The Indian Ocean nodule field: geology and resource potential

V O LU M E T E N HANDBOOK OF EXPLORATION AND ENVIRONMENTAL GEOCHEMISTRY The Indian Ocean Nodule Field: Geology and R...

Author: Ranadhir Mukhopadhyay | Anil K. Ghosh | Sridhar D. Iyer

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V O LU M E

T E N

HANDBOOK OF EXPLORATION AND ENVIRONMENTAL GEOCHEMISTRY

The Indian Ocean Nodule Field: Geology and Resource Potential

HANDBOOK OF EXPLORATION AND ENVIRONMENTAL GEOCHEMISTRY Series Editor

MARTIN HALE 1.

Analytical methods in geochemical prospecting

2.

Stastistics and data analysis in geochemical prospecting

3. Rock geochemistry in mineral exploration 4.

Regolith exploration geochemistry in tropical and sub-tropical terrains

5.

Regolith exploration geochemistry in arctic and temperate terrains

6.

Drainage geochemistry

7.

Geochemical remote sensing of the sub-surface

8. Life cycle of the phosphoria formation: From deposition to the post-mining environment 9. 10.

Biogeochemistry in mineral exploration The Indian Ocean nodule field: Geology and resource potential

V O LU M E

T E N

HANDBOOK OF EXPLORATION AND ENVIRONMENTAL GEOCHEMISTRY

The Indian Ocean Nodule Field: Geology and Resource Potential RANADHIR MUKHOPADHYAY National Institute of Oceanography, Goa, India

ANIL K. GHOSH University of Calcutta, Kolkata, India

SRIDHAR D. IYER National Institute of Oceanography, Goa, India

AMSTERDAM • BOSTON • HEIDELBERG • LONDON NEW YORK • OXFORD • PARIS • SAN DIEGO SAN FRANCISCO • SINGAPORE • SYDNEY • TOKYO

Elsevier Radarweg 29, PO Box 211, 1000 AE Amsterdam, The Netherlands Linacre House, Jordan Hill, Oxford OX2 8DP, UK

First edition 2008 Copyright # 2008 Elsevier B.V. All rights reserved. No part of this publication may be reproduced, stored in a retrieval system or transmitted in any form or by any means electronic, mechanical, photocopying, recording or otherwise without the prior written permission of the Publisher. Permissions may be sought directly from Elsevier’s Science & Technology Rights Department in Oxford, UK: phone: (þ44) 1865 843830, fax: (þ44) 1865 853333, E-mail: permissions@elsevier. com. Alternatively you can submit your request online by visiting the Elsevier web site at http:// elsevier.com/locate/permissions, and selecting Obtaining permission to use Elsevier material. Notice: No responsibility is assumed by the publisher for any injury and/or damage to persons or property as a matter of products liability, negligence or otherwise, or from any use or operation of any methods, products, instructions or ideas contained in the material herein. Because of rapid advances in the medical sciences, in particular, independent verification of diagnoses and drug dosages should be made. Library of Congress Cataloging-in-Publication Data A catalog record for this book is available from the Library of Congress British Library Cataloguing in Publication Data A catalogue record for this book is available from the British Library ISBN 978-0-444-52959-6 ISSN 1874-2734 For information on all Elsevier publications visit our website at books.elsevier.com

PRINTED AND BOUND IN THE NETHERLANDS 08 09 10 11 12 10 9 8 7 6 5 4 3 2 1

CONTENTS

Preface Series Editor’s Foreword About the Authors

1. The Indian Ocean Nodule Field 1. Physiography and Geology 2. Physics, Chemistry and Biology 3. Evolution of the Indian Ocean Nodule Field

2. Tectonics and Geomorphology 1. Ridge-Normal Lineaments 2. Ridge Parallel Lineations and Anomalies 3. Seamounts

3. Volcanics 1. 2. 3. 4.

Major Volcanics Minor Volcanics Alteration of Volcanics Tectonics and Volcanics: Interrelations

4. Sediments 1. Distribution and Source 2. Sedimentary Processes

5. Ferromanganese Deposits 1. 2. 3. 4. 5.

Nodule Characteristics Factors Influencing Nodule Formation Dynamics of Nodule Formation Ferromanganese Encrustation The World Oceans Scenario

vii xi xiii

1 5 25 31

37 40 45 54

67 69 86 103 112

115 117 135

155 157 184 194 205 211

v

vi

Contents

6. Resource Management 1. 2. 3. 4. 5. 6.

Resource Identification Mining Technology Environmental Impact Assessment Metal Extraction and Processing Law of the Sea Global Perspectives

