Western Continental Region Lithosphere beneath this region, seems to have suffered maximum post break-up changes among all other Gondwanic continents (Pandey and Negi 1987a; Pandey and Agrawal 1999; Agrawal and Pandey 2002; Krishna et al. 2006; Kumar et al. 2007, 2013; Pandey 2016). Pandey et al. (1996) analysed 1° × 1° Bouguer gravity anomaly map of the western India and the adjacent Arabian sea, adapted from Woollard et al. (1969) and Geological-Geophysical Atlas of the Indian Ocean (UNESCO 1975), respectively (Fig. 6.19). It was subjected to Bouguer correction using the infinite slab formula (P = 2.67 g/cm3 ). High frequencies are automatically filtered, thereby giving rise to long-wavelength anomalies, which signify deep sources. While the contour interval is interpolated at 20 mgal in the oceanic region, it is 10 mGal in the

Fig. 6.19 Bouguer gravity anomaly map in mGal of the Indian subcontinent (1° × 1°) (Woollard et al. 1969) and adjacent Arabian Sea (UNESCO 1975). A-A shows the location of lithospheric profile as shown in Fig. 6.20. Modified after Fig. 3, Pandey et al. (1996)

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continental region. The map, with fairly good continuity of the gravity field in ocean as well as continent, depicts large variation. Two 1700 km long gravity profiles that extend for about 700 km in the oceanic region and 1000 km in the continental region, were chosen for investigating crustmantle characteristics, one of which (A-A ) passed through the impact site near offshore Bombay (Fig. 6.20), which suggests presence of about 9 km thick low density layer (~2.95 g/cm3 ) below the Moho in the Eastern basin of the Arabian sea at depths between 11 and 20 km. The thickness of this layer increases to about 15 km beneath western continental region, where it is sandwiched between 45 and 60 km depth in the mantle lithosphere. A low-velocity layer at similar depth is reported below the Koyna region of the Deccan Traps (Krishna et al. 1991) (Fig. 6.21), where

Fig. 6.20 Lithospheric model along Profile A-A , location of which is shown in Fig. 6.19. Dots represent the region characterised by low-density sub-crustal underplated magma layer. Densities of different layers are in g/cm3 . After Fig. 4, Pandey et al. (1996)

6.7 Seychelles Breakup and Initiation of Laxmi and Calsberg Ridges

193

Fig. 6.21 Travel-time model of the sub-crustal lithosphere beneath Koyna region of Deccan Traps, based on DSS studies. After Fig. 8, Krishna et al. (1991)

the velocity drop is almost 10% (from 8.3 to 7.4 km/s), that would correspond to a lower density of about 0.3 g/cm3 . A layer with similar velocity (7.2–7.4 km/s) and density (3.15 g/cm3 ) also occurs below the Cambay graben at depth between 24 and 32 km along Tharad-Degam DSS profile (Kaila et al. 1990) (location shown in Fig. 6.11). Such low velocity layers are present in many other geotectonic segments of India (Behera et al. 2004; Satyavani et al. 2004; Tewari et al. 2002; Murty et al. 2004 etc.)

6.7.2.1

Characteristics of Anomalous Density Layer

The estimated densities (2.95–3.05 g/cm3 ) and velocities (7.2–7.4 km/s) that characterise the anomalous layers, are often been seen below many active rift-graben systems world-wide (e.g. Ervin and McGinnis 1975; Mooney et al. 1983; Kaila et al. 1990; Mickus and Keller 1992). These layers are related to injection of high-density magmatic material from the mantle beneath stretched lithosphere. Such regions are associated with high gravity anomalies in contrast to negative anomalies over the Laxmi ridge. Negi et al. (1989) reported that more than two-thirds of the Indian subcontinent is underlain by low-density material. Since the asthenosphere beneath Laxmi ridge has risen to a shallower depth of about 35 km, it would indicate lithospheric erosion beneath the Eastern basin, which conforms with the observation of severe degeneration and thinning of Indian lithosphere (Negi et al.1986; Polet and Anderson 1995; Kumar et al. 2007, 2013). Geochemically, this postulation has been found viable. Ellam (1992) indicated that the lithospheric thinning process might have been accelerated prior to Deccan Trap eruption, which allowed partial melting of the lithospheric mantle at a lower depth range of tholeiitic magma generation. This may be a viable reason, why Deccan volcanic magma is so differentiated.

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There could be so many other factors for low density and low velocity layers, but mantle metasomatism and serpentinization appear far more reasonable (Pandey et al. 1995, 1996, 2016; Vasanthi and Satish Kumar 2016; Vedanti et al. 2018). Mantle metasomatism is found to be very common in Deccan volcanic basement (Pandey et al. 2014, 2016) as discussed in Chap. 7 on petrophysical properties. Serpentinisation is often seen in basic and ultrabasic rocks of oceanic environment, where hightemperature primary ferromagnesian minerals are converted to low-temperature, secondary, serpentine-group minerals. Serpentinization is predominant in olivine-rich rock, which induces reduction in density and a decrease in compressional wave velocity. Some island flows close to Bombay exhibited partial or complete olivine alteration at the surface (Deshmukh 1984; Subba Raju et al. 1992). Considering all these aspects, it is felt that the low-density layer could either be (i) serpentinized olivine basalt-peridotite or (ii) low-density fractionated magma in a solid state, related to Deccan volcanism, which ultimately originated from a shallow depth and got substantially fractionated before extrusion and underplating.

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Chapter 7

Seismic, Elastic and Petrophysical Properties of Crustal Rocks: Deccan Volcanic Province

7.1 Introduction The knowledge of seismic, elastic and physical properties of the deep-seated rocks, like density, porosity, P- and S-wave velocities, elastic modulii and seismic wave attenuation, is extremely important to study various geophysical and geodynamical problems. The major part of the Earth’s continental crust, remains inaccessible to direct observations. Even the deepest drilling attempts like Kola peninsula Super Deep Borehole, could not reach beyond 12.262 km (Ault 2015). Fragments carried out from great depths to the surface, like crustal and mantle xenoliths, did provide some useful information. However, our main understanding about the nature of such properties improved only after the advent of seismic and seismological measurements, particularly the deep crustal seismic refraction and reflection profiles. They provided an insight into subsurface velocity-depth distribution, to infer deep-seated petrological and compositional nature. Some important treatise on the subject can be found in Rudnick and Fountain (1995) and Christensen and Mooney (1995). Nevertheless, even the seismic methods are loaded with lots of assumptions and uncertainities. Therefore, the laboratory measured data on rock properties becomes important. They have been widely used to model the crustal composition, rheological behaviour and physical and thermal state of the crust and the underlying mantle lithosphere. Due to relative absence of such informations, even the nature and bulk composition of the Earth’s crust remain poorly understood. This is specially true for the areas which have been geodynamically active, like Deccan Volcanic Province of India, populary known as Deccan traps, basement of which is covered by 1–2 km thick suite of basaltic lavas. Deccan volcanic region forms one of the prominent large Igneous Provinces (LIP) on the surface of the Earth (Fig. 7.1), covering almost 20% (or nearly half a million km2 ) of the Indian landmass, both onshore and offshore, across the western margin. This terrain, which has long been considered as a stable intra-plate region, is experiencing much seismic activity and some events are even disastrous in nature (Pandey 2009; Ramalingeswara Rao 2000); for example, the 1993 Killari earthquake © Springer Nature Switzerland AG 2020 O. P. Pandey, Geodynamic Evolution of the Indian Shield: Geophysical Aspects, Society of Earth Scientists Series, https://doi.org/10.1007/978-3-030-40597-7_7

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Fig. 7.1 Tectonic map of the Deccan Volcanic Province (shaded) covering western and central part of India. Locations of 1993 Killari and 1967 Koyna earthquake, together with prominent rift valleys and faults, are also shown

(Mw 6.2), 2001 Bhuj earthquake (Mw 7.7), 1997 Jabalpur earthquake (Mw 5.8) and the 1967 Koyna (Mw 6.3) earthquake. Recurring seismic activity is still continuing since decades, in places like Koyna and Bhuj. It is thus apparent that the crust–mantle compositional nature, as influenced by the Deccan plume activity beneath these regions, may be playing a key role in the seismogenesis of such earthquakes. Further, this terrain poses a complex problem in view of geophysical studies (Negi et al. 1983, 1986; Agrawal et al. 2004). Many geophysical methods, including seismics and magnetics, fail to provide desired results while mapping the sub-basalt formations underneath, due to the highly heterogeneous nature of the overburden (Vedanti et al. 2018). The fact is that in spite of several geological and geophysical investigations over this terrain since last four decades (Baumbach et al. 1994; Singh et al. 1999; Mandal et al. 2000; Gupta et al. 2001; Kayal and Mukhopadhyay 2002; Gupta and Gupta 2003; Mandal and Pandey 2011; Kayal 2007, 2008), its seismotectonics and the crustal structures, remain least understood. We do not even know what kind of geologic terrain is concealed below these volcanic sequences.

7.2 Deep Scientific Drilling

203

7.2 Deep Scientific Drilling A deep scientific drilling programme was undertaken in two prominent seismic regions; first, in Killari (Gupta and Dwivedi 1996; Gupta et al. 2003), and currently in Koyna (Gupta et al. 2015). The latter region is characterised by intense Reservoir Triggered Seismicity (RTS) since almost five decades (Gupta 1992; Talwani 1997; Rastogi and Mandal 1999; Rastogi 2001; Agrawal et al. 2004; Yadav et al. 2016; Gupta et al. 2017). Similarly, the other location Killari was devastated in 1993 by one of the deadliest earthquakes of modern times, killing almost 10,000 people. Here, four boreholes were drilled in the surface rupture zone, including 617 m deep KLR-1 borehole, which was drilled 80 m south of surface scarp on the hanging wall near the Killari village (18° 03 07 N, 76° 33 20 E) (Gupta and Dwivedi 1996; Gupta et al. 2003) and it was fully cored. It penetrated 338 m thick column of Deccan basalts, 8 m of infratrappean sediments and a further 270 m of the 2.57 Ga old (Zacharaiah 1998) Neoarchean crystalline basement (Figs. 7.2 and 7.3) belonging to eastern part of the Dharwar craton (EDC). Fig. 7.2 Depths of the basement core samples from the KLR-1 borehole, as discussed in the text. This borehole was drilled in the epicentral zone of 1993 Killari earthquake, Maharashtra India. Location of Gondwana Infratrappeans occurrences are also shown. After Fig. 1.3, Tripathi (2015)

204

7 Seismic, Elastic and Petrophysical Properties of Crustal Rocks: …

Fig. 7.3 Depths of the Deccan basalt samples from the KLR-1 borehole. In this figure, samples above 173 m depth belong to Ambenali Formation and below this depth, to the Poladpur Formation. Sample positions are shown on left side and their corresponding number as well as depths, are shown on the right side. After Fig. 2, Vedanti et al. (2018)

7.3 Measurement of Seismic, Elastic and Petrophysical Properties Tripathi et al. (2012a) made an attempt to measure seismic, elastic and petrophysical properties on 29 cores of crystalline basement rocks from the KLR-1 borehole. Later, Pandey et al. (2016) measured another batch of 14 geologically anomalous samples. Location of these samples are shown in Fig. 7.2. Vedanti et al. (2018) recently studied 35 carefully selected cores of Deccan basalts from the same borehole (Fig. 7.3). In order to comprehend the evolutionary nature of these samples, geochemical and petrological studies were also conducted. Adequate care is taken to insure that the selected cores are unweathered, fresh, dry, crack-free, compact and sufficiently homogeneous and represent typical lithologies of the drilled column.

7.3 Measurement of Seismic, Elastic and Petrophysical Properties

205

7.3.1 Experimental Techniques The measured samples had a fixed diameter of 30 mm, but their lengh varied up to a maximum of about 75 mm. Both the faces of the samples were machined and polished, to ensure that their end-faces were flat and parallel in order to provide a good coupling between the interfaces and the transducers and to measure the bulk volume, using physical dimensions of the sample. All the measurements were carried out at the Rock Mechanics Laboratory of CSIR-National Geophysical Research Institute, Hyderabad, under the ambient room temperature and pressure conditions, following the standard procedures recommended by the International Society for Rock Mechanics (Brown 1981).

7.3.1.1

Density and Porosity

Estimations of density and porosity, requires measurement of the bulk volume, which is done either by fluid displacement (or volumetric) method or alternately, by direct calculation. The former approach involves gravimetric (or Archimedes) method, which is somewhat similar to the triple-weight method of porosity measurement, or alternatively, measurements through pycnometer. This method is normally preferred in case of irregularly shaped samples, but then, it is not suitable for porous (vesicular) samples, which can absorb water. The direct method in comparison, is suitable for regularly shaped samples, like cylindrical cores, wherein the bulk volume can be determined from their physical dimensions. Before measuring density (ρ), the core samples are placed in an electric oven at ~80 °C for about an hour, in order to completely remove the moisture. The diameter and length of the samples, as measured by Vernier Caliper with a precision of 0.02 mm, are used for computing the bulk volume (v). Similarly, the weight (mass, m) of the samples are measured by an electronic balance, which usually have a precision of 0.001 g. The density (ρ) of the specimen can then be calculated by the formula, ρ = m/v. Since the pore volume existing in the core samples, can have an important bearing on petrophysical properties, porosity of the samples are also measured together with density, following the water saturation method. In this method, the mass of the core specimen is measured before and after the saturation of the samples. The difference and the volume of the core specimen, are then used to obtain the porosity. The measured porosity using this technique is considered as the effective porosity, which relates to the connected pores within the measured specimen.

7.3.1.2

P- and S-Wave Velocity

To facilitate the transmission of sound energy between the transducer and the test sample, the piezoelectric transducers are coupled to the specimen on both sides.

206

7 Seismic, Elastic and Petrophysical Properties of Crustal Rocks: …

A very thin film of the machine oil (for P-waves) and honey (for S-waves), are usually used as acoustic couplants between transducers and the core specimen, so that no air remains between the transducers and the face of the polished core. Further, while measuring, the transmitter is pressed to the center of a plane normal to the direction of wave propagation, which generates a stress of about 10 N/cm2 (Brown 1981). Two different sets of transducers are used for measuring Vp and Vs . Both, the transmitting and receiving transducers, are normally kept in a specially made housings, having suitable damping material. A high energy pulser (Olympus Model No. 5058 PR) is then used for generating electrical pulses. The transmitting transducer converts these electrical pulses into stress pulses, that travel through the measured specimen. At the receiving end, the stress pulses are converted into electrical pulses by the receiving transducer. The electric pulse is viewed using a 2-channel digital storage oscilloscope for travel time measurement. In this method, the arrival times of P- and S-waves are measured with a precision of ±1% and ±2% respectively, at 1 MHz frequency. The velocities can then be calculated using the sample length and ultrasonic travel time in the studied specimen. Such measuring approach is quite common worldwide (e.g. Birch 1960; Ivankina et al. 2005; Ji et al. 2009; Kern et al. 2009). However, the main differences in such measurements can usually be the frequencies, that can vary from 1 to 2 MHz. In these type of studies, precision of timing measurements is generally about ±5 ns, and the timing accuracy, better than ±0.5% (Kern et al. 2009). After measuring the velocity and taking into account the earlier measured density, elastic moduli can be calculated using the following relationships: Poisson’s ration (γ) = ½ [(Vp /Vs )2 − 2)]/[(Vp /Vs )2 − 1)] Young’s modulus (E) = ρ (1 + γ) (1 − 2γ) V2p /(1 − γ) Bulk modulus (K) = E/3(1 − 2γ) Shear or rigidity modulus (G) = E/2(1 + γ) where, Vp and Vs are P- and S-wave velocity respectively of the studied rock sample, while, ρ is density of the rock specimen in g/cm3 .

7.3.1.3

P- and S-Wave Attenuation (Qp , Qs )

Seismic attenuation is an important parameter, which defines the loss of energy per cycle. It helps to understand the intrinsic rock properties. A wave whichhas traveled −

fπ

x

some distance x will suffer a loss in its amplitude by a factor given by e QV , where Q is attenuation, f the frequency, and V is the velocity. The attenuation is usually quantified by 1/Q, where Q is the quality factor (total energy/energy lost during one cycle). It can also be written as Q = 2π. Smaller the Q, higher is the attenuation. When the seismic wave propagates through a geological medium, the elastic energy associated with the wave is gradually absorbed by the medium and eventually gets converted to heat, and can be termed as anelastic attenuation. Seismic energy is also lost in scattering. Due to these two mechanisms, intrinsic attenuation and scattering,

7.3 Measurement of Seismic, Elastic and Petrophysical Properties

207

total disappearance of the seismic wave can also happen. In seismic exploration experiments, attenuation of shear wave results from relaxation of the shear modulus, while compressional wave attenuation results from the relaxation of both the shear as well as bulk modulus. For the Deccan volcanic rocks, the nature of seismic wave attenuation has hardly been understood due to lack of laboratory data. One of the prominent investigations over the Deccan basalts include Vedanti et al. (2018), who attempted to measure such properties using the ‘pulse broadening’ technique that involves the use of the pulse ultrasonic double transducer experimental arrangement, as explained in detail by Ramana and Rao (1974). This method enables to determine Q over a wide range of frequencies.

7.3.2 Integrated Geological Studies Petrological and integrated geochemical studies were undertaken on all the chosen core samples from Killari borehole. Whole rock major oxides are measured by Phillips Magix PRO model PW 2440 wavelength dispersive X-ray fluorescence spectrometer, while trace and rare earth elements by microwave acid digestion method and ICP-MS. Modal analysis of the samples have been done using standard point counting method on thin sections.

7.4 Crystalline Basement 7.4.1 Geological Nature of Crystalline Basement Integrated geological investigations on crystalline basement below Killari, indicated the 2.57 Ga old basement to be made up of metasomatised and retrogressed amphibolite to granulite facies transitional rocks. These basement rocks are characterised by about 5–7 kb pressure and corresponds to mid-crustal origin (Pandey et al. 2009, 2014, 2016; Tripathi et al. 2012a, b; Tripathi 2015). Out of the 43 chosen cores from depths between 351.5 and 616.8 m, four samples were of granulites, 22 samples of amphibolite, while the rest seventeen belong to totonalite/granodiorite catagory (Pandey 2016) (Fig. 7.4).

7.4.2 Elastic and Petrophysical Properties Comprehensive results are reproduced in Table 7.1; it reveals that their densities varied widely from 2.67 to 3.11 g/cm3 with a mean of 2.82 ± 0.11 g/cm3 . A large

208

7 Seismic, Elastic and Petrophysical Properties of Crustal Rocks: …

Fig. 7.4 Lithological section pierced by KLR-1 borehole, based on petrophysical, elastic, geological and mineralogical studies (Pandey et al. 2009, 2014; Tripathi et al. 2012a, b; Tripathi 2015). After Fig. 4, Pandey (2016)

variation is also seen in Vp from 5.82 to 6.61 km/s (mean: 6.17 ± 0.21), and in Vs from 3.21 to 4.03 km/s (mean: 3.61 ± 0.16). The other elastic parameters, like Poisson’s ratio, Young’s modulus, Bulk modulus and Shear or rigidity modulus, however, show variations in narrow range. The SiO2 and FeOT variations in these samples, are included in Table 7.2. Theses samples are grouped into three categories of the rocks, basic, intermediate and acid, based on their SiO2 content. Their characteristic variations are summarised in Table 7.3, together with their Si O2 and total FeOT contents. Further, means and details of other elastic parameters are included in Table 7.4. As can be seen from the Table 7.3, basic rocks are characterized by a much higher average density of 2.96 g/cm3 (range: 2.81–3.11 g/cm3 ), which are normally associated with lower crustal rocks. Correspondingly, Vp and Vs show a wider range, averaging around

432

439

440.5

451

459.4

467.6

475

486

493

499.6

505

508.5

520.3

532

Kil-11

Kil-12

Kil-13

Kil-14

Kil-15

Kil-16

Kil-17

Kil-18

Kil-19

Kil-20

Kil-21

Kil-22

Kil-23

Kil-24

351.7

Kil-8

359

589.7

Kil-3

397.3

592.2

Kil-2

Kil-10

568.2

Kil-1

Kil-9

Depth (m)

Sample no.

– –

2.85a

2.80a

2.75

6.15

6.45

2.87a

6.09 6.61

2.71a

2.94

6.09

6.09

2.77

2.9

6.12

6.22

2.84a

2.75

6.21

2.83

6.04

–

2.79a

2.77

6.18

2.75

5.83

6.19

2.97

6.27

2.82a

6.48

2.92

2.87

6.45

–

2.81a

2.95

Vp (km/s)

Density (g/cm3 )

3.49

3.62

3.58

3.58

3.59

–

3.45

3.5

–

3.56

3.31

3.68

–

3.61

–

3.66

3.54

–

–

–

Vs (km/s)

1.89

1.68

1.7

1.7

1.7

–

1.8

1.73

–

1.74

1.76

1.75

–

1.74

–

1.68

1.82

–

–

–

Vp /Vs

0.31

0.23

0.24

0.24

0.24

–

0.277

0.249

–

0.253

0.262

0.258

–

0.253

–

0.226

0.284

–

–

–

Poisson’s ratio

93.79

87.24

87.9

92.03

88.13

–

85.97

84.37

–

87.03

82.17

97.98

–

95.12

–

90.48

95.07

–

–

–

E (GPa)

80.58

53.07

55.28

57.88

55.43

–

64.25

56.02

–

58.72

57.54

67.48

–

63.18

–

54.91

73.36

–

–

–

K (GPa)

(continued)

35.91

35.58

35.59

37.26

35.68

–

33.66

33.78

–

34.73

32.56

38.94

–

37.96

–

36.82

37.02

–

–

–

G (GPa)

Table 7.1 Density, seismic wave velocities and moduli of crystalline basement samples from KLR-1 borehole, drilled in the epicentral zone of the Killari earthquake

7.4 Crystalline Basement 209

Depth (m)

540

546.5

561

565

579.8

596.5

603

613.7

616.8

351.5

367.75

372

380.3

412.2

438

499.5

505.7

507

526.95

540.4

Sample no.

Kil-25

Kil-26

Kil-27

Kil-28

Kil-29

Kil-30

Kil-31

Kil-32

Kil-33

Kil-34

Kil-35

Kil-36

Kil-37

Kil-38

Kil-39

Kil-40

Kil-41

Kil-42

Kil-43

Kil-44

Table 7.1 (continued)

2.72

2.67

2.92

3.07

2.69

3.11

2.75

2.81

2.75

2.84

2.85

2.69

6.11

5.98

5.86

5.86

5.91

6.51

6.23

6.37

6.1

6.09

6.58

6.2

6.18

6.49

2.69

6.19

2.79a

6.09

6.14

5.89

6.1

6.2

Vp (km/s)

2.84

2.72

2.77

3.06

2.76

2.76

Density (g/cm3 )

3.77

3.4

3.6

3.55

3.57

3.77

3.48

4.03

3.59

3.7

3.91

3.63

3.67

3.96

3.64

3.65

3.58

3.21

3.55

3.58

Vs (km/s)

1.62

1.76

1.63

1.65

1.66

1.73

1.79

1.58

1.7

1.64

1.68

1.71

1.68

1.64

1.7

1.67

1.71

1.83

1.72

1.73

Vp /Vs

0.19

0.26

0.2

0.21

0.21

0.25

0.27

0.17

0.24

0.21

0.23

0.24

0.23

0.20

0.24

0.22

0.24

0.29

0.25

0.25

Poisson’s ratio

92.18

78.06

90.43

93.52

83.01

110.19

84.77

106.64

87.58

93.7

107.05

87.73

89.18

105.23

93.11

88.36

88.6

81.63

86.37

88.58

E (GPa)

50.22

54.24

49.8

53.64

48.37

72.97

62.19

53.14

55.05

53.24

65.51

56.24

54.25

59.25

58.56

52.6

56.79

63.87

57.2

58.82

K (GPa)

(continued)

38.6

30.97

37.76

38.66

34.19

44.14

33.3

45.75

35.46

38.83

43.6

35.38

36.37

43.7

37.7

36.21

35.73

31.71

34.69

35.46

G (GPa)

210 7 Seismic, Elastic and Petrophysical Properties of Crustal Rocks: …

546.6

560

585

Kil-45

Kil-46

Kil-47

6.07 6.17 ± 0.21

2.82 ± 0.11

5.82

6.03

Vp (km/s)

2.75

3.08

2.72

Density (g/cm3 )

After Table 6.1, Tripathi (2015) a Indicate density measurement by Archimedes principle

Average (±sd)

Depth (m)

Sample no.

Table 7.1 (continued)

3.61 ± 0.16

3.62

3.66

3.65

Vs (km/s)

1.71 ± 0.07

1.68

1.59

1.65

Vp /Vs

0.24 ± 0.03

0.22

0.17

0.21

Poisson’s ratio

90.99 ± 7.28

88.18

96.54

87.68

E (GPa)

57.97 ± 7.17

53.16

49.25

50.77

K (GPa)

36.86 ± 3.39

36.04

41.14

36.17

G (GPa)

7.4 Crystalline Basement 211

212

7 Seismic, Elastic and Petrophysical Properties of Crustal Rocks: …

Table 7.2 Variation of SiO2 and FeOT content (wt%) in Killari crystalline basement samples Sample

SiO2

FeOT

Sample

SiO2

FeOT

Sample

SiO2

FeOT

Kil-1

60.57

5.36

Kil-20

64.74

6.39

Kil-34

51.85

9.22

Kil-2

56.14

7.18

Kil-21

52.62

18.42

Kil-35

49.96

10.27

Kil-3

57.79

7.16

Kil-22

66.57

5.85

Kil-36

52.32

11.63

Kil-8

53.44

11.30

Kil-23

69.59

3.77

Kil-37

51.49

10.29

Kil-9

52.21

14.72

Kil-24

62.07

7.25

Kil-38

55.91

6.11

Kil-10

58.06

8.50

Kil-25

62.78

6.74

Kil-39

47.78

15.08

Kil-11

54.56

11.67

Kil-26

62.73

7.10

Kil-40

60.45

6.76

Kil-12

55.66

8.63

Kil-27

52.82

13.11

Kil-41

45.12

20.90

Kil-13

55.29

10.55

Kil-28

61.08

8.07

Kil-42

52.47

22.92

Kil-14

52.97

12.13

Kil-29

59.27

8.96

Kil-43

64.56

0.34

Kil-15

63.80

6.22

Kil-30

63.52

6.37

Kil-44

60.85

4.10

Kil-16

64.78

5.90

Kil-31

57.77

8.93

Kil-45

60.57

4.99

Kil-17

64.04

6.50

Kil-32

64.40

6.80

Kil-46

46.10

13.75

Kil-18

57.84

9.86

Kil-33

69.96

3.33

Kil-47

59.96

4.91

Kil-19

62.09

9.50

Source Tripathi (2015) Table 7.3 Mean density and seismic velocity variations in all the three group of basement rocks from KLR-1 borehole, together with their Si O2 and FeOT contents. Sample no.

Density (g/cm3 )

Vp (km/s)

Vs (km/s)

SiO2 (wt%)

FeO (T) (wt%)

Basic rock n

6

6

6

6

6

Mean ± SD (δn−1)

2.96 ± 0.13

6.20 ± 0.30

3.77 ± 0.16

48.72 ± 2.57

13.25 ± 4

Range

(2.81–3.11)

(5.82–6.58)

(3.55–4.03)

(45.12–51.85)

(9.22–20.90)

Intermediate rock n

27

24

21

27

Mean ± SD (δn − 1)

2.83 ± 0.09

6.17 ± 0.20

3.58 ± 0.15

57.42 ± 3.56

Range

(2.69–3.06)

(5.83–6.61)

(3.21–3.96)

(52.21–62.78)

27 9.50 ± 4.13 (4.10–22.92)

Acidic rock n

10

9

9

10

Mean ± SD (δn − 1)

2.74 ± 0.05

6.12 ± 0.07

3.58 ± 0.08

65.60 ± 2.23

Range

(2.67–2.84)

(5.98–6.20)

(3.40–3.67)

(63.52–69.96)

After Table 6.2, Tripathi (2015)

10 5.15 ± 1.95 (0.34–6.80)

7.4 Crystalline Basement

213

Table 7.4 Mean Vp /Vs , Poisson’s ration (γ), Young’s modulus (E), Bulk modulus (K) and Shear or rigidity modulus (G) in three groups of basement rocks from KLR-1 borehole Sample no.

Vp /Vs

Poission’s ratio (γ)

E (GPa)

K (GPa)

G (GPa)

n

6

6

6

6

6

Mean ± SD (δn − 1)

1.64 ± 0.05

0.21 ± 0.03

101.27 ± 6.85

57.96 ± 8.39

42.02 ± 2.68

Range

(1.58–1.73)

(0.17–0.25)

(93.52–110.19)

(49.25–72.97)

(38.66–45.75)

Basic rock

Intermediate rock n

21

21

21

21

21

Mean ± SD (δn − 1)

1.72 ± 0.07

0.24 ± 0.03

89.77 ± 5.47

58.92 ± 7.82

36.15 ± 2.52

Range

(1.62–1.89)

(0.19–0.31)

(81.63–105.23)

(48.37–80.58)

(31.71–43.70)

n

9

9

9

9

9

Mean ± SD (δn-1)

1.71 ± 0.03

0.24 ± 0.01

86.97 ± 3.82

55.76 ± 1.80

35.09 ± 1.77

Range

(1.68–1.76)

(0.23–0.26)

(78.06–93.11)

(53.07-58.72)

(30.97–37.70)

Acidic rock

After Table 6.3, Tripathi (2015)

6.20 (Range: 5.82–6.58 km/s) and 3.77 km/s (Range: 3.55–4.03 km/s) respectively. Geological studies (specially petrography) indicate that such large variations in velocities can be attributed to metasomatic alteration as seen in thin sections of the rocks (Tripathi et al. 2012b; Pandey et al. 2016; Tripathi 2015). This category of rock are highly enriched in iron contents, in which FeOT varies from 9.22 to 20.90 wt%, with an average of 13.25 wt%. Such level of concentration far exceeds to that found even in the lower crustal rocks (8.4 wt%; Rudnick and Fountain 1995). In comparison to the basic rocks, the dominantly occurring intermediate rocks, have a mean density of 2.83 g/cm3 . (Range: 2.69–3.06 g/cm3 ). Correspondingly, average Vp and Vs are found to be 6.17 km/s (Range: 5.83–6.61 km/s) and 3.58 km/s (Range: 3.21– 3.96 km/s) respectively (Table 7.3). Even in these samples, total FeO content is much higher (average: 9.50 wt%), than the average lower crust. As expected, acid rocks show much lower values, wherein mean density, Vp and Vs are 2.74 g/cm3 (range: 2.67–2.84 g/cm3 ), 6.12 km/s (5.98–6.20 km/s) and 3.58 km/s (3.40–3.67 km/s) respectively. They also contain more silica (average: 65.6 wt%) and correspondingly, lesser FeOT (5.15 wt%) than the basic and intermediate category. Estimated density from the measured mean velocity, following Christensen and Mooney (1995) and Barton (1986), comes to about 2.77 to 2.82 g/cm3 , which is in agreement with the lab measured mean density of 2.82 g/cm3 .

214

7.4.2.1

7 Seismic, Elastic and Petrophysical Properties of Crustal Rocks: …

Elastic Moduli

All the elastic moduli parameters like Young’s modulus (E), Bulk modulus (K) and Shear or rigidity modulus (G), fall in a narrow range with respective means 90.99 ± 7.28, 57.97 ± 7.17 and 36.86 ± 3.39 Gpa (Table 7.1); these higher values correspond to basic to intermediate nature of the rocks. These mafic rocks are associated with increase in grain contact area and grain contact stress with negligible porosity. Some samples are characterized by quite higher values of Poisson’s ration (γ) due to inherent high density and high velocity. The means of elastic moduli, as included in Table 7.4, provide the same inference. Metasomatism related chemical alterations and microcracks, can substantially lower the elastic moduli of the rock. The average density of basic (2.96 g/cm3 ) as well as intermediate (2.83 g/cm3 ) samples (Table 7.3), are much higher than acid and intermediate charnockites of south Indian shield and global felsic and para granulites (Table 7.5), but they are reasonably close to the intermediate granulites.

7.4.3 Mineralogical Effects on Density, Velocity and Elastic Moduli The measured velocities and densities in rocks can vary widely, consequent to changes in petrological and chemical composition and as such, the minerals formed by retrograde reactions and metasomatic alterations, can affect these parameters in an extreme manner. In general, there is a tendency of increase in Vp with increasing density in rocks, which are largely unaffected by metasomatic alteration. However, such relationship is not visible in case of Vs and Vp /Vs versus density plots. Similarly, no clear relationship is observed between elastic moduli and density. However, measured densities exhibit a positive correlation with total FeO content (Fig. 7.5a), while a negative association is seen with SiO2 which varies in a large range between Table 7.5 Densities of granulite facies rocks from various terrains Rock type

n

Density (g/cm3 )

Source

Acid charnockites (South Indian shield)

140

2.71 (2.57–2.89)

Subrahmanyam and Verma (1981)

Intermediate charnockites (South Indian shield)

114

2.78 (2.66–3.15)

Subrahmanyam and Verma (1981)

Charnockites (SGT, India)

60

2.73 (2.63–2.77)

Ray et al. (2003)

Felsic granulites (Global)

96

2.76

Christensen and Mooney (1995)

Felsic/para granulites (Global)

42

2.76

Christensen and Mooney (1995)

Felsic granulites

26

2.71 (2.62–2.82)

Rudnick and Fountain (1995)

9

2.86 (2.74–3.02)

Rudnick and Fountain (1995)

Intermediate granulites After Table 5, Tripathi et al. (2012a)

7.4 Crystalline Basement

215

Fig. 7.5 Variation of measured density of amphibolite to granulite facies basement rocks from KLR-1 borehole, against a total FeO, and b SiO2 contents. After Fig. 6.3, Tripathi (2015)

45.12 and 70.0 wt% (Fig. 7.5b). Rudnick and Fountain (1995) indicated a negative relationship between Vp and SiO2 . No such relationship is, however, seen between Vp and FeOT or even SiO2 , possibly due to inherent metasomatic alteration in many of these samples.

216

7 Seismic, Elastic and Petrophysical Properties of Crustal Rocks: …

7.4.4 Mantle Metasomatism and Anomalous Velocity Drop in Iron and Biotite-Rich Rocks Pandey et al. (2016) reported that quite a few Killari basement samples, which are characterised by quite high density, have shown quite low velocities (Table 7.6), together with anomalous geochemical and petrological characters. Most of these samples contained low SiO2 (as low as 45%), but quite high FeOT content up to 23 wt% (Table 7.7). It can be seen from Table 7.6 that the measured Vp is much lower (range: 5.82–6.22 km/s; mean: 6.02 km/s) in these mid crustal rocks. For example, five samples (KIL-14, 27, 41,42 and 46) in this table, with lowest measured velocities ranging between 5.82 and 5.89 km/s (average of 5.85 km/s), have on an average, high density of 3.02 g/cm3 , which are the typical of lower crust. Against this density, the usual range of Vp should have been 6.9–7.0 km/s (Barton 1986; Christensen and Mooney 1995), but the measured average Vp of 5.85 km/s, shows an unusual velocity drop of more than 15%. Such low measured velocities are known to conform with upper crustal granitic velocities. Similarly, values of S-wave velocity and Poisson’s ratio are also much lower with respective means of 3.50 km/s (3.21–3.66 km/s) and 0.24 (0.17–0.29). Universely, a positive relationship exists between density and velocity in crust/mantle rocks (Christenson and mooney 1995; Rudnick and Fountain 1995), but in case of these anomalous basement samples, an inverse relationship is seen (Fig. 7.6a). As mentioned earlier, all these samples are iron-rich. In a few samles, FeOT even exceeded 20%, e.g. KIL-41 and 42. These two iron-rich high density samples, exhibit extremely low velocities of 5.86 km/s (Table 7.6). Petrological studies indicate that they contain big clusters of the opaque minerals (Figs. 7.7), subsequently added after the formation of the crust in the region due to mantle metasomatic activity. Scanning Electron Microscopic studies indicate that most of these opaque minerals are magnetite, apart from a few grains of pyrite and chalcopyrite. The magnetite mineral is usually formd as a result of breakdown of water bearing minerals like biotite and hornblende, during the late stage of magma crystallisation.

7.4.4.1

Mantle Metasomatism

Mantle metasomatism is a prominent geological process, which envolvs mass influx and infiltration of the gaseous-rich hydrothermal fluid from the mantle, that cause widespread chemical alteration, partial replacement and further additions of the minerals under high stress and strain conditions. This process which alters the basic lithological fabric of the rock, often takes place during exhumation and retrogression of the deep seated crustal layers. The Killari basement has been pervasively envolved with such processes. In order to determine the volatile components, Pandey et al. (2009) carried out, a high-temperature differential thermal analyses (DTA) and thermo-gravimetric studies (TG) on some of these rocks, which indicated presence of about 2 wt% of CO2 , besides significant amount of chlorine. These were contributed

560

451

486

493

505

540

561

372

507

KIL-14

KIL-18

KIL-19

KIL-21

KIL-25

KIL-27

KIL-36

KIL-42

2.92

2.75

3.06

2.76

5.86

6.10

5.89

6.20

6.09

6.22

2.90

6.21

2.84a

5.83

6.19

5.82

5.86

Vp (km/s)

2.83

2.97

2.82a

3.08

3.07

Density (g/cm3 )

3.60

3.59

3.21

3.58

3.58

–

3.45

3.31

–

3.66

3.55

Vs (km/s)

E, K and G refer to Young’s, bulk and rigidity (or shear) modulus respectively ‘m’ is mean atomic weight a Density measured by the Archimedes Principle After Table 1, Pandey et al. (2016)

439

KIL-12

Intermediate

505.7

KIL-46

Depth (m)

KIL-41

Basic

Sample no.

1.63

1.70

1.83

1.73

1.70

–

1.80

1.76

–

1.59

1.65

Vp /Vs

0.20

0.24

0.29

0.25

0.24

–

0.28

0.26

–

0.17

0.21

Poisson’s ratio

90.43

87.58

81.63

88.58

92.03

–

85.97

82.17

–

96.54

93.52

E (GPa)

Table 7.6 Depth, density, elastic wave velocities and moduli of 11 anomalous basement samples from KLR-1 borehole

49.80

55.05

63.87

58.82

57.88

–

64.25

57.54

–

49.25

53.64

K (GPa)

37.76

35.46

31.71

35.46

37.26

–

33.66

32.56

–

41.14

38.66

G (GPa)

23.92

22.45

23.02

21. 77

23.51

22.02

22.35

22.62

22.28

22.91

24.05

m

7.4 Crystalline Basement 217

2.64

10.01

20.9

0.22

7.43

4.09

2.32

2.89

0.36

98.29

24.05

TiO2

AI2 O3

FeOT

MnO

MgO

CaO

Na2 O

K2 O

P2 O5

Total

m

22.91

99.55

0.27

2.01

2.34

9.97

10.7

0.23

13.75

11.22

1.43

46.1

Basic KIL-46

‘m’ refers to mean atomic weight After Table 2, Pandey et al. (2016)

45.12

SiO2

Basic KIL-41

22.28

98.39

0.06

1.41

1.91

9.07

6.73

0.21

8.63

13.03

0.72

55.66

Int KIL-12

22.62

98.91

0.09

0.91

2.4

9.37

5.56

0.17

12.13

13.13

0.84

52.97

Int KIL-14

22.35

98.41

0.96

2.05

3.18

4.95

2.29

0.06

9.86

14.98

1.14

57.84

Int KIL-18

22.02

98.73

0.19

2.2

2.98

3.62

1.21

0.04

9.5

15.07

0.77

62.09

Int KIL-19

23.51

98.35

0.26

1.75

2.25

5.08

2.72

0.22

18.42

11.49

1.5

52.62

Int KIL-21

Table 7.7 Whole rock major oxides (wt%), in anomalous basement samples from KLR-1 borehole

21.77

98.49

0.19

2.28

3.25

4.47

1.43

0.07

6.74

15.93

0.61

62.78

Int KIL-25

23.02

98.06

0.33

1.63

1.64

8.27

6.36

0.24

13.11

10.98

1.23

52.82

Int KIL-27

22.45

98.8

0.24

0.82

3.94

7.25

8.11

0.13

11.63

12.2

0.87

52.32

Int KIL-36

23.92

98.67

0.26

1.15

1.93

3.76

5.85

0.16

22.92

6.18

1.44

52.47

Int KIL-42

218 7 Seismic, Elastic and Petrophysical Properties of Crustal Rocks: …

7.4 Crystalline Basement

219

Fig. 7.6 Variation of measured Vp with density (a), and total FeO contents (b), for the anomalous basement cores of KLR-1 borehole as listed in Table 7.6. These samples have shown considerably low velocities against their high densities. After Fig. 7 Pandey et al. (2016).

220

7 Seismic, Elastic and Petrophysical Properties of Crustal Rocks: …

Fig. 7.7 Photomicrographs showing hornblende (Hbl), plagioclase (Plg), titanite (TnT) and opaques (Opq) in sample KIL-41 (a) and elongated deformed quartz (Qtz) and cluster of opaques (Opq) in sample KIL-42 (b). Reaction between opaques, titanite and hornblende is clearly visible in sample KIL-41. After Fig. 6, Pandey et al. (2016)

by the mantle fluids, as deduced from carbon and oxygen isotopic studies (Tripathi et al. 2012b; Pandey et al. 2014). Metasomatic process is also rampant in many other areas of Deccan Traps as well (Karmalkar and Rege 2002; Desai et al. 2004; Sen et al. 2009). After effects of metasomatic processes can be seen in thin sections in the form of biotitization, saussuritization, sericitization, chloritization and uralitization, apart from addition of iron contents. Possibly, no other viable process can enrich the basement with such a high level of FeOT . ‘Biotitisation’ envolves conversion of mafic minerals into biotite, which is a common rock forming and weakly bound iron-rich mineral. It can contribute to significant velocity drop, as interatomic bonds in the crystal structure of the biotites, are known to be weaker than those of other mineral constituents, like pyroxene, plagioclase and hornblende (Alexandrov and Ryzhova 1961). Biotitization also results into lower acoustic impedance, and its higher abundances is known to drag down P-wave velocities to lower levels in middle to lower crustal basic rocks (Kitamura et al. 2001; Ishikawa et al. 2008). The KIL-46 sample, which is basic in nature and has the lowest velocity (5.82 km/s), but high density

7.4 Crystalline Basement

221

(3.08 g/cm3 ), contains 26% biotite (Table 7.8). Mafic rocks, which undergo retrogressive metamorphism, contain large amount of biotites formed at low strain conditions, produced either directly from pyroxene or more commonly from hornblende, as P and T continuously shifts to lower grade due to exhumation of deep seated rocks. Apart from this, there are other velocity reducing metasomaic processes, which too act under retrogressive regime, like saussuritization, sericitization, uralitization and chloritization (Tripathi et al. 2012a; Pandey et al. 2016). Other elastic parameters like, E, K, G and Poisson’ ratios also get effected by such metamorphic transformations.

7.4.4.2

Effect of Microcracks

Elastic wave propagation through the borehole sampled rocks, can also be very sensitive to the insitu microcracks (or microfractures), which often result into the lowering of velocity as well as elastic moduli (Moose and Zoback 1983; Kern et al. 2009; Ullemeyer et al. 2011; Sun et al. 2012). They are produced when the local stresses exceed strength of the rock. Localised deformation, grain boundary stress concentration and temporal temperature changes, can also induce microcracking in the rock matrix (Kranz 1983). Besides, it can also develop en route, while taking out samples from the very deep sections of the boreholes due to ambient pressure change. Majority of stress induced microcracks are extensional in nature. In case of Killari, such microcracks appear to be sealed by chemically saturated fluids, say in the form of calcite. Usually, most of these cracks virtually get closed at the pressure of 2 kb or more and do not reopen completely during subsequent exhumation and retrogression (Kranz 1983; Sun et al. 2012).

7.4.4.3

Role of Mean Atomic Weight

As mentioned before, Vp and density seems to be considerably affected by FeO contents in the rock. Besides observing a negative relationship between density and Vp (Fig. 7.6a), a similar negative relationship is also seen between Vp and total FeO content (Fig. 7.6b). This may be due to the fact that FeO content affect the mean atomic weight considerably, which in turn known to directly effect the intrinsic velocity of the rocks. To know mean atomic weight of any compound, its molecular weight can be divided by the number of particles in the molecular formula. For example, the molecular weight of SiO2 is 60.09, and the number of particles are 3, then the mean atomic weight of the compound will be 20.03. In geological cases, the rock composition is usually expressed in proportions by weight of the oxides. If xi is the proportion by weight and mi the mean atomic weight of oxide i, the mean atomic weight m of the rock is given by (Birch 1961a) as m=



  −1 x i mi

560

KIL-46

451

486

493

505

540

561

372

507

KIL-14

KIL-18

KIL-19

KIL-21

KIL-25

KIL-27

KIL-36

KIL-42

30

42

18

33

40

43

63

45

29

21

34

Plag

20

6

1

28

12

18

6

3

1

1

6

Qtz

26

35

55

18

12

17

2

47

45

45

25

Amph

14

12

21

8

33

19

27

22

26

25

Bt

3

2

10

0.5

2

Kfsp

3

1

Cpx

3

Ep

6

2

1

2

2

2

1

1

1

5

7

Mag

0.5

0.5

0.5

0.5

1

0.5

0.5

0.5

0.5

0.5

Tnt

0.5

0.5

0.5

1

0.5

Py

0.5

0.5

0.5

0.5

0.5

Ccp

0.5

Ap

2.92

2.75

3.06

2.76

2.90

2.84

2.83

2.97

2.82

3.08

3.07

Density (g/cm3 )

After Table 3, Pandey et al. (2016) Amph amphibole; Plag plagioclase; Bt biotite; Qtz quartz; Cpx clinopyroxene; Ep epidote; Kfsp K feldspar; Tnt titanite; Ap apatite; Mag magnetite; Py pyrite; Ccp chalcopyrite

439

KIL-12

Intermediate rock

505.7

Sample depth (m)

KIL-41

Basic rock

Sample no.

Table 7.8 Modal compositions (vol.%) of anomalous basement cores from Killari borehole (KLR-1)

222 7 Seismic, Elastic and Petrophysical Properties of Crustal Rocks: …

7.4 Crystalline Basement Table 7.9 Calculated mean atomic weight of different major oxides, as used in the calculation of mean atomic weight of different rock samples

223 Calculated ‘m’ value SiO2

20.03

TiO2

26.62

Al2 O3

20.39

FeOT

35.92

MnO

35.47

MgO

20.15

CaO

28.04

Na2 O

20.66

K2 O

31.40

P2 O5

40.56

After Table 6, Pandey et al. (2016)

The calculated mean atomic weights for different major oxides, are given in Table 7.9, which are used to estimate the mean atomic weight of diffrent rock samples, as listed in Table 7.6 as ‘m.’ What iron exactly does is that it controls mean atomic weight of the rock (Birch 1961a, b; Christensen 1968). There seems to be one to one relationship between FeOT and the mean atomic weight of the rocks (Fig. 7.8). If the sample is enriched in the iron content, it will have higher mean atomic weight (Birch 1961a, b) and vice versa. All the studied samples are in general characterized by higher FeOT content (Table 7.7). Fig. 7.8 FeOT variation with calculated mean atomic weight in anomalous basement samples (as listed in Table 7.6), showing positive linear relationship. After Fig. 9, Pandey et al. (2016)

224

7 Seismic, Elastic and Petrophysical Properties of Crustal Rocks: …

As per Christensen (1968), for a constant density, higher mean atomic weight will result into lower velocity. In Fig. 7.9a, a plot is made between Vp and density together with the mean atomic weight of the samples, which indicates decrease in Vp with the increase in density of the samples, having higher mean atomic weight. In fact, significant drop in velocity is also well correlated with the increase in mean atomic weight. Presence of opaques like pyrite, which is an iron sulphide mineral itself and also has high density (4.95–5.10 g/cm3 ), will on one side increase the density, but on the other side, will further lower the velocity (Kern et al. 1993). However, there seems to be no such relationship between Vs and density or the iron content.

7.4.4.4

Estimating Velocities and Densities from Mineral Modal Composition

The seismic properties of a particular rock type are mainly dependent on the petrophysical characteristics of the rock matrix and closely related to the elastic properties of the constituent minerals with respect to their volume percentages (Kern et al. 2009). Pandey et al. (2016) combined the mineral modal data (Table 7.8) with the respective velocities and densities of the isotropic monomineralic aggregates of the constituent minerals (Table 7.10) and estimated the average (isotropic) elastic properties of the studied cores, using the formula (Kern et al. 2009): Vr ock =

n 

X i Vi

i=1

where n is the number of minerals in the rock, X i is the volume fraction of each mineral, and V i is the average aggregate velocity of each mineral. In a similar manner, one can estimate the density also. The calculated elastic properties are included in Table 7.11, together with the measured velocities and densities for comparison. The calculated values would, in principle, correspond to the seismic properties of crack-free, non-porous and isotropic aggregate of the basement rocks at ambient temperature-pressure conditions. As can be seen from the Table 7.11, high P-wave velocities are estimated for amphibolites and granulites. In general, abundance of modal hornblende appear to increase the P-wave velocities considerably. In comparison, tonalities are associated with lower velocities. This table, clearly indicates that the calculated velocities from the modal composition of the rocks are markedly higher than the laboratory measured velocities, specially in case of Vp , where the velocity increase is almost 8% (from an average of 6.02 to 6.51 km/s). However in case of Vs , such increase is negligible (only about 1.4%). Further, when a plot is made between calculated Vp and density, based on mineral modes and single crystal properties (Fig. 7.9b), a normal positive trend emerges. The comparison of Fig. 7.9a and b, convincingly demonstrates the marked effects

7.4 Crystalline Basement

225

Fig. 7.9 Plot of measured Vp versus density together with their mean atomic weight (a), and calculated Vp versus density (Table 7.11) from the modal data for the same samples (b). Significant decrease in Vp with increasing density and mean atomic weight (a) can be ascribed to mantle metasomatism. After Fig. 10, Pandey et al. (2016).

226

7 Seismic, Elastic and Petrophysical Properties of Crustal Rocks: …

Table 7.10 Velocities and densities of the isotropic rock forming minerals (Kern et al. 2009) used for the calculation of elastic properties Minerals

Vp (km/s)

ρ (g/cm3 )

Vs (km/s)

Quartz

6.05

4.09

2.648

Plagioclase

6.30

3.45

2.662

K’feldspar

5.93

3.26

2.555

Biotite

6.01

3.00

3.05

Amphibole

7.20

3.77

3.074

Chlorite

6.01

3.00

3.00

Fe-sulfides

7.92

5.06

5.01

Clinopyroxene

7.70

4.38

3.31

Fe-oxides

6.95

4.20

5.201

Table 7.11 Calculated elastic properties of anomalous basement samples from the modal data. Laboratory measured values are also included for comparison Sample no.

Measured density (g/cm3 )

Measured Vp (km/s)

Measured Vs (km/s)

Calculated density (g/cm3 )

Calculated Vp (km/s)

Calculated Vs (km/s)

KIL-41

3.07

5.86

3.55

3.05

6.48

3.51

KIL-46

3.08

5.82

3.66

3.11

6.68

3.54

Basic rock

Intermediate Rock KIL-12

2.82

6.19

–

2.98

6.67

3.53

KIL-14

2.97

5.83

3.31

2.91

6.77

3.66

KIL-18

2.83

6.21

3.45

2.80

6.24

3.38

KIL-19

2.84

6.22

–

2.87

6.37

3.56

KIL-21

2.90

6.09

3.58

2.89

6.30

3.43

KIL-25

2.76

6.20

3.58

2.82

6.35

3.65

KIL-27

3.06

5.89

3.21

3.02

6.74

3.54

KIL-36

2.75

6.10

3.59

2.91

6.58

3.57

KIL-42

2.92

5.86

3.60

2.98

6.48

3.65

After Table 5, Pandey et al. (2016) Note In calculation of Vp , Vs and density, the values of amphibole, chlorite and plagioclase (Table 7.10) have been respectively used for apatite, epidote and titanite in Table 7.8

of metasomatism over the physical properties of the rocks. Quite likely regional metasomatism may be one of the main cause of crustal low velocities zones (not always fluids), which are often referred to be associated with earthquake nucleation.

7.5 Deccan Volcanics

227

7.5 Deccan Volcanics Western part of India exposes thick sequence of Deccan volcanic rocks. These are characterized by number of volcanic formations (Ramakrishnan and Vaidyanadhan 2008), which are made up of as many as 46 individual basaltic flows (Mishra et al. 2017) whose cumulative thickness exceeds 3 km (Table 7.12). These volcanics, mainly tholeiitic in composition, are highly differentiated and partially eroded at the surface, and possibly erupted in a quick succession of less than one million year (Courtillot et al. 1988; Renne et al. 2015) across Cretaceous-Paleocene (K/Pb) boundary (Gradstein et al. 2012). However, there may be some lava sequences (like Mandla lavas), which may be relatively younger (64.21 ± 0.33 Ma) than the main Deccan volcanic event (Shrivastava et al. 2014, 2015, 2017). Palaeomagnetic data fix their eruption period mainly during chron 29r (Schoene et al. 2015). Recent studies, based on palaeomagnetic data and other geological evidences on cores from Koyna deep borehole, too indicate their eruption in a single geomagnetic polarity transition period (Radhakrishna et al. 2019).

7.5.1 Geological Nature of Basaltic Column Two major formations, Ambenali and Poladpur, were penetrated by the KLR-1 borehole, which contained eight flows of aa-type (blocky lavas), having an amygdaloidal and/or vuggy top, a massive interior and chilled glassy basal part (Gupta et al. 2003). Both of the formations belong to the Wai Subgroup in Deccan volcanic chronology (Mitchell and Widdowson 1991). Out of the eight flows, the first four flows lying above the red bole layers (encountered between 173 and 178 m depth), belong to Ambenali Formation, while the rest four flows occurring below it, belong to Poladpur Formation (Fig. 7.3). Recently, Vedanti et al. (2018) investigated in detail ultrasonic Table 7.12 Approximate age and thicknesses of different Deccan volcanic formations, based on studies of Schoene et al. (2015)

Formation

Approximate thickness (m)

Chron

Age (Ma)

Mahabaleshwar

280

C29n

65.5–65.7

Ambenali

550

C29r

65.5–65.7

Poladpur

375

C29r

66.0–66.3

Bushe

325

C29r

66.0–66.3

Khandala

140

C29r

66.0–66.3

Bhimash

150

C29r

66.0–66.3

Thakurvadi

650

C29r

66.0–66.3

Neral

100

C29r

66.0–66.3

Igatpuri

600

C29r

66.0–66.3

Jawhar

200

C29r

66.0–66.3

228

7 Seismic, Elastic and Petrophysical Properties of Crustal Rocks: …

P- and S-wave attenuations and other seismic and petrophysical properties on 35 selected basalt cores. These samples are compact and sufficiently homogeneous and came from the depths between 20.5 and 332.5 m (Fig. 7.3). Thus these samples cover the entire column of the Ambenali and Poladpur volcanic formations, stacked over the Killari basement. The basalts penetrated by the borehole, are mostly fine to medium grained, rarely coarse-grained, and highly massive to vesicular in nature. Based on density measurements on samples from boreholes KLR-1, 2, 3, Reddy et al. (1998) earlier reported that 53% of volcanic sequence is made up of massive variety (meaning high density and negligible porosity), while the rest 47%, belong to vesicular and amygdaloidal (non-massive) type. Out of these two varieties, chosen massive basalt cores are heavy and greenish black to dark black in color with metallic lustre, while the vesicular ones, found at the top of the flows, are greyish brown to dark brown in color (Vedanti et al. 2018). In general, these basalts are relatively Fe and Mg-rich and silica deficient in composition. Petrologically, these basalts contain plagioclase, pyroxene phenocrysts, and microphenocrysts and occasionally olivine as major constituents and magnetite and secondary silicates, as accessory minerals. Quite a few samples are extremely glassy in nature, while many others contained abundant microlites and plagioclase laths. They also contained secondary minerals, like chlorophaeite and other forms of silicates, formed due to alteration of pyroxene and plagioclase grains. Mineralogical studies indicate that these basalts also contained rare minerals like moganite, ferrous saponite and zonation in zeolites (Parthasarathy 2006; Parthasarathy et al. 2001, 2003, 2007). Geochemically, SiO2 content varied in a narrow range, from 45.74 to 47.36 wt% with a mean of 46.64 wt% in the Ambenali Formation (first half of the borehole), while, in Poladpur Formation (bottom half of the sequence), it was almost 2.0 wt% higher at 48.51 wt% (range: 47.06–50.35 wt%). This indicates that the Poladpur formation basalts are a little more silicic than the Ambenali basalts. These basalts are also rich in MgO contents, which suggest that the original source rock was rich in pyroxene and olivine. This inference is supported by petrological studies. The MgO contents average around 8.11 wt% (range: 7.49–9.42 wt%) for Ambenali Formation, and with slightly lower average of about 7.19 wt% (range: 5.39–9.14 wt%) in Poladpur Formation. Apart from generally higher MgO, average FeOT in Ambenali and Poladpur Formations are also on higher side around 13.57 wt% (range: 12.60– 14.73 wt%) and 13.25 wt% (range: 12.66–14.24 wt%), respectively, than the usually occurring Deccan tholeiitic basalts. Intriguingly, these basalts are highly rich in TiO2 contents (2.18–3.74 wt%) as well, than the known tholeiitic basalts of the Deccan traps.

7.5 Deccan Volcanics

229

7.5.2 Seismic and Petrophysical Properties Petrophysically, this volcanic terrain is not too well studied. In fact, seismic attenuation study was hardly undertaken, except by Weiner et al. (1987), who studied ultrasonic P- and S-wave attenuation characteristics and estimated Qp for two tholeiitic basalt samples from Deccan Traps region, using the torsional pendulum technique. Their measurements carried out at very low frequencies (1.3 Hz), produced Qs (shear wave quality factor) ranging from 25 to >1000, which was similar to that determined by Volarovitch and Gurvitch (1957) for basalt samples from USSR. Weiner et al. (1987) concluded that viscous sliding at grain boundaries causes relaxation and anelastic energy loss in the core samples. As mentioned earlier, study by Vedanti et al. (2018) included density, porosity, P- and S-wave velocities together with seismic wave attenuation. Such information are considered crucial for seismic modeling of the sub-basalt formations, as well as underlying deep structural features. All the measurements were made both in dry, as well as in saturated states, and at the ambient room temperature/pressure. Results of the measurements along with estimated elastic modulus, are summarised in Table 7.13 and their Formation-wise averages, along with characteristic SiO2 and FeOT contents, are included in Table 7.14.

7.5.2.1

Density and Porosity

The dry and saturated densities of 13 massive basalt cores from the Ambenali Formation varied in a limited range from 2.81 to 3.00 g/cm3 (mean: 2.91 g/cm3 ) and 2.83 to 3.01 g/cm3 (mean: 2.92 g/cm3 ) respectively. The values are almost identical due to their much lower porosity (average: 1.2%). The dry and saturated density of the sole vesicular basalt sample from this formation is 2.24 and 2.40 g/cm3 respectively, which is related to its higher porosity of 16.6%. Measured dry and saturated densities of the 14 massive basalt cores from the Poladpur Formation, varied from 2.80 to 2.94 g/cm3 (mean: 2.88 g/cm3 ) and 2.80– 2.95 g/cm3 (mean: 2.90 g/cm3 ) respectively. They are almost similar due to their intrinsically negligible porosity, averaging 1.08%. The dry and saturated densities in vesicular basalts of this formation, on the other hand, varied from 2.28 to 2.71 g/cm3 (mean: 2.54 g/cm3 ) and 2.43 to 2.79 g/cm3 (means: 2.65 g/cm3 ) respectively.

7.5.2.2

P- and S-Wave Velocity

Average P-wave velocities in dry and saturated samples of massive Ambenali basalts are estimated to be 5.78 (range: 4.91–6.48) and 5.93 (range: 5.03–6.50) km/s respectively, with corresponding Vs of 3.36 (range: 3.02–3.68) and 3.45 (3.10–3.64) km/s. However, in case of vesicular sample, dry and saturated Vp for this formation are much low at 3.76 and 3.67 km/s respectively, with corresponding Vs of 2.57 and 2.47 km/s.

Depth (m)

30.5

39.6

67.8

70.8

124.4

134.6

145

148

152

164.4

168

171.5

96

KIL-51

KIL-67

KIL-68

KIL-52

KIL-69

KIL-70

KIL-71

KIL-72

KIL-55

KIL-73

KIL-74

KIL-56

KIL-53

188.5

209.6

213.8

216.3

KIL-57

KIL-58

KIL76

KIL-77

Poladpur Formation

20.4

KIL-66

Ambenali Formation

Sample

M

M

M

M

V

M

M

M

M

M

M

M

M

M

M

M

M

M

2.94

2.92

2.90

2.88

2.24

2.86

2.87

2.81

2.91

2.91

2.90

2.95

2.94

2.96

3.00

2.91

2.87

2.89

Dry

2.95

2.94

2.92

2.92

2.40

2.87

2.88

2.83

2.93

2.91

2.90

2.96

2.95

2.97

3.01

2.92

2.88

2.91

Sat

ρ (g/cm3 )

2.97

2.97

2.97

3.00

2.68

2.90

2.91

2.87

2.97

2.92

2.92

2.96

2.96

3.00

3.02

2.94

2.90

2.96

Part

1.3

1.7

2.4

3.9

16.6

1.6

1.5

1.8

1.8

0.5

0.7

0.3

0.8

1.4

0.6

1.0

1.0

2.4

φ (%)

5.70

5.81

5.48

4.76

3.76

5.11

4.99

4.91

5.99

5.96

5.81

6.27

6.26

6.03

6.48

5.90

5.84

5.63

Dry

5.69

5.83

5.65

5.01

3.67

5.34

5.49

5.03

6.07

6.14

6.14

6.32

6.31

6.04

6.50

5.94

5.89

5.84

Sat

Vp (km/s)

3.35

3.41

3.48

2.88

2.57

3.07

3.08

3.02

3.49

3.53

3.12

3.64

3.63

3.51

3.68

3.41

3.38

3.14

Dry

3.48

3.54

3.47

2.84

2.47

3.15

3.29

3.10

3.44

3.64

3.43

3.66

3.69

3.50

3.64

3.42

3.39

3.44

Sat

V s (km/s)

0.24

0.24

0.16

0.21

0.06

0.22

0.19

0.20

0.24

0.23

0.30

0.25

0.25

0.24

0.26

0.25

0.25

0.28

Dry

0.20

0.21

0.20

0.26

0.09

0.23

0.22

0.20

0.26

0.23

0.27

0.25

0.24

0.25

0.27

0.25

0.25

0.24

Sat

Poisson’s ratio

Table 7.13 Petrophysical, elastic and attenuation properties of Deccan basalt cores from KLR-1 borehole

81

84

82

58

31

66

65

61

88

89

73

98

97

91

102

84

82

72

Dry

86

89

84

59

30

70

76

65

87

95

87

99

99

91

101

85

83

84

Sat

E (GPa)

116

464

779

93

7

95

51

42

–

61

177

33

1196

136

330

630

713

1015

Dry

Qp

136

1960

715

88

6

91

60

44

–

696

286

102

–

228

–

1002

1304

881

Sat

78

49

287

35

6

42

43

49

–

66

170

53

48

82

–

318

124

61

Sat

(continued)

57

61

306

65

6

67

49

38

158

38

110

31

108

59

177

503

148

410

Dry

Qs

230 7 Seismic, Elastic and Petrophysical Properties of Crustal Rocks: …

258

265

269.7

271.85

279.6

287

317

327

327.1

332.45

196

220

226.6

231

241

243

250.2

KIL-85

KIL-62

KIL-82

KIL-83

KIL-63

KIL-84

KIL-87

KIL-65

KIL-88

KIL-89

KIL-75

KIL-59

KIL-78

KIL-79

KIL-60

KIL-61

KIL-80

V

V

V

V

V

V

V

M

M

M

M

M

M

M

M

M

M

2.30

2.47

2.28

2.73

2.71

2.68

2.65

2.91

2.92

2.89

2.82

2.91

2.92

2.87

2.80

2.80

2.90

Dry

2.46

2.60

2.43

2.79

2.78

2.75

2.74

2.91

2.93

2.90

2.84

2.91

2.93

2.88

2.82

2.80

2.90

Sat

ρ (g/cm3 ) Part

2.75

2.83

2.68

2.91

2.91

2.89

2.92

2.92

2.94

2.91

2.87

2.92

2.94

2.89

2.84

2.81

2.91

16.6

12.7

15.0

6.3

6.8

7.3

9.4

0.5

0.6

0.6

0.5

0.3

0.4

0.6

1.4

0.6

0.3

φ (%)

3.30

3.91

3.14

4.63

4.39

4.00

3.64

5.93

5.88

5.97

5.17

6.23

6.15

5.86

5.48

5.71

6.18

Dry

3.28

3.97

3.41

4.78

4.64

4.36

3.91

5.95

5.93

6.03

6.03

6.23

6.15

5.89

5.60

5.72

6.23

Sat

Vp (km/s)

1.99

2.60

2.17

2.88

2.56

2.62

2.19

3.36

3.47

3.45

3.10

3.59

3.48

3.31

3.37

3.33

3.51

Dry

1.70

2.44

2.19

2.95

2.70

2.46

2.04

3.46

3.49

3.49

3.43

3.59

3.48

3.36

3.34

3.34

3.51

Sat

V s (km/s)

0.22

0.11

0.04

0.18

0.24

0.12

0.22

0.26

0.23

0.25

0.22

0.25

0.26

0.27

0.20

0.24

0.26

Dry

0.32

0.20

0.15

0.19

0.24

0.27

0.31

0.24

0.24

0.25

0.26

0.25

0.26

0.26

0.22

0.24

0.27

Sat

Poisson’s ratio

M and V refer to massive and vesicular varieties of Deccan basalts and Part refers to particle density After Table 1, Vedanti et al. (2018)

Depth (m)

Sample

Table 7.13 (continued)

22

37

22

54

44

41

31

83

87

86

66

94

90

79

76

77

90

Dry

17

35

25

56

49

41

29

87

88

86

86

94

90

82

77

78

91

Sat

E (GPa)

7

10

9

42

27

12

13

2855

288

225

56

–

339

53

23

64

–

Dry

Qp Sat

6

19

11

44

46

17

30

–

270

437

65

–

498

44

33

61

–

7

10

5

29

33

15

11

1132

50

75

40

47

129

21

24

59

949

Dry

Qs Sat

5

5

6

49

49

13

12

–

79

47

41

83

130

26

47

55

506

7.5 Deccan Volcanics 231

ρ (g/cm3 )

φ (%)

2.24 1

2.86 14

Vesicular (n)

Avg. (n)

2.30 14

16.55 1

1.20 13

3.47 35

φ (%)

2.77 21

2.81 35

ρ (g/cm3 )

Avg. n

Total Avg. n

Sample

2.92 13

2.40 1

Massive n

Vesicular n

16.55 1

1.20 13

Ambenali Formation (saturated samples)

4.25 21

10.58 7

2.54 7

Vesicular n

1.08 14

2.88 14

Massive n

Poladpur Formation (dry samples)

2.91 13

Massive (n)

Ambenali Formation (dry samples)

Sample

3.67 1

5.93 13

V p (km/s)

5.32 35

5.11 21

3.86 7

5.74 14

5.64 14

3.76 1

5.78 13

V p (km/s)

2.47 1

3.45 13

vs (km/s)

3.15 35

3.05 21

2.43 7

3.36 14

3.30 14

2.57 1

3.36 13

V s (km/s)

13.71 1

13.56 13

FeOT (wt%)

13.38 35

13.25 21

13.28 7

13.24 14

13.57 14

13.71 1

13.56 13

FeOT (wt%)

46.62 1

46.64 13

SiO2 (wt%)

47.77 35

48.51 21

47.95 7

48.80 14

46.64 14

46.62 1

46.64 13

SiO2 (wt%)

0.09 1

0.24 13

γ

0.22 35

0.21 21

0.16 7

0.24 14

0.23 14

0.06 1

0.24 13

γ

30 1

86 13

E (GPa)

71 35

66 21

36 7

81 14

79 14

31 1

82 13

E (GPa)

6 1

469 10

Qp

311 32

288 19

17 7

446 12

345 13

7 1

373 12

Qp

(continued)

6 1

96 11

Qs

144 35

149 21

16 7

215 14

136 14

6 1

146 13

Qs

Table 7.14 Average density, porosity, elastic moduli, P- and S-wave velocities and attenuation, together with FeOT and SiO2 contents in samples of Ambenali and Poladpur formations in KLR-1 borehole

232 7 Seismic, Elastic and Petrophysical Properties of Crustal Rocks: …

2.88 14

Avg. n

2.30 14

φ (%)

2.65 7

2.81 21

2.84 35

Vesicular n

Avg. n

Total Avg. n

3.47 35

4.25 21

10.58 7

1.08 14

After Table 2, Vedanti et al. (2018)

2.90 14

Massive n

Poladpur Formation (saturated samples)

ρ (g/cm3 )

Sample

Table 7.14 (continued)

5.46 35

5.25 21

4.05 7

5.85 14

5.77 14

V p (km/s)

3.19 35

3.06 21

2.36 7

3.42 14

3.38 14

vs (km/s)

13.38 35

13.25 21

13.28 7

13.24 14

13.57 14

FeOT (wt%)

47.77 35

48.51 21

47.95 7

48.80 14

46.64 14

SiO2 (wt%)

0.24 35

0.24 21

0.24 7

0.24 14

0.23 14

γ

74 35

68 21

36 7

84 14

82 14

E (GPa)

317 29

249 18

25 7

392 11

427 11

Qp

83 32

80 20

20 7

112 13

89 12

Qs

7.5 Deccan Volcanics 233

234

7 Seismic, Elastic and Petrophysical Properties of Crustal Rocks: …

Average dry and saturated Vp in massive basalts of the Poladpur Formation, are a little lower at 5.74 (range: 4.76–6.23) and 5.85 (range: 5.01–6.23) km/s respectively, with corresponding Vs of 3.36 (range: 2.88–3.59) and 3.42 (range: 2.84–3.59) km/s. In case of vesicular basalts, averaged dry and saturated Vp are found to be 3.86 (range: 3.14–4.63) and 4.05 (range: 3.28–4.78) km/s respectively, with correspondingly low Vs of 2.43 (range: 1.99–2.88) and 2.36 (range: 1.70–2.95) km/s.

7.5.2.3

Elastic Moduli

Estimated Young’s modulus in massive basalt of both the formations (Ambenali as well as Poladpur) varied from 58 to 102 GPa in case of dry samples and 59– 101 GPa in saturated samples. These values are reasonably high and can be ascribed to compactness and resultant increase in grain contact area and grain contact stress in highly mafic massive core samples with negligible porosity. Corresponding values of Young’s modulus are quite low in case of vesicular basalts; 22–54 GPa in dry samples, and 17–56 GPa in case of saturated samples, indicating dominance of porosity in such samples. Similarly, Poisson’s ratio in massive basalts ranged from 0.16 to 0.30 in the dry state and 0.20–0.27 in saturated state. In case of vesicular basalts, Poisson’s ratio are much lower; 0.04–0.24 in dry state and 0.09–0.31 in saturated state (Table 7.13). Effect of chemical alterations, microcracks, and vesicles filled by secondary minerals, are seen in lowering the elastic moduli. As can be seen from the Table 7.14, a distinction between Ambenalli and Poladpur Formations is quite evident. The massive cores of Ambenali Formation are characterised by a little higher average saturated density (2.92 g/cm3 ) and Vp (5.93 km/s) compared to 2.90 g/cm3 and 5.85 km/s respectively in Poladpur Formation (Table 7.14). This can be ascribed partly to the non-contaminated nature (Cox and Hawkesworth 1985; Beane et al. 1986; Mahoney 1988; Cox and Mitchell 1988; Sano et al. 2001), and almost 2 wt% lesser SiO2 content (46.64 wt%) as well as higher MgO contents (8.11 wt%) of the Ambenali Formation, which makes it more mafic than the Poladpur Formation that has relatively lower MgO (6.8 wt%) and higher SiO2 content (48.80 wt%). Petrophysical properties suggest relatively slow rate of crystallization for the Ambenali Formation, which is also one of the thickest formations in Deccan chronology. Contaminated nature of the Poladpur Formation is consistent with their REE abundances. Usually, Ba elemental concentrations and Ba/Nb ratios are found to be high in crustally contaminated rocks. In massive samples of Poladpur Formation, Ba concentrations and Ba/Nb ratios are quite high at 199.3 ppm and 16.3 respectively, compared to 79.5 ppm and 9.42 in Ambenali Formations. Consequently, Poladpur Formation is silica-rich by almost 2 wt% than the Ambenali Formation (Table 7.14). It is generally known that Vp has an inverse relationship with SiO2 content (e.g., Rudnick and Fountain 1995), while MgO content which is generally related to olivine, has a positive impact on velocities.

7.5 Deccan Volcanics

235

Further, petrophysically, the VIth flow from the top (i.e. 218.8–254.5 m) belonging to Poladpur Formation, appears to be quite anomalous. It is characterised by much low density, Vp , Vs and E, apart from higher seismic attenuation. A petrographic study carried out on a thin section of one of the samples (KIL-79) of this flow, confirms its highly altered nature, containing a large amount of secondary silicate minerals as well as chlorophaeite (Fig. 7.10b). The considerable lowering in petrophysical attributes of this particular flow, may be closely related to the higher porosity, due to dominance of vesicular nature. Petrophysical and geochemical nature of this flow, suggest that the magma was volatile-rich, got deposited in quick succession and consequently, volatiles could not fully escape. This flow subsequently underwent hydrothermal alteration and recrystallization, both of which have influenced the elastic parameters. Considering its contaminated nature, it must have erupted through secondary magma chambers at upper crustal depths. Figure 7.11 would supFig. 7.10 Photomicrographs of Deccan basalts from thin sections, showing glassy nature in sample KIL-68 (a), and highly altered plagioclase grain (Plg) and chlorophaeite (Chp) in KIL 79 (b). Modified after Fig. 3, Vedanti et al. (2018)

236

7 Seismic, Elastic and Petrophysical Properties of Crustal Rocks: …

Fig. 7.11 S-wave velocity variation with depth beneath Deccan Trap Volcanic (DTV) covered terrain of the Eastern Dharwar Craton (EDC). For comparison, average S-wave velocity distribution (converted from P-wave velocity; Christensen and Mooney 1995) beneath global shields and platforms is also included. Moho depth is shown by dashed line. Modified after Fig. 2, Pandey 2008

port it (Pandey 2008). This figure indicates that 4–5 km thick high-velocity primitive Deccan magma (peridotitic/eclogitic; Sen and Chandrasekharam 2011) underplated across the Moho. Besides, from 5 km down to 20 km depth, crustally contaminated unerupted Deccan magma is still present, as indicated by much higher velocities compared to global shields and platforms.

7.5.2.4

Relationships: Elastic and Petrophysical Parameters

Elastic and petrophysical properties are greatly affected by mineral composition, their geological evolution and subsequent alteration, and also pore fluid pressure (Kern et al. 1993, 2009; Tripathi et al. 2012a; Pandey et al. 2016). A linear relationship between density and porosity, may indicate a direct relationship between density of solid phase and the measured density. Since the basalts mainly constitutes pyroxene and plagioclase, it is susceptible to alteration due to hydrothermal fluids through the micro fractures. For example, plagioclase has two sets of cleavage (or weak zones) that translates into weak binding. Secondary hydrous minerals formed by chemical alteration, like chlorophaeite (Fig. 7.10b), clays, sericite, chlorite, epidote, etc. effectively seal such cracks. Their presence invariably lowers the particle (or grain) density as well as moduli of the rock and thus correlates with increasing porosity (Christensen et al. 1980; Carlson and Herrick 1990). The seismic velocities Vp and Vs too decrease linearly with increasing porosity, but only till the porosity does not exceed 10%, beyond that velocity drop is more gradual (Vedanti et al. 2018). This means that beyond this limit, the presence of secondary minerals in the rock matrix dominates the velocities, rather than saturation. Further, almost one to one relationship can be seen between Vp and Vs for dry (Fig. 7.12a) as well as saturated states (Fig. 7.12b). It appears that the porosity is a

7.5 Deccan Volcanics

237

Fig. 7.12 Plots between Vp and Vs for dry (a) and saturated states (b) in Deccan basalt samples of KLR-1 borehole. Solid and open blue circles refer respectively to massive and non-massive cores of Ambenali Formation, while similar red circles to Poladpur Formation. After Fig. 12, Vedanti et al. (2018)

major microstructural controlling parameter over both density and velocity. Addition of the alteration products to the rock matrix, can significantly lower the grain density and apparent grain bulk, and shear moduli of the rock and consequently, P-wave and S-wave velocities (Cerny and Carlson 1999).

7.5.3 Compressional and Shear Wave Attenuation Measured attenuation values and their ratios for both the formations studied by Vedanti et al. (2018) are listed in Tables 7.13, 7.14 and 7.15. As can be seen from Table 7.13, the measured Qp exhibits a wide range of values from 33 to 1960 in saturated massive basalts and 6–46 for vesicular basalts. The saturated Qs , on the other hand, varies in a range, 26–506 in massive basalts and 5–49 in vesicular samples.

238

7 Seismic, Elastic and Petrophysical Properties of Crustal Rocks: …

Table 7.15 Vp /Vs and Qs /Qp ratios in Deccan basalt cores from KLR-1 borehole S. No

Sample name

Depth

Nature of basalt

Vp /Vs dry

Vp/Vs sat

QS /QP dry

Qs /Qp sat

1.

KIL-66

20.4

Massive

1.80

1.70

0.40

0.07

2.

KIL-51

30.5

Massive

1.73

1.74

0.21

0.10

3.

KIL-67

39.6

Massive

1.73

1.74

0.80

0.32

4.

KIL-52

70.8

Massive

1.72

1.72

0.43

0.36

5.

KIL-70

134.6

Massive

1.72

1.73

0.93

0.52

6.

KIL-71

145

Massive

1.86

1.79

0.62

0.59

7.

KIL-72

148

Massive

1.69

1.68

0.62

0.10

8.

KIL-74

168

Massive

1.62

1.67

0.96

0.73

9.

KIL-56

171.5

Massive

1.67

1.69

0.70

0.47

10.

KIL-57

188.45

Massive

1.65

1.76

0.70

0.40

11.

KIL-58

209.55

Massive

1.57

1.63

0.39

0.40

12.

KIL76

213.8

Massive

1.70

1.64

0.13

0.02

13.

KIL-77

216.3

Massive

1.70

1.63

0.49

0.57

14.

KIL-62

265

Massive

1.71

1.71

0.92

0.90

15.

KIL-82

269.7

Massive

1.62

1.68

1.07

1.43

16.

KIL-83

271.85

Massive

1.77

1.75

0.39

0.58

17.

KIL-63

279.6

Massive

1.76

1.76

0.38

0.26

18.

KIL-87

317

Massive

1.67

1.76

0.72

0.63

19.

KIL-65

327

Massive

1.73

1.73

0.33

0.11

20.

KIL-88

327.1

Massive

1.70

1.70

0.17

0.29

21.

KIL-53

96

vesicular

1.46

1.49

0.82

0.87

22.

KIL-75

196

vesicular

1.66

1.91

0.83

0.38

23.

KIL-59

220

vesicular

1.53

1.77

1.22

0.78

24.

KIL-78

226.6

vesicular

1.72

1.72

1.21

1.06

25.

KIL-79

231

vesicular

1.61

1.62

0.69

1.10

26.

KIL-60

241

vesicular

1.45

1.56

0.57

0.52

27.

KIL-61

243

vesicular

1.51

1.63

1.03

0.27

28.

KIL-80

250.2

vesicular

1.66

1.93

0.98

0.84

After Table 3, Vedanti et al. (2018)

In case of dry samples, Qp ranges between 23 and 2855 for massive and from 7 to 42 for vesicular samples. Correspondingly, dry Qs varies from 21 to 1132 in massive and 5 to 33 in vesicular basalts.

7.5 Deccan Volcanics

7.5.3.1

239

Seismic Attenuation and Petrophysical Properties

In general, high seismic attenuation (Qp = 6–46, Qs = 5–49) is recorded in highly porous saturated vesicular cores. Some of the massive samples, however, exhibited high attenuation too (Table 7.13). These non-porous samples are often glassy in nature (Fig. 7.10a), sometime highly altered containing secondary minerals including chlorophaeite, and other forms of secondary silicates (Fig. 7.10b). which are hydrous in nature. Some of these samples have also shown conchoidal fractures. These factors appear responsible for seismic energy loss in such cases. In comparison to the Deccan basalts, oceanic basalts from Pacific (ODP holes 594B and 896A) and Atlantic (holes 395A) oceans, have shown Qp variation from 5 to 35, and Qs from 10 to 100 (Goldberg and Sun 1997). The Qp and Qs values for oceanic basalt, conform to the values for vesicular Deccan basalt, but not to the values for massive basalt. This would mean that the oceanic basalts are not massive, due to their possible formation by rapid cooling. Lewis and Jung (1989), Wepfer and Christensen (1990, 1991) and White and Clowes (1994) did indicate a close relation between low velocity, high crack/fracture porosity, a high degree of alteration and low Qp (i.e., high attenuation). Table 7.13 shows that seismic wave attenuation tend to increase with increase in porosity, but decreases with increase in density, as well as Vp and Vs . However, the seismic attenuation in high density and low porosity massive basalts may not be much sensitive to variation in density and porosity (Table 7.13), but possibly reflects dependence over composition of the rock matrix, which can play a major role in seismic energy loss.

7.5.4 Characteristic Density and P- and S-Wave Velocity for Deccan Basalts The Ambenali and Poladpur Formations are one of the thickest formations in Deccan volcanic stratigraphy. Their cumulative thickness in Western Ghats region reaches 800–900 m (Sano et al. 2001). Vedanti et al. (2018) attempted to estimate the characteristic velocity and density of basaltic formations based on Killari borehole data. They used the density log provided by Gupta et al. (2003) for delineating the thicknesses of massive and vesicular strata in each formation for assigning the measured average values for them (Tables 7.13 and 7.14). They reported a weighted mean saturated density of 2.74 g/cm3 for the entire Deccan volcanic sequence pierced by KLR-1 borehole. Corresponding, weighted mean values of saturated P- wave and S-wave velocities are found to be 5.00 km/s and 3.00 km/s respectively. The DSS studies carried out over this terrain too revealed similar low values of Vp (4.50– 5.25 km/s) both in offshore as well as the onshore basalt sections (Dixit et al. 2010; Murty et al. 2010, 2011). As mentioned earlier, average Vp for the Deccan volcanic sequence is only 5.00 km/s, while the weighted density is 2.74 g/cm3 . In normal crust, with such a

240

7 Seismic, Elastic and Petrophysical Properties of Crustal Rocks: …

density, P-wave velocity would be around 6.0 to 6.1 km/s (Christensen and Mooney 1995). Thus, in the Deccan volcanic sequence, the estimated Vp (5.0 km/s) is almost 16% lower than that of the normal crust. Even in the massive basalt cores, with density about 2.90 g/cm3 (Table 7.14), measured Vp is only 5.85–5.93 km/s. With such high-density, Vp value should have been around 6.60 km/s (Christensen and Mooney 1995). Low seismic velocities in high density Deccan basaltic rocks, is not well understood. Such lowering of velocities are often ascribed to fluids, chemical alterations, serpentinization, microcracking, high mean atomic weight, presence of glass contents and vesicles (Birch 1961a, b; Christensen 1968; Moose and Zoback 1983; Tompkins and Christensen 2001; Kern et al. 2009; Ullemeyer et al. 2011; Sun et al. 2012; Tripathi et al. 2012a; Pandey et al. 2016). However, Vedanti et al. (2018) suggested that the lowering of velocities, especially in massive basalt cores, may be attributed to the presence of fine-grained glassy material (Fig. 7.10a), high iron (magnetite) contents (Table 7.14) but more so due to presence of various secondary minerals as an altered product from the pyroxenes and plagioclase, which dominate the lithology (green coloured minerals in Fig. 7.10b). For example, chlorophaeite (Fig. 7.10b), which occurs commonly in the Deccan basalts, along with other secondary minerals, may contain 10–20% water (Parthasarathy et al. 2003; Parthasarathy 2006). Chlorophaeite is a devitrified form of glass, often formed due to hydration of the original glass during rapid cooling of the magma or hydrothermal alteration of ferromagnesium minerals. Samples containing glassy basalts have shown higher order of seismic attenuation. Presence of iron in such rocks increases the density but affect Vp adversely (Pandey et al. 2016). Its enrichment causes significant velocity drop even in high-density rocks, due to increase in mean atomic weight of the rock sample (Birch 1961a, b; Christensen 1968; Pandey et al. 2016). As per Christensen (1968), for a constant density, higher mean atomic weight results in a lower velocity. Almost all the samples studied here are in general characterized by quite high FeOT content (Table 7.14).

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Chapter 8

Seismic Instability and Major Intraplate Earthquakes

8.1 Introduction Consequent to active history of rifting, multiple plume interactions and a number of continental breakups, besides unusual crust-mantle structure, Indian Stable Continental Region (SCR) experiences moderate-to-large intraplate seismic activity since historical times. It contains several rift valleys and mega lineaments, which remain active since at least 1.5 Ga (Rogers and Callahan 1987). Most of the past earthquakes have occurred over the regions belonging to Deccan volcanic province, e.g., 1819 Kachchh (Mw 7.7), 1927 Son-valley (Mw 6.4), 1967 Koyna (Mw 6.3), 1970 Broach (Mw 5.4), 1993 Killari (Mw 6.3), 1997 Jabalpur and the 2001 Bhuj (Mw 7.7) among many others (Mukherjee 1942; Gupta et al. 1972; Mandal et al. 1997; Singh et al. 1999; Rajendran et al. 2008). Occurrence of such a high level of seismic activity is not surprising, as GPS and seismological studies clearly indicated high strain rates over the Indian terrain (Ramalingeswara Rao 2000; Talwani and Gangopadhyay 2001). For example, 0.22 × 10−7 /yr to 3.09 × 10−7 /yr in Bhuj (Kachchh) (Jade et al. 2002), 7 × 10−9 /yr in the exposed part of the Indian plate (Paul et al. 2001) and 2.5 × 10−10 /yr for Indian SCR region (Johnston 1994), compared to much lower strain rates (10−11 /yr–10−13 /yr) in other SCRs of the world (Johnston 1994). More than 20 events of magnitude 5 or more occurred in SCR India during the last 50 years. Many of these earthquakes were quite destructive in nature, leading to huge loss of human lives and properties, for example, the 1993 Killari and the 2001 Bhuj events, that claimed together more than 30,000 human lives (Gupta et al. 2001; Mandal et al. 2004). Locations of some 26 damaging earthquakes, M ≥ 5.0 since 1618, are shown in Fig. 8.1; these events mostly occurred within the intra-continental rifted zones, megasuture/lineaments or even paleo-mobile belts. Geological and geophysical information, however, are available only for a few events (Table 8.1). Seismotectonics of the three major events of recent times, which led to widespread damage to life and property, are discussed below. In some of these areas, recurring seismic activity is reported. © Springer Nature Switzerland AG 2020 O. P. Pandey, Geodynamic Evolution of the Indian Shield: Geophysical Aspects, Society of Earth Scientists Series, https://doi.org/10.1007/978-3-030-40597-7_8

247

248

8 Seismic Instability and Major Intraplate Earthquakes

Fig. 8.1 Map showing epicentres of the past damaging and moderate intensity earthquakes (Ramalingeswara Rao 2000), which occurred in Indian peninsula since 1618 and their relationships with the Deccan volcanic terrain and rift valleys. 1. Bhuj (2001, Mw 7.7)*; 2. Killari (1993, Mw 6.2)*; 3. Koyna (1967, Mw 6.3)*; 4. Jabalpur (1997, Mw 5.8)*; 5. Shimoga (1975); 6. Broach (1970); 7. Secunderabad (1876); 8. Mahad (1967); 9. Bhadrachalam (1969); 10. Khurja (1956)*; 11. Jaisalmer (1991); 12. Malani (1907); 13. Mt. Abu (1969); 14. Son valley (1927, Mw 6.4)*; 15. Kutch (1819, Mw 7.8)*; 16. Anjar (1956, Mw 6.0)*; 17. Paliyad (1938); 18. Bhavnagar (1993); 19. Bhavnagar (1919)*; 20. Midnapur (1964); 21. Satpura (1938, Mw 6.3)*; 22. Bombay (1618)*; 23. Sironcha (1872); 24. Mahabaleshwar (1764)*; 25. Bellari (1843)*; 26. Coimbatore 1900)*. Earthquake magnitudes are taken from the compilation of Mandal (1999) and Mandal and Pujol (2006). *Refers to damaging/disastrous earthquakes. KR: Kachchh rift, CG: Cambay graben, NSL: Narmada Son Lineament, MG: Mahanadi graben, GG: Godavari graben, WCF: West Coast Fault, EDC: Eastern Dharwar Craton, WDC: Western Dharwar Craton, CB: Cuddapah basin, SGT: Southern granulite terrain, BC: Bastar Craton, SC: Singhbhum Craton

8.1 Introduction

249

Table 8.1 Seismic parameters of some the selected Indian SCR earthquakes Earthquake

Lat° (N), Long° (E)

Magnitude

Velocity (km/s) and deptha

Velocity (km/s) at hypocentral deptha

Data source

1. Bhuj (2001)

23.41, 70.23

7.7

Vp = 6.31 at 0 km Vs = 3.85 at 0 km

Vp = 6.98 at 17–19 km Vs = 3.85 at 17–19 km

Mandal et al. (2004)

2. Killari (1993)

18.06, 76.55

6.3

Vs = 3.9 at 1.5 km

Vs = 3.9 at 6.5 km

Rai et al (2003)

3. Koyna (1967)

17.50, 73.70

6.3

Vp = 6.38 at 1.0 km Vs = 3.69 at 1.0 km

Vp = 6.44 at 8.5 km Vs = 3.79 at 8.5 km

Krishna (2006)

4. Jabalpur (1997)

23.00, 80.00

5.8

Vp = 6.7 at 1.0 km

Vp = 6.8 at 34.5 km

Murty et al. (2004)

5. Shimoga (1975)

13.80, 75.30

5.0

Vp = 6.4 at 6.0 km

Vp = 6.8 at 35 km

Kaila et al. (1979); Sarkar et al. (2001)

6. Broach (1970)

21.70, 73.00

5.4

Vp = 6.41 at 2.5 km

Vp = 6.57 at 8.5 km

Kaila et al. (1981b)

7. Secunderabad/Hyderabad (1876)

17.50, 78.50

–

Vs = 3.65 at 2.0 km

Vs = 3.65 at 3.0 km

Rai et al. (2003)

8. Mahad (1967)

18.20, 73.40

5.6

Vs = 3.65 at 1.0 km

Vs = 4.8 at 49 km

Ravi Kumar et al. (2001)

9. Bhadrachalam (1969)

17.90, 80.60

5.7

Vp = 6.2 at 3.0 km

Vp = 6.35 at 9 km

Kaila et al. (1990)

a Depths

are from crystalline basement

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8 Seismic Instability and Major Intraplate Earthquakes

8.2 1993 Killari Earthquake 8.2.1 Introduction The September 30, 1993 Killari earthquake (Mw 6.3), which occurred in the Latur district of Maharashtra, is considered to be one of the deadliest earthquakes of modern times, that killed more than 10,000 people and caused unprecedented damages to properties. The occurrence of this earthquake surprised many as it struck in an area, which was considered stable and free from shocks of high intensity. This region belongs to the northwest segment of the Eastern Dharwar Craton (EDC) of southern India, which is known for Late Archean–Early Proterozoic cratonic growth with low pressure metamorphism, episodic magmatic activity and intermittent remobilization of crustal blocks during the entire Proterozoic (Harish Kumar et al. 2003). After the occurrence of the Killari earthquake, a large number of geological, geophysical and aftershock studies, were undertaken by national and international agencies to study seismotectonics of this earthquake. Later, deep scientific drillings were carried out in the epicentral area of the earthquake. KLR-1 was the deepest borehole, which penetrated 338 m thick suite of volcanic rocks consisting some 4 flows of Ambenali formation and another 4 flows of Poladpur formation (Vedanti et al. 2018), underlain by a 8 m thick layer of Gondwana infratrappeans of non-marine fresh water origin (Parthasarathy et al. 2019), that rested directly over the Neoarchean basement of EDC (Pandey et al. 2009, 2014a, 2016; Tripathi et al. 2012a, b; Tripathi 2015; Pandey 2016). These basement rocks are primarily made up of high density and high velocity, amphibolite to granulite facies retrogressed, exhumed and pervasively metasomatised mid-crustal rocks, containing 2.0 wt% CO2 (Pandey et al. 2009). Detailed lineament patterns studies over this region suggests repeated reactivation of the basement below the earthquake epicentral area (Chetty and Rao 1994).

8.2.2 Geophysical Studies 8.2.2.1

Seismological Investigations

Although this part of the EDC was considered to be stable in the seismic zoning map of India before the 1993 Killari earthquake, paleoseismic evidences (Talwani 1994; Sukhija et al. 1998) indicate occurrence of a number of earthquakes in the past, including a major event during 190 BC–410 AD in the meizoseismal area of 1993 Killari earthquake region (Sukhija et al. 2006). Besides, there also have been several instances of felt earthquakes during 1962, 1967, 1980, 1984 and also between 1992 and 1993 (Gupta et al. 1993). In Indian shield, occurrences of almost all the earthquakes are associated with rift valleys, suture zones, water reservoirs or Quaternary tectonics. But no such association is established for this earthquake of Mw 6.3. Its epicentre was located

8.2 1993 Killari Earthquake

251

at 18.01° N, 76.56° E and estimated focal depth was very shallow at about 6.0 km (Baumbach et al. 1994). As this was one of the deadliest earthquakes ever felt in a stable continental region (SCR) anywhere in the world, it was studied in great detail by many workers (Baumbach et al. 1994; Rajendran et al. 1996; Seeber et al. 1996; Kayal and Mukhopadhyay 2002; Mukhopadhyay et al. 2002 etc.). Centroid focal parameters of this earthquake were estimated as, moment = 1.09 ± 0.07 × 1018 N m with focal depth, strike, dip and rake of 06.5 ± 0.1 km, 127 ± 0.5°, 47 ± 0.2° and 113 ± 0.5° respectively (Ramesh and Estabrook 1998). This event induced largescale surface deformation, groundwater table variation and gas emanations from the opened cracks in several nearby areas. Immediately after the earthquake, a mobile station network was installed by the CSIR-National Geophysical Research Institute, Hyderabad in collaboration with a German task force committee, comprising the Geo Forschungs Zentrum (GFZ), Potsdam and some German universities, to monitor the aftershock activity. Epicenters of the aftershocks, as located by a small aperture 3-station network (Baumbach et al. 1994), revealed that majority of the aftershocks epicentres were clustered in a rectangular area southwest of surface rupture zone (Fig. 8.2a) and a reliable depth section of the recorded events show strike of about 135° E and a dip of ~45° (Fig. 8.2b). Fault plane solutions revealed reverse faulting for the main shock as well as for the deeper aftershocks at depth range 6–10 km, and strike slip faulting for the shallower (depth < 5 km) aftershocks (Kayal 2000; Kayal and Mukhopadhyay 2002). Based on the aftershock trends, fault plane solutions and seismic tomography, a fault intersecting model was proposed. Using the aftershock data, Krishna et al. (1999) reported Vp and Vs of 6.1 km/s and 3.65 km/s, respectively for the uppermost part of the crystalline crust. Subsequently, Rai et al (2003) reported detailed velocity structure by receiver function studies, based on broadband data at the KIL seismic station. The S-wave velocity (Vs) distribution (Fig. 8.3), revealed that from almost 2 km depth downwards, the Vs is between 3.8 and 4.0 km/s down to the Moho (37 ± 2 km), which is equivalent to a high P-wave velocity (Vp) between 6.6 and almost 7.0 km/s. Average Vs for the crystalline crust was estimated at 3.93 km/s and Vp/Vs 1.739 ± 0.047. Expected errors in Moho depth is ±2 km, while in Vs ±0.1 km/s. The estimated velocities in the first 20 km can be considered much higher, compared to that usually encountered in Precambrian shields and platform world wide (Christensen and Mooney 1995). Such higher P-wave velocities (~6.6 to 7.0 km/s) as shown in Fig. 8.3, usually correspond mid to lower crust located at about 15–30 km depth (Christensen and Mooney 1995). This would signify presence of mafic components like granulite and amphibolite facies rocks (Rudnick and Fountain 1995) at basement depths itself. The measured crustal Vs below Killari are close or even higher than those obtained at TRVM (Trivandrum), MDRS (Madras) and KOD (Kodaikanal) seismic stations (Fig. 8.4), situated over the granulitic terrain of southern India (Ravi Kumar et al. 2001; Gupta and Rai 2005), which are largely made up of intermediate to felsic granulites, apart from Koyna.

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8 Seismic Instability and Major Intraplate Earthquakes

Fig. 8.2 a Locations of epicentres of aftershock as obtained from a small aperture 3-station network at Killari, and b SW–NE depth section of the events located within the rectangle shown in a, which is perpendicular to the strike of the fault plane (135° E). Hypocentre of the main shock is shown by an asterisk. Modified after Figs. 9 and 10, Baumbach et al. (1994)

In Fig. 8.3, Vp distribution with depth, as obtained by DSS along Kolattur–Palani section in the medium to high grade (10–11 kb) granulitic terrain of SGT, is also shown (Reddy et al. 2003). It is observed that the computed Vp below Killari is consistently higher down to depth of 30 km, compared to those observed below Kolattur–Palani section of SGT. Presence of similar rocks at shallower depths below Killari (Fig. 7.4, Chap. 7) is also reported by geological, seismic and petrophysical studies (Pandey et al. 2009, 2014a, 2016; Tripathi et al. 2012a, b; Tripathi 2015; Pandey 2016).

8.2 1993 Killari Earthquake

253

Fig. 8.3 P-wave velocity variation with depth beneath Killari (converted from Vs, Rai et al. 2003 using Vp/Vs = 1.739) and its comparison with global shields and platforms (Christensen and Mooney 1995) and Southern Granulite Terrain along Kolattur–Palani DSS section (Reddy et al. 2003). S-wave velocity variation with depth below Killari (Rai et al. 2003) is also included, where dashed line indicates the estimated average crustal Vs of 3.93 km/s Fig. 8.4 S-wave velocity variation with depth below Killari (Rai et al. 2003) and Koyna (converted from Rai et al. 1999 using Vp/Vs = 1.73), together with Madras (Chennai), Trivandrum (Thiruvananthapuram) and Kodaikanal seismic stations located on granulitic terrain of southern India (Ravi Kumar et al. 2001; Gupta and Rai 2005). Modified after Fig. 4, Pandey et al. (2009)

254

8.2.2.2

8 Seismic Instability and Major Intraplate Earthquakes

Gravity Studies

A detailed gravity survey in an area of 26 × 36 km in the Killari earthquake epicentre region was conducted using about one kilometre spacing (Mishra et al. 1994). A contour map, prepared with an interval of one mGal (Fig. 8.5), indicates presence of three broad gravity highs with one of the prominent highs lying south of Sastur village in the epicentral region. It has a gravity anomaly of about +9 to 10 mGal within a short distance of 15 km, which coincides with the upwarped high conductive body identified by MT studies. It also indicates presence of three fault zones and location of the Killari earthquake epicentre falls at the intersecting faults. Several juxtaposed blocks, bounded by deep fracture/fault zones in the epicentral area are

Fig. 8.5 Bouguer gravity anomaly map (in mGal) of the Killari earthquake region. AB is the profile location along which a gravity model is presented in Fig. 8.6. After Fig. 2, Mishra et al. (1994)

8.2 1993 Killari Earthquake

255

Fig. 8.6 Shallow crustal density model along profile AB (Fig. 8.5), indicating presence of mafic rocks (amphibolite to granulite facies) with density 2.79 g/cm3 , lying immediately below Deccan volcanics. Dotted circles and solid line represent the observed and calculated gravity field respectively. Location of KLR-1 borehole is also shown. After Fig. 13, Pandey et al. (2009)

seen to be present here. Subsurface structure, as modelled along a gravity profile AB (Fig. 8.6), indicates upwarping of high density crystalline basement below the Killari seismogenic region.

8.2.2.3

Magnetotelluric Studies

Earthquake activity in seismogenic areas, is often related to the presence of fluids at crustal depths (Gupta et al. 1996; Zhao et al. 1996; Sarma et al. 2004; Mahesh et al. 2012; Shelly et al. 2013). Sarma et al. (1994) and Ramaprasad Rao et al. (2003) conducted MT studies in the epicentral area. An anomalous conductive body (15– 25 m) in the depth ranging between 6 and 10 km, was ascribed to possible presence of fluid-filled rock matrix in the focal depth region (Sarma et al. 1994). A large conductive 2-D structure (100–200 m) was also delineated by Ramaprasad Rao et al. (2003), which extended from the near surface to a depth of 25 km (Fig. 8.7), and located below Sastur village (near the station K3 in Fig. 8.7). High seismic velocities are observed below these conductive features.

8.2.2.4

Terrestrial Heat Flow Studies

Roy and Rao (1999) measured heat flow at four locations in the epicentral area. Out of these, three boreholes were shallow (up to ~180 m depth) and penetrated only the Deccan basalts. These were located 4–8 km away from the surface rupture zone. A heat flow of 33–40 mW/m2 was observed from these boreholes. The fourth

256

8 Seismic Instability and Major Intraplate Earthquakes

Fig. 8.7 Geoelectric depth section below the Killari earthquake region. Modified after Fig. 10, Ramaprasad Rao et al. (2003)

Table 8.2 Summary of heat flow measurements from KLR-1 borehole at Killari Depth interval (m)

Rock types

Temperature gradient (°C/km)

Thermal conductivity (W/m°C)

Heat flow (mW/m2 )

312–336

Basalt

25.4 (0.4)

1.70 (0.03)

43.4 (1.0)

350–549

Granite-gneiss?

14.7 (0.01)

2.92 (0.04)

42.9 (0.6)

549–600

Granite-gneiss?

15.5 (0.04)

2.84 (0.04)

44.0 (0.6)

Mean heat flow = 43 mW/m2 Data source Roy and Rao (1999). Numbers in brackets represent standard errors

borehole (KLR-1), which was the deepest (617 m) and specially drilled in the surface rupture zone, yielded a consistent heat flow of 43 mW/m2 in both basaltic as well as in the basement columns (Roy and Rao 1999). Details of the estimated geothermal parameters in KLR-1 borehole are summarized in Table 8.2. During this investigation, radio elemental (U, Th, K) concentrations were also measured on the core samples by gamma ray spectrometry, which yield a heat production of 0.5 ± 0.2 µW/m3 for migmatitic gneisses and 2.6 ± 0.6 µW/m3 for granites (Roy and Rao 1999). A low mantle heat flow of 12 mW/m2 (defined here as the input of heat flow coming from below the Moho discontinuity) as well as low Moho temperatures of only about 350 °C, was estimated for this location by these

8.2 1993 Killari Earthquake

257

authors. Based on the reasonable crustal heat production model, well constrained by integrated geological and geophysical studies (Table 8.3), temperature-depth variation is estimated below the earthquake epicentral area (Fig. 8.8). This resulted into a high Moho temperature and mantle heat flow of about 540 °C and 32 mW/m2 respectively (Pandey 2016; Pandey et al. 2017), which is much more than estimated earlier by Roy and Rao (1999). Further, the asthenosphere was found at a shallow depth of about 100 km below this region (Fig. 8.8). Table 8.3 Crustal heat production models and adopted geothermal parameters for temperaturedepth estimation beneath Killari, Bhuj and Koyna earthquake regions Depth range (km)

Rock type

Heat production (µW/m3 )

Killari earthquake region (Maharashtra)

Thermal conductivity (W/m°C) Surface heat flow: 43 mW/m2

0–0.34

Deccan volcanics

0.02

1.7

0.34–8.0

Upper amphibolite-granulite

0.78

2.88

8.0–18.0

Granulite

0.2

2.50

18.0–37.0

Mafic granulite

0.16

2.50

>37.0

Ultramafic mantle

0.01

3.0

Data source Roy and Rao (1999), Rai et al. (2003), Ray et al. (2003), Pandey et al. (2017) Bhuj earthquake region (Kachchh, Gujarat) 0–4.0

Sediments

Surface heat flow: 61 mW/m2 1.5

3.5

4.0–6.0

Granite-gneiss

1.82

3.0

6.0–14.0

Amphibolite-granulite

0.78

2.5

14.0–24.0

Granulite

0.16

2.5

24.0–34.0

Magmatic crust

0.02

2.6

34.0–42.0

Magmatic mantle

0.02

3.0

>42.0

Ultramafic mantle

0.01

3.0

Data source Liu and Zoback (1997), Rathore et al. (2004), Ray et al. (2003), Mandal and Pandey (2011), Vedanti et al. (2011), Pandey et al. (2017) Koyna earthquake region (Maharashtra) 0–1.0

Surface heat flow: 41 mW/m2

Deccan volcanics

0.02

1.7

1.0–4.0

Amphibolite-granulite

0.78

2.88

4.0–12.0

Granulite

0.50

2.5

12.0–36.0

Mafic granulite

0.16

2.5

36.0–43.0

Mafic magma

0.02

2.6

>43.0

Ultramafics

0.01

3.0

Data source Gupta and Gaur (1984), Rai et al. (1999), Roy and Rao (1999), Ray et al. (2003), Krishna (2006), Pandey et al. (2017)

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8 Seismic Instability and Major Intraplate Earthquakes

Fig. 8.8 Estimated temperature-depth distribution beneath Killari, Bhuj and Koyna earthquake regions of Deccan volcanic province, based on crustal heat production models as given in Table 8.3

8.3 2001 Bhuj Earthquake 8.3.1 Introduction Small to large magnitude deep crustal earthquakes, occur in several continental rift zones, including Narmada-Son and Kachchh rift zones of India (Mukherjee 1942; Mooney et al. 1983; Prodehl et al. 1994; Johnston 1996; Liu and Zoback 1997; Singh et al. 1999; Kruger et al. 2002; Wilson et al. 2003; Gao et al. 2004; Mandal and Pandey 2010, 2011 etc.). Such rift zones are often associated with large scale magma underplating, crustal thinning and asthenosphere upwarping. The Kachchh rift zone in northwest India is one of them (Fig. 8.9a). It is an east–west trending asymmetrical active continental rift basin, situated in the Kachchh district of Gujarat and filled with 3–5 km thick Tertiary and Mesozoic sediments, which overlie the concealed Precambrian crystalline basement. This basin is severely faulted and bounded by two major faults, north-dipping Nagar Parkar fault in the north and south-dipping Kathiwar fault in the south (Biswas 2005; Merh 1995). Besides, there are several other faults in the region, like E–W trending Allah Bund Fault (ABF), Island Belt Fault (IBF), Kachchh Mainland Fault (KMF) and Katrol Hill Fault (KHF) (Fig. 8.9a, b). Geodynamically, this region is severely affected by the NE–SW compressive stress due to collision of Indian and Eurasian plates. The region is also covered by Deccan volcanics on its western and southern sides. Being an ancient rift basin, it has been

8.3 2001 Bhuj Earthquake

259

Fig. 8.9 a Location of seismic stations (marked by open black triangles) along with the 2001 Bhuj main shock epicentre (star symbol) (details given in Mandal and Pandey 2010, 2011). KMU, Kachchh mainland uplift; ABF, Allah bund fault; IBF, Island belt fault; KMF, Kachchh mainland fault; KHF, Katrol hill fault; NPF, Nagar Parkar fault; BF, Banni fault; GF, Gedi fault and NWF, North Wagad fault. An elliptical area marks the location of central Kachchh rift zone covering KNP, TAP, BHA, NDD and VJP sites. Modified after Fig. 1, Vedanti et al. (2011). b Location of after shocks. The causative fault for 2001 Bhuj earthquake and Gedi fault, are shown by dotted line. Solid star shows the epicenter for the 2001 Bhuj earthquake main shock. Modified after Fig. 2, Mandal and Pandey (2011)

260

8 Seismic Instability and Major Intraplate Earthquakes

persistently active prior to and during the Deccan volcanic episode (Biswas 1987). Local earthquake tomography and seismic refraction studies indicated presence of crustal intrusives and underplated magma at lower crustal levels (Sarkar et al. 2007; Kayal et al. 2002; Mandal and Chadha 2008), besides a prominent low-seismic velocity zone at the depths between 200 and 600 km beneath the Kachchh region (Kennett and Widiyantoro 1999).

8.3.2 Seismic Activity This rift zone remains one of the most prominent earthquake sites since historical times (Rajendran and Rajendran 2001), where deep crustal earthquakes keep on occurring (Mandal et al. 2007; Mandal and Pandey 2010, 2011). Recent paleoseismological investigations indicate that the Kachchh region has been experiencing large magnitude earthquakes (M > 7.0) since 325 BC (Rajendran et al. 2008). The region witnessed two large magnitude earthquakes during the last two hundred years that include 1819 Kachchh earthquake, Mw: 7.7, and the 2001 Bhuj earthquake, Mw: 7.7 (Rajendran et al. 2008). The region was also rocked by a large earthquake in 1956 at Anjar (Mw: 6.0), whose epicentre lied south of the 2001 Bhuj earthquake epicentre (Chung and Gao 1995). Rajendran and Rajendran (2001) reported some 15 historic and recent earthquakes of M 5–6 that took place in this region earlier. In the recent time, the January 26, 2001 Bhuj earthquake (Mw: 7.7, depth 23 km) has been the most devastating SCR earthquake in the world. It was located north of the Kutch Mainland Fault (KMF), about 60 km east of Bhuj (23.412° N, 70.232° E). The fault plane solution of the main event indicated a reverse faulting with a strikeslip component (Kayal et al. 2002). Maximum reported intensity was XI near the epicenter (Hough et al. 2002). It caused unprecedented damage and killed almost 20,000 people (Gupta et al. 2001). Here, the aftershocks are continuing even after two decades of the main event.

8.3.2.1

Recent Seismological Investigations

Seismotectonics of this region, specially the genesis of the 2001 Bhuj earthquake, is reported by many workers based on temporary seismic network data (Kayal et al. 2002; Mishra and Zhao 2003; Bodin and Horton 2004; Mandal 2006, 2007; Mandal and Pujol 2006; Mandal and Chadha 2008); all of them reported large variations in crustal velocities across the aftershock zone. However, a large quantity of aftershock data, some 5000 aftershocks (Mw ≥ 2.0) recorded during 2001–2007 by a semipermanent seismic network of 5–18 broadband stations, were studied by Mandal and Pandey (2010, 2011). This provided high quality arrival times of 13,862 P-waves and 13,766 S-waves from the seismograms of 2303 well-located aftershocks. The data collected by the digital networks of CERI (USA) and Hirosaki University (Japan) were also added to this analysis, thus 58 seismic stations covered the aftershock

8.3 2001 Bhuj Earthquake

261

zone (Fig. 8.9b). First, a 1D velocity was estimated by inversion technique, which was used for 3D Vp and Vs imaging below the epicentral zone using simultaneous inversion tomographic method as described in detail by (Mandal and Pandey (2010). In order to obtain a detailed 1D velocity structure, 3D P- and S-wave velocity distribution were determined below the epicentral zone using tomographic method (Mandal and Pandey 2010). In this method, Vp and Vs tomograms were prepared at different depth slices. These tomograms indicated lateral velocity heterogeneity in the form of high and low velocity patches. The high velocity zones correspond to crustal mafic and ultramafic intrusive bodies, while low velocity zones to entrapped volatiles like CO2 and aqueous fluids. The volatiles and fluids are released through the ongoing degassing of the mantle from the asthenospheric upwarp and consequent magma underplating.

8.3.2.2

Seismic Velocity Structure and Crustal Composition

An average seismic velocity structure, based on Vp and Vs tomograms (Mandal and Pandey 2010) is shown in Fig. 8.10a, which represents the epicentral area between latitude 23.3 and 23.7° N and longitude 69.9 and 70.2° E. The crustal structure beneath the central part of the Kachchh rift zone (Fig. 8.10b) contains 4 km thick sediments, resting over a thin (only 2 km) layer of the granitic-gneissic upper crust characterised by Vp 5.81 km/s. This thin veneer of the upper crust is underlain by an intermediate crust between 6 and 14 km depth and characterised by Vp 6.27– 6.53 km/s. It is followed by a higher velocity (6.62–6.98 km/s) mafic lower crust to a depth of 24 km. This higher velocity would correspond to high grade amphibolite to granulite facies rocks. The lower crust is further underlain by an unprecedented 18 km thick mafic to ultramafic underplated magmatic layer, stacked between 24 and 42 km depth, above the normal ultramafic mantle. This magmatic layer appears to be made up of two parts. First part lies between 24 and 34 km depth above the present Moho. In this layer, Vp increases gradually from 7.15 to 7.63 km/s. The other part of the magma layer is located between the present Moho and the normal ultramafic mantle corresponding to the depth 34–42 km. This mantle magma layer, appears differentiated in which Vp increases from 7.75 to 8.11 km/s. This subcrustal magma layer may contain ultramafic eclogite segregations, as evidenced by sub-Moho reflections (Sarkar et al. 2007).

8.3.2.3

Relocation of Aftershocks

Some 1400 aftershocks are relocated with higher precision, using the velocity model evolved by Mandal and Pandey (2010) (Fig. 8.11). The detailed method and its application is given in Mandal and Pandey (2010, 2011). The average relative uncertainties in aftershocks relocation was found to be about 30 m in epicentral location and 50 m in focal depth estimation. The relocated epicentres are shown in Fig. 8.11a.

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8 Seismic Instability and Major Intraplate Earthquakes

Fig. 8.10 a One-dimensional velocity model for Vp and Vs (marked by thick solid line) derived from the three-dimensional velocity tomograms for an area of 33 km × 44 km beneath the epicentral zone of 2001 Bhuj main shock (23.3°–23.7° N; 69.9°–70.2° E). The values of standard deviation of Vp and Vs (in km/s) are also shown by dashed lines (Mandal and Pandey 2010, 2011). b Inferred crustal cross-section from the 1-D velocity structure. Modified after Fig. 3, Mandal and Pandey (2011)

It reveals aftershock zone encompassing an area of 55 km × 30 km, and a possible NNE trending transverse feature that covers an area of about 25 km × 10 km. Two cluster-trends, one in NE side and the other in NW side, can be seen from this figure which seemingly form a triangular shape. This figure also includes inferred as well as mapped faults/lineaments as shown by pink/red dotted lines. A good correlation is seen between the relocated aftershock epicenters and the geologically mapped faults/major lineaments (Ghevariya and Sahu 2001). Further, a major lineament running from Rapar to Bhachau (RPR–BHA) (Singh and Singh 2005) seems to clearly mark the limit of the relocated aftershocks. Figure 8.11b which shows a N–S hypocentral depth section, based on relocated aftershocks, seem to reflect the geometry of NWF, SWF1 and SWF2 faults (which are marked by pink dotted lines), as well as that of the ITF (a possible transverse

8.3 2001 Bhuj Earthquake

263

Fig. 8.11 a Relocation of 1403 aftershocks of the 2001 Bhuj earthquake (blue open circles) except the events on the inferred transverse fault (ITF) shown by red dots. Big red dot indicates 2001 Bhuj main shock. Solid black lines show geologically mapped major faults. Geologically mapped minor faults, inferred faults and lineaments are marked by the pink/red dotted lines. The inferred transverse fault is marked by black dotted lines. Dashed rectangular area contains number of major lineaments (Ghevariya and Sahu 2001), b N–S hypocentral depth section of relocated aftershocks. Inferred geometry of NWF (north Wagad fault), SWF1 (south Wagad fault 1) and SWF2 (south Wagad fault 2) (locations shown in Fig. a) are marked by pink dotted lines, whereas, that of the ITF is shown by black dotted lines. Crustal layer boundaries are marked by solid grey lines: Sed.: sedimentary layer; Uc: upper granitic–gneissic crust; Ic: intermediate crust; ULc: upper lower crust; LMc: lower mafic crust; M: Moho. Modified after Fig. 2, Mandal and Pandey (2010)

fault), which is shown by black dotted lines. South-dipping NWF extends to about 36 km depth, while the transverse fault ITF, penetrates up to 30 km depth. Similarly, the south-dipping south Wagad faults 1 and 2 (SWF1, SWF2), extend up to 24 km depth.

8.3.2.4

Moho and Lithosphere Thickness

Mandal and Pandey (2011), using local earthquake velocity tomography, and Joint inversion of the receiver functions and surface wave group velocity dispersion, showed strong and clear P- to S-conversions with positive peaks associated with

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8 Seismic Instability and Major Intraplate Earthquakes

top sediments (at 1–3 s) as well as Moho (PS, at 4.5–6 s), and negative peak (Pls at 8.0–12.5 s) corresponding to the LAB (Lithosphere-Asthenosphere Boundary) or the thickness of the lithosphere on the radial RFs. The receiver functions showed a strong negative phase (Pls) at 8.0–12.5 s, which corresponded to LABs. The thicknesses of the crust and the lithosphere were then estimated using the joint inversion of stacked radial receiver functions and surface wave group velocity dispersion data (Julia et al. 2000) for the twelve broadband sites, as given in Table 8.4. This indicates a shallow depth to the Moho (32–37 km; mean: 35.9 km) in the central part of rift Zone, compared to 38–42 km in adjacent areas. Further, the thickness of the lithosphere varied from 62 to 76 km, with an average of 63.1 km in the central rift zone and about 69.0 km in surrounding areas. Estimated crustal and lithospheric thickness below Vajepar (VJP) seismic station in the Kachchh rift zone is shown in Fig. 8.12. The elevated Moho by almost 4–7 km and asthenosphere by about 6–10 km, below the central Kachchh rift zone, may be attributed to anomalous rise of mantle solidus, and presence of a confined body of partial melts at shallow depth below the region. Table 8.4 Estimated crustal and lithospheric thicknesses below the Kachchh seismogenic region Seismic station

Lat. (°N)

Long. (°E)

Moho depth (km)

Lithospheric thickness (km)

BHA

23.28

70.34

36.0

63.0

CHP

23.27

70.28

36.0

63.0

VJP

23.56

70.50

32.0

65.0

NDD

23.32

70.14

36.0

63.0

RAM

23.55

70.47

37.0

64.0

TAP

23.24

70.12

37.0

62.0

KNP

23.40

69.91

37.0

62.0

35.9

63.1

Central rift zone

Mean Surrounding regions VND

23.39

70.40

42.0

68.0

GDM

23.06

70.11

38.0

68.0

GDD

23.87

70.37

39.0

65.0

MTP

23.86

69.78

42.0

68.0

NPR

23.11

69.58

40.0

76.0

40.2

69.0

Mean

Source Mandal and Pandey (2011), Vedanti et al. (2011)

8.3 2001 Bhuj Earthquake

265

Fig. 8.12 Estimation of crustal and lithospheric thickness at Vajepar, located in the central Kachchh rift zone. M represents the Moho and L represents the LAB. Modified after Fig. 4, Vedanti et al. (2011)

8.3.2.5

Terrestrial Heat Flow

Although, the Bhuj region is a unique site of intraplate seismic activity, no heat flow study is made in this region. Vedanti et al. (2011) developed an inverse recurrence method to estimate a first order surface heat flow for this area. This methodology requires the knowledge of depth to the LAB, seismic structure of the crust and the radioactive heat production in the exposed crystalline basement rocks. All such parameters are available in Bhuj region. Following the procedure developed by Vedanti et al. (2011), we arrive at the surface heat flow of about 61.3 mW/m2 for the central part of the Kachchh seismic zone. Although no direct heat flow measurement was reported, based on mantle-derived paleo-geotherm on mantle xenoliths from the alkali rocks, Mukherjee and Biswas (1988) estimated a minimum heat flow of 60 and 70 mW/m2 for this region, following Pollack and Chapman (1977). Empirical relationships (Negi et al. 1987; Chapman and Pollack 1977) would also result surface heat flow of 60–75 mW/m2 , against a lithospheric thickness of about 70 km. As per Sen et al. (2009), the lithosphere may have been still thinner during the alkalic volcanism period in the Kachchh region. Further, presence of high Moho temperatures of ~625 °C is envisaged at the depth of about 34 km, besides a high mantle heat flow of about 43 mW/m2 beneath epicentral zone. This would indicate massive thermal restructuring of the crust and the mantle lithosphere underneath, due to influx of ultramafic melts from below during the course of lithospheric stretching, consequent subcrustal erosion and rise of isotherms. Thermal lithosphere is estimated at the depth of about 70 km, similar to that found by receiver function studies, as mentioned above. Details of the thermal calculations can be found in Pandey et al. (2017). Estimated temperature-depth profile is shown in Fig. 8.8. Measured heat flow averages about 77 mW/m2 in the adjacent Cambay graben (Gupta 1981), having similar crustal structure.

266

8.3.2.6

8 Seismic Instability and Major Intraplate Earthquakes

Seismic Velocity and Earthquake Nucleation

Seismic tomograms (Mandal and Pandey 2010) indicate occurrence of distinct zones of low Vp, low Vs and large Vp/Vs in the mid-level crust, which appear related to the presence of trapped gaseous fluids like CO2 emanating from the underlying mantle (Miller et al. 2004). Pandey et al. (2009) reported that almost 2 wt% CO2 is indeed present in the crystalline basement rocks of Killari. The eclogitisation of the lower crustal olivine-rich gabbroic rocks, is likely to provide aqueous/gaseous fluids (like CO2 and N2 ) in the intermediate to lower crust. Large amounts of CO2 can also be transported upwards in the crust through melts generated from a CO2 -rich lherzolite mantle (Sen et al. 2009), where it can be intrapped in weaker zones. It appears that the unusual crust-mantle structure, deepening of brittle-ductile transition and a high input of volatiles and heat flow from the mantle, control the nucleation of the lower crustal earthquakes below the Kachchh seismogenic region (Mandal and Pandey 2011).

8.3.3 Seismic Reflection Studies Sarkar et al. (2007) processed three 35-km-long seismic-reflection profiles in order to image the variations in regional crustal structure and the Moho beneath this region. A zone of high reflectivity is observed in the lower crust, beginning at ~22 km depth. The crust-mantle boundary is located at a depth of about 35 km at the coast to 45 km in the immediate epicentral region. Broadband seismic study, on the other hand, has shown presence of a much thinner Moho in the epicentral region (Table 8.4). Seismic reflection studies, however, could not image the fault associated with the 2001 Bhuj earthquake. Seismic reflections were observed even from 10 to 15 km depth below the Moho, which could be ascribed to presence of mafic igneous intrusions caused by rifting and lithospheric stretching. As per these authors, thickening of the Moho could be a result of magmatic intrusions related to Mesozoic rifting associated with the breakup of Gondwanaland.

8.3.4 Gravity Field Several attempts have been made to study the gravity field of the Kachchh region (Khan et al. 2016; Chandrasekhar and Mishra 2002; Chandrasekhar et al. 2005; Mishra et al. 2005; Tewari et al. 2009). Recently, Khan et al. (2016) constructed a Bouguer gravity anomaly map of this region based on the data from several sources, which covered Bhuj earthquake region also. It revealed a large variations in the Bouguer gravity anomaly from −24 to +48 mGal (Fig. 8.13). This region is dotted by several gravity highs and lows, which can be primarily attributed to fault controlled basement uplift and depressions. Gravity highs of circular shape can specially be seen

8.3 2001 Bhuj Earthquake

267

Fig. 8.13 Bouguer gravity anomaly map of Bhuj and surrounding areas (based on Mishra et al. 2005; Chandrasekhar and Mishra 2002; Tewari et al. 2009). Red solid star represents the location of the 2001 main shock. Morpho-tectonic features are indicated by dotted lines. BU: Bela uplift, GF: Gedi fault, KHF: Katrol hill fault, KMF: Kachchh mainland fault, KU: Khadir uplift, KWU: Kathiawar uplift, PU: Pachham uplift, SWF: South Wagad fault, VGF: Vigodi fault, WU: Wagad uplift. Blue open circles indicate earthquakes of magnitude more than 5. Modified after Fig. 8, Khan et al. (2016)

in the northern part of the Kachchh region, which in general coincides with the Island Belt Fault. Apart from this, the gravity highs also characterize the western part of the Kachchh Mainland, NE part of the Kathiawar uplift, and several other uplifted regions, which seem related to the occurrences of volcanic plugs or presence of mafic rocks at sub surface depths. Low Bouguer gravity anomalies occur in the south as well as northeastern part of this region, for example, Bhachau Tertiary basin. Sediment thicknesses are expected to be much larger in the southern part which are intimately associated with post-rift vertical tectonics, controlled by thrust faults, having large throw (Chandrasekhar and Mishra 2002; Khan et al. 2016). A steep gravity gradient, roughly coincide with the strike of the Katrol Hill Fault. The location of the Bhuj earthquake as well as the aftershocks, correlate fairly well with the intersection of thrust faults like Kachchh Mainland Fault, south and north Wagad Faults and NW–SE and NE–SW trending structural features (Chandrasekhar and Mishra 2002).

8.3.5 MT Studies Several MT studies are made over this region (Sastry et al. 2008; Naganjaneyulu et al. 2010; Chandrasekhar et al. 2012; Mohan et al. 2015). Abdul Azeez et al. (2018) recently mapped detailed 3D resistivity structure of the aftershock zone. Three dimensional modeling of MT data, were carried out for 37 sites, which provided

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8 Seismic Instability and Major Intraplate Earthquakes

Fig. 8.14 Representative depth slice of the 3D resistivity image between 17 and 27 km below the Wagad aftershock zone of the 2001 Bhuj earthquake. Surface trace of various faults and lineaments (solid/broken lines) mapped in the area are also shown. C1 and C2 are highly conductive zones, while, R1 is resistive zone. Similarly, R2 is moderately resistive zone while zone marked by C3 too is moderately conductive. GF: Gedi fault; KMF: Kutch mainland fault; NWF: North Wagad fault; SWF: South Wagad fault; IBF: Island belt fault. Modified after Fig. 6, Abdul Azeez et al. (2018)

lateral resistivity variation at different crustal depths. The 3D conductivity model revealed a homogeneous and conductive shallow crust mainly representing thick Tertiary and Mesozoic sediments, while the upper and middle parts of the crust was found to be highly heterogeneous showing a blend of conductive (300 m) features. Two major enhanced conductive zones (C1 and C2 in Fig. 8.14) are imaged in the upper to deep crust below the two transverse faults, KTF (Khadir Transverse Fault) and MF (Manfara Fault). The high conductivity zones indicate the loss of compact nature of crystalline crustal rocks, facilitating the accumulation of fluids phases in its pores and fracture zones. These fluids may have infiltrated from the deep mantle through faults and fractures under the seismically active zone, which has been supported through many investigations (e.g. Kayal et al. 2002; Mishra and Zhao 2003; Mandal and Pandey 2010, 2011; Rodkin and Mandal 2012). Conductivity variations in the depth range of 17–27 km within which main earthquake event occurred, is shown in Fig. 8.14. The zone with moderate conductivity (100–200 m) may be related to conductive fluids in the crust which percolated through deep extending fault/fracture systems down to the Moho depth. These anomalies could also be caused by the weak bindings in the rock due to metasomatic alteration, which is also mantle fluid driven and capable of creating low velocity zones (Pandey et al. 2016).

8.4 1967 Koyna Earthquake

269

8.4 1967 Koyna Earthquake 8.4.1 Introduction Reservoir Triggered Seismicity (RTS), is often observed in the area of man-made reservoirs, thereby making it as an anthropogenic problem. Such activities are found worldwide, like Lake Mead (Colorado, USA), Hsingfenking (China), Kariba in the vicinity of Zambia-Zimbabwe and Kremasta in Greece among many others. The Koyna RTS zone of Maharashtra, located near the west coast of India, is also one such region, which came into focus, when it was struck by an earthquake of magnitude 6.3 on December 10, 1967 at Lat. 17.50° N and Long. 73.70° E and claimed about 200 human lives. It devastated Koyna township. It can be considered as one of the best studied RTS site in the world, where seismicity has been continuing since 50 years now, after the impounding in the Shivaji Sagar Lake (Talwani 1997; Rastogi et al. 1997; Talwani et al. 1996; Rai et al. 1999; Agrawal et al. 2004; Mandal et al. 2000; Yadav et al. 2013, 2016; Gupta 2017, Gupta et al. 2017a, b). RTS activity, which earlier occurred mainly in the Koyna region, has now progressively migrated southward around Warna Reservoir, which was impounded in 1985. Every year following the monsoon, water level rises in these reservoirs, which then lead to increase in triggered seismicity. In the last fifty years, 22 earthquakes of M ≥ 5, some 200 earthquakes of M ≥ 4 and several thousands of smaller earthquakes have occurred in this region at the shallow depth between 2 and 9 km within a small area of about 20 km × 30 km (Gupta 2017). Active seismicity of the Koyna–Warna region, together with the location of 23 broad-band seismic stations as well as the 6 borehole seismic stations, is shown in Fig. 8.15. Although, RTS in the Koyna–Warna regions is well established, but the triggering mechanism and the genesis of such earthquakes are still poorly understood, in spite of large number of geoscientific studies carried out in the last five decades. Some researchers (e.g. Lee and Raleigh 1969; Krishna Brahmam and Negi 1973; Kailasam et al. 1976; Rai et al. 1999; Agrawal et al. 2004; Pandey et al. 2009) argued that the occurrences of these earthquakes are related to tectonic origin. This controversy has led to numerous detailed geological and geophysical investigations over the area, including deep scientific drilling initiated with the help of Ministry of Earth Sciences (New Delhi). Till date, nine exploratory boreholes of about 1.5 km depth are drilled which penetrated the basement, including drilling of a 3 km pilot borehole, which was completed in June 2017, where the Deccan volcanics are 1247 m thick. In order to estimate the earthquake parameters in a better way, 6 bore holes were fitted with 3-component seismometers at depths of around 1500 m (Fig. 8.15).

270

8 Seismic Instability and Major Intraplate Earthquakes

Fig. 8.15 Site of Koyna–Warna seismogenic region near the west coast of India. Green curve refer to Western Ghat escarpment. After Fig. 1, Misra et al. (2017)

8.4.2 Scientific Deep Drilling and Nature of Crystalline Basement Nine boreholes (KBH-1 to KBH-9) are so far drilled in the Koyna–Warna seismic zone, which penetrated 479–1251 m thick basaltic column, lying above the crystalline basement (Sinha et al. 2017). However, no infratrappean sediments were found between the basaltic column and the basement. Some 46 flows are recorded in the KBH-7 borehole, where basaltic column is thickest at 1251 m. Out of the nine boreholes, the KBH-1 (17.377° N; 73.741° E, elevation 601 m) penetrated 932 m of basalt, above the Neoarchean eastern Dharwar cratonic basement. The core samples comprise predominantly grey migmatite gneisses, tonalite and quartz monzodiorite. Zircons from granodiorite and monzogranite, yielded a consistent U–Pb ages 2710 ± 63 Ma and 2700 ± 49 Ma (Bhaskar Rao et al. 2017). Thin sections studies commonly show hipidiomorphic granular and porphyroblastic textures. The melanosomes were

8.4 1967 Koyna Earthquake

271

found to be medium grained, with amphibole, epidote, chlorite quartz, plagioclase and opaques. Effect of saussuritization, kaolinization and chloritization are also ubiquitous (Bhaskar Rao et al. 2017), which have commonly been seen in exhumed and retrogressed metasomatised rocks of Killari earthquake region (Tripathi et al. 2012a, b; Pandey et al. 2014a, 2016). The boreholes KBH-5, 6 and 7 too penetrated migmatitic gneisses and amphibolites, apart from granite and granite gneisses. Thin section studies on sheared granitegneiss of KBH-5, indicated presence of highly strained and dynamically recrystallized quartz grains (Misra et al. 2017). It also contained pyroxenes, hornblende, biotite and epidote and showed signatures of saussuritization. Misra et al. (2017) reported that the great majority of basement granitoids beneath the Koyna seismic zone are mainly composed of quartz, feldspar, pyroxene, hornblende, biotite, which may conform to mid-crustal assemblages. These minerals have developed under the amphibolite facies condition, over which subsequent overprinting of recrystallized, muscovite and chlorite as well as of epidotes are often seen. It suggests subsequent deformation under the greenschist facies condition, possibly during retrogressive metasomatic phase. Like the Killari seismogenic region, the Koyna region too seems to have undergone crustal exhumation, causing removal of almost entire sedimentary as well as large chunk of granitic upper crust, and bringing denser mid-crustal lithological facies close to the surface.

8.4.3 Crustal Seismic Studies Two East–west DSS profiles, were shot during 1975–1978 across this region. One of them, between Guhagar on the west coast and Chorochi in the east, passed through the seismically active Koyna region (Kaila et al. 1981a). This profile delineated a deep fault about 5 km west of Koyna Dam (Fig. 8.16); the local seismicity is related to this fault. The velocity (Vp) in the Deccan volcanic sequence varied from 4.7 to 4.9 km/s, while in the basement, 5.9–6.1 km/s. In an another study, Krishna et al. (1989) used normalized digital outputs of analog records and derived a seismic model which showed a Vp 6.2 km/s at a depth of 2.8 km and the Moho at 37 km. This work was subsequently followed by a much detailed investigation, which used digital network of 20 seismic stations. Rai et al. (1999) utilized these data for seismic tomography and provided 1-D velocity model with high Vp of 6.30 km/s at a shallower depth of 2.5 km only. Using the inverse vertical seismic profiling (INVSP) geometry, Krishna (2006) reported a modified velocity model for Koyna region (Fig. 8.17), which indicates much higher Vp of 6.4 km/s at a depth of around 1 km only. Presence of higher velocities are further supported by seismic tomographic studies of Sringesh et al. (2000), who reported 2–5% high velocity anomaly in the crust below Koyna seismic zone. Velocity-depth model produced by Shashidhar et al. (2011), who used VELEST programme, showed lower Vp of around 6.0 km/s in the upper crustal rocks down to 10 km depth, but almost upper mantle-type velocity till the Moho.

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8 Seismic Instability and Major Intraplate Earthquakes

Fig. 8.16 Crustal seismic section along Guhagar-Chorochi deep seismic sounding profile (Kaila et al. 1981a). SP refers to shot points. Modified after Fig. 2, Sarkar and Sain (2017)

Fig. 8.17 P- and S-wave velocity models for the upper crust beneath the seismograph station WR (Koyna–Warna seismic region), as deduced by modeling the INVSP gather. After Fig. 5, Krishna (2006)

8.4.3.1

Receiver Function Studies

Receiver function studies (Satyanarayana 2010), indicated presence of a high Vs 4.0 km/s in the upper crustal section down to 10 km depth, which was attributed to ultra-mafic composition. This was followed by another recent study, where Rohilla et al. (2018) carried out P-receiver function analysis for 18 seismic stations, located in Koyna–Warna region, using teleseismic earthquakes at a range of 30°–90°. They too obtained an unusually high upper crustal Vs of about 4 km/s at 5 km depth, which is comparable to that of lower crustal velocities. These velocities are similar to those

8.4 1967 Koyna Earthquake

273

found by Rai et al. (2003) for the Killari region. Such high velocities would conform to amphibolite facies basement found below Koyna by Misra et al. (2017). Further, high shear-wave velocities would also indicate higher rock strength in the Koyna– Warna region, capable of sustaining higher stresses. The Moho below Koyna–Warna region was found to lie between 37.7 and 42 km, with a small upwarping below seismogenic region.

8.4.4 Gravity Field Gravity studies over the Koyna and other parts of Deccan volcanic terrain was carried out by many researchers (Kailasam et al. 1972; Krishna Brahmam and Negi 1973; Tiwari et al. 2001; Vasanthi and Kumar 2016). Bouguer gravity anomaly map prepared by Tiwari et al. (2001) is reproduced in Fig. 8.18 which shows a large gravity low (around −110 mGal) in Koyna-Karad area (anomaly A). Another prominent NW–SE gravity low of almost similar magnitude is observed near Kurduwadi (anomaly B). The WNW–ESE trending gravity high (anomaly C) is another interesting feature on the map, which has been ascribed to a major uplift around Sangola region (Krishna Brahmam and Negi 1973). It lies between the two negative gravity

Fig. 8.18 Modified Bouguer gravity anomaly map (in mGal) of Deccan volcanic province. Location of Killari and Koyna seismogenic regions are shown by stars. Solid broken line show the location of Kurduwadi rift (marked as 1) and Koyna rift (marked as 2) as postulated by Krishna Brahmam and Negi (1973). Solid line shows the gravity profile from Guhagar-Vikarabad along which the deduced crustal model is shown in Fig. 6.19. (A) Koyna gravity low. (B) Kurduwadi gravity low. (C) Sangola gravity high. Modified after Fig. 3, Tiwari et al. (2001)

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8 Seismic Instability and Major Intraplate Earthquakes

Fig. 8.19 Crustal density model along Guhagar-Vikarabad gravity profile (location shown in Fig. 8.18), showing presence of mafic crust (density 2.79 g/cm3 ) immediately below Deccan volcanics. A 10 km thick underplated layer possibly underlie the Moho beneath Koyna seismic zone. Dotted circles and solid line represent the observed and calculated gravity field respectively. After Fig. 11, Pandey et al. (2009)

anomalies, A and B. and has a magnitude of −65 mGal. A crustal density model along Guhagar-Vikarabad gravity profile (Fig. 8.18) that cuts cross the Koyna gravity low is shown in Fig. 8.19. Pandey et al. (2009) attribute this gravity low to low density Deccan volcanics (2.61–2.65 g/cm3 ) at the surface and a thick (about 10 km) low density (3.20 g/cm3 ) underplated layer just below the Moho. The Koyna gravity low has earlier been attributed mainly to (i) variation in Deccan trap thickness (Kailasam et al. 1972), (ii) low density Gondwana sediments in the Koyna rift valley (Krishna Brahmam and Negi 1973), (iii) variation in crustal layers (Mishra 1989) and (iv) isostatic compensation (Tiwari et al. 2001). Vasanthi and Kumar (2016) also modelled the residual gravity field across Koyna gravity low, and found presence of two thick low-density/low-velocity crustal zones below the seismic zone, the shallower one lying between 5 and 13 km depth while the deeper one, lies just above the Moho between 35 and 43 km depth. The shallower low velocity zone possibly contained mantle metasomatised rocks, to which petrological studies on Koyna basement cores from the boreholes KBH-1 (Bhaskar Rao et al. 2017) and KBH-5 (Misra et al. 2017) supported. Similarly, the LVZ between 35 and 43 km depth, was attributed to highly differentiated/fractionated Deccan volcanic material, not directly derived from the mantle (Pandey and Negi 1987; Negi et al.

8.4 1967 Koyna Earthquake

275

1992, 1993). All the Koyna earthquakes M ≥ 5.0 occurred within the upper lowvelocity/low-density zone, which could be zone of metasomatic alteration.

8.4.4.1

Koyna Seismic Region—A Possible Paleo-Rift

Majority of the earthquakes seems to indicate an intimate relationship with rift valleys and mega shear zones and lineaments (Fig. 8.1). Such regions are often characterised by thick magma underplating, mafic crustal exhumation, retrogression and upper crustal erosion. A graben structure may have existed below this region, but due to regional uplift, the sediments as well as the underlying granitic crust are largely eroded. Available geological signatures, like occurrence of amphibolite facies basement at subsurface depths (Misra et al. 2017) and Quaternary uplifting (Agrawal et al. 2004) support it. It is also consistent with the model of Krishna Brahmam and Negi (1973), as well as with the recent tectonic model (Catchings et al. 2015). RTS may be a reality in Koyna–Warna region, but larger magnitude earthquakes to occur, suitable tectonic environment does exist here to store the stress.

8.4.5 Thermal Regime and Lithosphere Structure There is only one reliable heat flow value available for the Koyna region, which is about of 41 mW/m2 (Gupta and Gaur 1984). Based on heat flow and heat generation studies, Roy and Rao (1999) and Senthil Kumar et al. (2007), advocated presence of low crustal temperatures below the Deccan Volcanic Province. However, thermal studies in recently drilled KBH-1 borehole (Gupta et al. 2015), indicated possibility of higher heat flow up to 50 mW/m2 , with expected temperatures of 130–150 °C at 6 km and 165–225 °C at 10 km depth (Fig. 8.20), which is in departure to that computed by Roy and Rao (1999) and Senthil Kumar et al. (2007). In view the availability of good quality crustal seismic and heat production data, a temperaturedepth estimation was recently made (Fig. 8.8) by Pandey et al. (2017), which reveals Moho temperature of about 605 °C and high mantle heat flow of 31 mW/m2 , using a heat flow value of 41 mW/m2 . The LAB was estimated at around 100 km depth, which is similar to that of Killari. Temperatures estimated by Gupta et al. (2015) conform with the temperatures estimated by Pandey et al. (2017). Details of the parameters used in these calculations are given in Table 8.3 and discussed in detail in Pandey et al. (2017).

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8 Seismic Instability and Major Intraplate Earthquakes

Fig. 8.20 Temperature-depth profile in the upper crust below Koyna seismic zone, for possible range of heat flow, as estimated from the deep borehole KBH-1. After Fig. 7, Gupta et al. (2015)

8.4.6 MT Studies Sarma et al. (2004) interpreted the broad band MT data along a 192 km east–west traverse from Guhagar to Sangole, that covered 16 MT locations. This profile coincides with the DSS profile (Fig. 8.16) and it covers Koyna seismic zone, associated with high order gravity low (Fig. 8.18). The acquired data were subjected to 1D and 2D interpretation. The results of the 2-D modeling are shown in Fig. 8.21, which indicate presence of block structure in the upper crust, with resistive blocks interspersed with low to moderately conductive features at the depths between 5 and 8 km. The upper crustal layers between the depths 10 and 15 km, contains high resistive section (104 m) on the western side that covers Koyna region also. Lower resistive blocks are found in the eastern part (150 km) persisted in southern India till about 90 Ma. Thinning of the Indian lithosphere is combination of both (pre- as well as post Deccan volcanic events), specially when almost all the cratons and their surroundings

9.8 Lithospheric Mantle Deformation Beneath Indian Cratons

319

Table 9.4 Some recent estimates of lithospheric thickness in various geotectonic segments of India Region

Lithospheric thickness/LAB (km)

Method

Sources

Indian shield lithosphere

70–140

P- and S-receiver function, seismological data

Kumar et al. (2013)

Coorg block (western Dharwar craton)

125

MT studies

Abdul Azeez et al. (2015)

Eastern segment of western Dharwar craton

190

MT studies

Abdul Azeez et al. (2015)

Western Dharwar craton

200

MT studies

Malleswari et al. (2019)

Eastern Ghats belt

120

MT studies

Malleswari et al. (2019)

Gravity

Kumar et al. (2014)

MT studies

Gokarn et al. (2004)

Dharwar craton Eastern Dharwar craton

130–180 100

Singhbhum craton

130–140

Gravity

Singh et al. (2015)

Singhbhum craton

95

MT studies

Shalivahan et al. (2014)

Singhbhum craton

86 ± 8

Receiver function studies

Mandal (2017)

Kachchh region (Gujarat)

63–69

Receiver function studies

Mandal and Pandey (2011)

Eastern Dharwar craton

>150

Geological studies

Dongre et al. (2017)

140

Geological studies

Lehmann et al. (2010)

Central India

were continuously remobilized throughout the Proterozoic (Naqvi and Rogers 1987; Rogers and Callahan 1987; Acharya 1997; Pandey and Agrawal 1999), resulting into the formation of mobile belts and later on rift valleys, which were active since at least Mesoproterozoic (Rogers and Callahan 1987). There are also suggestions that a thermal anomaly persisted below the Gondwana lithosphere for a long period of ~150 Ma, before the actual dispersal of the Gondwana fragments (Rogers 1993). This anomaly must have resulted into rising of the mantle solidus, and continental stretching/rifting. Besides, Indian subcontinent traversed over four mantle plumes in quick succession and passed through super mobile phase (20 cm/yr) during 90–53 Ma (Jurdy and Gordon 1984). Such mantle plumes existed in close proximity to cratonic keels (Storey 1995). It also experienced large scale magmatic extrusions (Deccan and Rajmahal Traps), together with a possible K-T boundary bolide impact offshore near Bombay (Negi et al. 1993; Chatterjee et al. 2006), when plate boundaries were getting reorganised in the Indian Ocean, with

320

9 Heat Flow and Lithospheric Thermal Structure

Fig. 9.6 Distribution of the depth to the LAB (lithosphere–asthenosphere boundary) beneath India as inferred by S-receiver function data (Kumar et al. 2013). PCSZ: Palghat-Cauvery shear zone, NSL: Narmada-Son lineament, MBT: Main boundary thrust, SGT: Southern granulite terrain

break up of continental segments. Successive geodynamic events thus led to almost total degeneration of the Indian cratonic lithospheric roots (Fig. 9.7). Due to such tectonothermal events, Indian shield contained enriched, fertile and juvenile mantle underneath (Rogers and Callahan 1987; Chalapathi Rao and Lehmann 2011). In fact, Indian cratons were not strong enough to resist such extensional tectonism. Low surface heat flow, observed in certain ancient cratonic segments like Dharwar craton and Deccan Volcanic Province, does not necessarily mean cool and thick lithosphere underneath as has been widely understood. In any continental region,

9.8 Lithospheric Mantle Deformation Beneath Indian Cratons

321

Fig. 9.7 Schematic diagram showing possible deformation/destruction of subcrustal cratonic roots, caused by upwelling mantle plume. Interaction between the mantle plume and continental lithosphere is far more under the mobile belts. PMZ: Partial Melting Zone, LAB: Lithosphere–asthenosphere boundary. Modified after Fig. 3a, Veeraswamy and Raval (2004)

deep thermal regime is invariably controlled by upper crustal radioactivity containing granitic-gneissic rocks. If it is thick, mantle heat flow component will be much smaller, leading to a thick and cold lithosphere, and if the granitic crust is altogether absent, or very thin as seen below Deccan Volcanic Province, the underlying lithosphere will be thin and warm, with higher component of mantle heat flow. The cause of low heat flow in such cases are explicitly related to low crustal radiogenic contribution towards the build up of surface heat flow. The SGT of the south Indian shield is another finest example. Had there been no erosion of the 20–25 km thick radioactively-rich granitic crust from this region, the heat flow observed at the surface would have been much elevated (Ray et al. 2003). Therefore, it is rather mantle heat flow (not surface) that govern the thickness of the lithosphere. Such a relationship is shown in Fig. 9.8 for some prominent geotectonic localities of the world (Vedanti et al. 2011b). Further, sometimes we tend to estimate lithospheric thickness from xenolithic/kimberlitic data (Ganguly and Bhattacharya 1987, Lehmann et al. 2010; Dongre et al. 2017), but again, they would tell us the thickness of the lithosphere at the time of xenoliths formation and not for the present time, as many geodynamic phenomena apparently took place between the xenolith formation and now.

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9 Heat Flow and Lithospheric Thermal Structure

Fig. 9.8 Relationship between mantle heat flow and the depth to the lithosphere–asthenosphere boundary (LAB), based on different geotectonic units of the world. Modified after Fig. 5, Vedanti et al. (2011b)

9.9 Distribution of Geothermal Springs Large number of geothermal spring occur over the Indian subcontinental terrain (Fig. 9.9), in which temperatures vary from 30 to 100 °C. These non-volcanic type of geothermal springs are mainly distributed along continental margins, tectonic plate boundaries and rifted basins. Most prominent thermal manifestations may be grouped into three regions, NW-NE Himalayan region, Son-Narmada Tapti rift zone, and west coast continental margin. In some geothermal areas, reservoir temperature could reach between 100 and 200 °C, thus quite exploitable. Total geothermal power potential as assessed by number of geoscience organisations, is of the order of 2000– 10,000 MW; detailed reviews can be found in Ravi Shankar et al. (1991), Pandey and Negi (1995), Pandey (1996) and Chandrasekharam and Chandrasekhar (2015). (i) NW and NE Himalayan Region This region contains number of hot springs, specially in NW Himalayan segment, which are suitable for both direct and indirect use. They were extensively studied by Ravi Shankar et al. (1991). The region contains several promising areas like PugaChhumathang, Beas, Parbati, Satluj and Spiti valleys and Tapoban region. Out of these, Manikaran geothermal area, situated in the Parbati valley, is studied extensively. Exploratory drilling indicated temperature of the thermal water discharge at 86–94 °C, indicating a promising source for geothermal electricity generation. (ii) Son-Narmada Tapti Rift Zone This region has a long history of tectonic reactivation and magmatic activity. It contains 46 thermal spring areas, the most promising being Tatapani region, which is situated in Surguja district of Madhya Pradesh. Here, thermal spring temperatures vary from 50 to 98 °C. EM data modelling indicates presence of a hot water fracture zone at a depth of about 400 m (Harinarayana et al. 1988). A deep thermal reservoir is likely to exist here between the depths of 1 and 3 km.

9.9 Distribution of Geothermal Springs

323

Fig. 9.9 Distribution of thermal spring localities (as listed in Ravi Shankar et al. 1991) shown by solid dots. Son-Narmada-Tapti lineament rift zone is represented by thick broken line along which large number of hot springs occur. ADFB: Aravalli-Delhi fold belt

(iii) West Coast Thermal Province This is anomalously hot geothermal province, covering the areas of Konkan and Cambay geothermal provinces (Fig. 6.11, Chap. 6). Out of these two thermal fields, Konkan region is located in a narrow belt between the Western Ghats and the west coast of India. It contains about 60 thermal springs, having a temperature of 34–71 °C which spread for a distance of almost 300 km. The Cambay region is a prominent rift structure, crossed by several deep seated faults. It too contains a number of hot springs. As mentioned above, heat flow is quite high averaging 83 mW/m2 in

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its northern part. During drilling, hot water with steam was encountered in a few boreholes. At 3 km depth, expected temperature is about 175 ± 25 °C. Both the regions are associated with shallow asthenosphere lying at the depths of about 30– 70 km only (Fig. 6.11, Chap. 6; Table 9.3).

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Chapter 10

Indian Crust

10.1 Introduction Perceptions about the nature of the Earth’s crust in general and the continental crust in particular, is of critical importance to our understanding of its time-bound evolution, as well as interpretation of acquired geological and geophysical data. Almost 5– 20 km thick (or even more) upper portion of the continental crust is eroded away from several Archean-Proterozoic terrains in India. Similar situation may possibly exist in other shields and platforms as well, where structural fabric and composition, have undergone noticeable change. All the geologic terrains globally, other than shields and platforms, are composed of felsic upper crust, characterized by an average Pwave velocity of less than 6.2 km/s (Rudnick and Fountain 1995). In comparison, upper parts of the crust in shields and platform regions, are associated with a much higher velocity of 6.3 ± 0.2 km/s. Similarly, middle crust in shields and platforms is also characterized by higher velocities of 6.5–6.9 km/s compared to 6.2–6.5 in other terrains. An averaged crustal seismic section below shields and platforms, based on large number of profiles (Rudnick and Fountain 1995), is shown in Fig. 10.1. Fig. 10.1 Average continental crustal section below shields and platforms and associated P-wave velocities (in km/s), based on Rudnick and Fountain (1995)

© Springer Nature Switzerland AG 2020 O. P. Pandey, Geodynamic Evolution of the Indian Shield: Geophysical Aspects, Society of Earth Scientists Series, https://doi.org/10.1007/978-3-030-40597-7_10

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10.2 What Is Crust? Untill 1909, it was not known that the Earth has a crustal layer. It was discovered by a seismologist named Andrija Mohorovicic, who observed a sudden change in seismic velocity across a discontinuity located at about 50 km depth. He also realised that this outer layer, or the crust, was made up of less dense material (rock), while the next underlying layer, the mantle, was far more denser. This discontinuity, separating the crust and the underlying mantle, was later on named after him as Mohorovicic discontinuity or “Moho”. The crust may be defined as a thin outer-most rigid layer of the Earth, constituting less than 1% of the Earth’s volume. It consists all the three major variety of rocks viz., igneous, metamorphic, and sedimentary, and the silicate minerals are its main constituent. Almost every known rock types on the surface of the Earth, occurs in the continental crust. The crust is underlain by the upper mantle, which is largely made up of the rocks like peridotites and dunites. The underlying upper mantle, is considered highly buoyant and responsible for the plate tectonic processes, which is accountable for creating three types of crust, oceanic, continental and transitional, the latter one being mostly associated with ocean-continent boundaries. Formation of the crust is an ongoing process and thus, it is still growing slowly.

10.2.1 Oceanic Crust It is hidden beneath the ocean and occupies almost 71% of the total Earth’s area. On an average, it is no more than 7–10 km thick beneath the ocean floor and no older than about 180 million years either. Oceanic crust is formed by the infusion of mantlederived high density (~3.0 g/cm3 ) basalt and gabbroic rocks at the spreading axis of the Mid Oceanic Ridges, where plates are pulled apart. The oceanic crust, together with its mantle lithosphere, gets subducted at the oceanic trenches, which are 5–11 km deep, and situated close to active plate margins, popularly known as subduction zones (Stern 2002), as discussed in detail in Chap. 1. One of such prominent regions include Pacific plate margin (Tonga-Kermadec, Izu-bonin, Japan, Kurile, Aleutians, North Island of New Zealand etc.), which are still active even now. Andes of south America can be considered another example (Fig. 1.3, Chap. 1). Magmatism associated with subducting slabs (Fig. 1.4, Chap. 1), are solely responsible for building island arcs. This would mean that the oceanic crust goes through a cycle of creation as well as destruction. Among the terrestrial planets, possibly only the Earth appears to have subduction zones and plate tectonic process. Being mafic in nature, the oceanic crust contains much lesser SiO2 (48–52 wt%) than the continental crust (Table 10.1). Its average density is close to 2.90 g/cm3. Further, if we leave the thin sedimentary column, the oceanic crust can be divided into two main parts, upper oceanic crust and the lower oceanic crust. Upper oceanic crust is composed of basalts, pillow lavas and dykes, formed due to cooling of the

10.2 What Is Crust?

333

Table 10.1 Average measured SiO2 and heat production in three-layered continental crust. Estimated thicknesses of these layers in shield and platform areas are also included Crustal section

Thicknesses in shield and platform areas (km)

SiO2 (wt%)

Heat production (µW/m3 )

Continental upper crust

13.7

67

1.6–1.7

Continental middle crust

13

64

0.7–1.0

Continental lower crust

12.1

53

0.2

Continental upper mantle (peridotite)

40–700

69

13

>40

0–5

0–2

Moho

Depth (km)

Pressure (kb)

Upper mantle

Underplated

Lowermost

Lower

Middle

Upper

Division

Ultramafic

Mafic

Mafic

Mafic

Intermediate

Felsic

Nature

Peridotitic

Magmatic

Granulitic

Granulitic (and amphibolitic)

Greenschist/amphibolitic (and granulitic)

Granitic-gneissic

Composition

7.7 or more

7.0–7.6

6.9–7.0

6.6–6.9

6.3–6.6

5.7–6.3

P-wave velocity (km/s)

3.3 or more

3.05–3.20

3.0–3.05

2.9–3.0

2.8–2.9

2.67

Density (g/cm3 )

Table 10.7 Generalised sub-division of Indian continental crust, based on recently carried out integrated geological, geophysical, seismic and petrophysical studies, as discussed in various chapters

344 10 Indian Crust

10.7 Sub-division of the Indian Crust

345

In general, Tertiary sediments are characterized by slower velocities of 2–3 km/s (Tewari et al. 1991). The Gondwana and Proterozoic sediments, on the other hand, are characterized by much higher P-wave velocity ranging from 3.8 to 4.2 km/s (Table 1.1, Chap. 1) and 5.2–5.6 km/s (Chandrakala et al. 2017) respectively. Density does not exceed 2.7 g/cm3 , unless these are covered by thick overburden of sediments for example, the Bay of Bengal. Due to lower density, these regions are associated with negative gravity anomalies, unless they are infected by magmatic extrusions, for example, the Cambay graben in northwest India. (ii) Upper crust If we exclude sedimentary strata, in general, upper crust is made up dominantly of felsic rocks like granite and gneisses. However, in due course of time, significant part of the upper crust has been eroded away from the many segments (Fig. 10.5). Thus, some segments of the Indian shield like, Southern Granulite Terrain (SGT) (located in southernmost part of South Indian shield), Eastern Ghats Belt located close to eastern margin, and the Sandmata complex of the Aravalli Fold Belt (northwest India), directly expose the lower parts of the crust at the surface. In the granitic-gneissic upper crust, P-wave velocities vary from 5.7 to less than 6.3 km/s, while S-wave velocities seldom exceed 3.6 km/s (Table 10.2). Being felsic in nature, average heat production is considerably high at about 1.6–1.8 µW/m3 (Rudnick and Fountain 1995) due primarily to rich-ness in uranium, thorium and potassium contents. In most of the Indian terrain, its thickness, on an average, may not exceed 5 km, as most of it has been eroded away. (iii) Middle crust The middle crust lies in between the granitic-gneissic upper crust and the underlying mafic lower crust, and thus it has an intermediate composition. Geophysically, it is not well defined. Most of our informations about this layer, are based on studies on exposed sections of high grade terrains and deep crustal seismic studies. The middle crust is highly evolved due to compressional uplift, exhumation and retrogression. It is characterised mainly by amphibolite to lower-grade granulite facies rocks, with minor amount of tonalite, granodiorites and migmatitic gneisses (Fig. 7.4, Chap. 7) but, these can also contain small quantities of metapelite (Christensen and Mooney 1995). The amphibolite facies rocks are, however, the dominant constituent, although there is not much difference between the bulk composition of amphibolites and granulites. This layer is characterised by a lower heat production averaging around 0.78 µW/m3 (Pandey et al. 2017), but it can range between 0.7 and 1.00 µW/m3 (Table 10.1). Seismic velocities, Vp and Vs usually range from 6.3 to 6.6 km/s and 3.6 to 3.8 km/s respectively (Table 10.2). In earlier crustal models, it represented 15–25 km depth (5–8 kb), but as per our present understanding, mid-level crust, can be seen even exposed in many Indian terrains, including, Deccan Volcanic Province. This layer is very prone to metasomatic alteration, and in such cases, in situ velocities could be much lower. In case of Killari, basement rocks have shown on average a P-wave velocity of only 6.17 km/s, consequent to metasomatic effects.

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10 Indian Crust

(iv) Lower crust Lower crust is partially exposed at several places on the surface of the Earth, including in India. Southern Granulite Terrain (SGT), located in the southernmost part of Indian shield, is one of the finest example. It is characterised by 8–10 kb pressure and thus represent 25–30 km deep section. As mentioned earlier, the Sandmata complex in the Aravalli Fold Belt, is another prime example (Figs. 4.1 and 4.3, Chap. 4). This part of the crust is highly mafic in nature and largely composed of highgrade mafic granulites and a small quantities of amphibolite, metapelite and lower facies of granulites. It is characterised by extremely low heat production of about 0.20 µW/m3 (Table 10.1) (Ray et al. 2003; Rudnick and Fountain 1995). The Vp in this layer typically range from 6.6 to 6.9 km/s and Vs , 3.8–3.95 km/s. (v) Lowermost High velocity crust In some geologic terrains, a high velocity layer lies just below the lower crust, and it is difficult to distinguish. It is mainly composed of Mafic granulites and amphibolites and high grade (granulite facies) metapelites. In view of high radioactivity, metapelites may be present only in minor quantities (Rudnick and Fountain 1995). In this layer, Vp range from 6.9 to 7.0 km/s and Vs 3.95 to 4.0 km/s respectively. (vi) Magmatic Lower crust In many geotectonically active areas, lower crust is often seen underplated by thick mafic magmatic layer, formed by the stacking of fractionated basaltic magma, infused from the upper mantle. It is characterized by high Vp in the range 7.0–7.6 km/s. In some parts of the Indian terrain, this layer is as thick as 15–20 km (Fig. 2.12, Chap. 2). Its presence is fairly common in geodynamically active area like Aravalli fold belt, Kachchh rift zone, Cambay graben, Cuddapah basin and Gondwana grabens. In fact, many of the geotectonic active areas of India are characterised by such high velocity layers above the Moho (Fig. 10.5). Many times, however, no magma underplating is recorded between the lower crust and the Moho. Under such situations, the region immediately below the Moho, is usually characterised by a sharp velocity jump in Vp from 6.9–7.0 to 7.7 km/s or more. (vii) Mantle Magma layer Mantle magma layer is a new find by Mandal and Pandey (2011) and Pandey et al. (2013). This layer is located between Moho and the normal ultramafic mantle. The nature of this part of the mantle is often gradational due to strong crust–mantle interaction caused by rising thermal plume. This section of the transitional mantle, which is largely made up of magmatically altered peridotite rocks is characterized by Vp between 7.6 and 8.0 km/s and Vs from 4.35 to 4.6 km/s. In India, occurrence of such layers have been found below the Kachchh rift region of western India, where 2001 Bhuj earthquake occurred (Fig. 10.5a), and also in south Indian shield (Fig. 10.5c). Below this magma layer, normal ultramafic mantle has a constant P-wave velocity hovering around 8.0–8.1 km/s.

10.7 Sub-division of the Indian Crust

347

As discussed in various chapters, numerous crustal seismic sections indicate that the Indian crust, on an average, is largely mafic. It suggests that a large part of the Indian crust from the top, is eroded away due to geodynamic processes, thereby exposing the mid-level crust in various segments. It was earlier believed that the Deccan mantle plume caused massive crustal uplift and lithospheric erosion, but deep borehole studies have now confirmed that a large chunk of the upper crust in peninsular shield, had been eroded away even before the commencement of Deccan volcanism (Pandey 2016).

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