References Author Index Subject Index Colour Insert between pages 18 and 19

225 225 231 233 249 254 255 261 283 289

PREFACE

Oceans, with their enormity and mystery covering 259 million km2 and some 72% of the earth’s surface, have always been a part of human consciousness. Climate and weather, and even the quality of the air people breathe, depend in great measure on the interplay between the ocean and the atmosphere in ways still not fully understood. Besides being a prime source of nourishment for life it helped to generate, the ocean is considered as a storehouse for several living and non-living resources. Making use of such resources, in general, and minerals, in particular, essentially needs intelligent and sustainable treatment. The 1982 United Nations Convention on the Law of the Sea, since its entry into force on 16 November 1994, provided for the first time a universal legal framework for the rational management of marine resources and their conservation. The Central Indian Ocean Basin (CIOB), bordered by the Indian Subcontinent, the Ninetyeast Ridge and the Indian Ocean Ridge system, is the largest and most unique of all the basins in the Indian Ocean. Because of its complex tectonic fabric, vastness and ferromanganese deposits, this basin has attracted the attention of oceanographers since long. However, in comparison to the quantum of investigations carried out in the Pacific and the Atlantic oceans, the study in the CIOB has left much to be desired. Even the international drilling programmes (Deep Sea Drilling Project, Ocean Drilling Programme) have not extensively covered this basin. As a result, a paucity of data has limited a proper understanding of the basin and the evaluation of its resources. During the last two decades, the CIOB was explored in some detail for its mineral resource potential, with India taking the lead. The broad objectives were to explore and delineate economically feasible manganese nodulebearing areas, and ultimately concentrate on resource exploration from the Self-allocated area. Extensive exploration for the resources saw more than 50 oceanographic voyages to this basin. The amount and type of underway data collected are immense, ranging from single-beam to multi-beam bathymetry, to seafloor magnetism and gravity and to underwater photography. A large number of manganese nodules, rocks, sediments and water column samples were recovered from an area more than 700,000 km2, bordered between 9 S and 16 300 S and 72 E and 80 E. This helped delineate a nodule-rich area to be later known as the Indian Ocean Nodule Field (IONF). Subsequently, close-grid exploration, followed by the critical resource assessment of the IONF, identified the ‘first generation mine site’. Globally, the IONF represents the second largest and second richest deep-sea manganese nodule resources after those of the Equatorial North Pacific. vii

viii

Preface

A conservative estimate places the total manganese nodules availability in the IONF as more than 1400 million tons, with an average abundance of little more than 4.5 kg of nodules per square metre. Considering the recovery capacity of available mining technology and metal extraction processes, the nodule resources are expected to last for many hundreds of years, and resource mining would be economically viable too. The publications resulting from the field and laboratory studies, and covering diverse aspects (such as evolution of the CIOB, structure and tectonics, volcanism, sedimentation, and resource characterisation and potential), are indeed numerous. Several workers from different institutions in various countries have published their important findings, but these are scattered in different journals, all of which are not easily accessible to researchers interested in the nodule resources of the Indian Ocean. In this scenario, the basic purpose of the present book is to collate the available information in a concise and systematic manner, and carry out a critical evaluation of such information. The necessity of this book is all the more relevant because the currently available books on manganese nodules largely pertain to the Pacific and the Atlantic oceans, making inadequate coverage of the Indian Ocean, in general, and the IONF, in particular. Hence, this book is targeted at students and researchers at the master’s and doctoral levels, and for all those interested in the marine geology and resource potentiality of the IONF. We hope that this book will serve as a ready reference to those involved in the development and management of geo-resources, and to economists and policy planners. After introducing the IONF in the first chapter, we discuss the geomorphology and tectonics of this field in the second and volcanics in the third chapter. The bottom sedimentary regimes—sediment source and its characteristics, dynamics of sedimentation—are discussed in Chapter 4. The distribution, grade and processes of formation of ferromanganese deposits (both nodules and crusts) during the last 5–10 million years are discussed in Chapter 5. And finally in Chapter 6, we describe the equipment used for exploration, sketch on probable mining methods, assess environmental implications concerning mining, discuss metallurgical techniques and converse on international legal constraints on exploration and exploitation of the nodules. While writing this book, we received help from many quarters. The Council of Scientific & Industrial Research (CSIR) and the Ministry of Earth Sciences, both in New Delhi, India, are thanked for all the logistic and financial support—in the laboratory and at sea—during the last two decades through grants to the National Institute of Oceanography (NIO), Goa. The director, National Institute of Oceanography, and the vice-chancellor, University of Calcutta, are specially thanked for their support and encouragement. We place on record with thanks the generous support of all the members of the project ‘Surveys for Polymetallic Nodules’ in helping one way or the other during collection and analysis of samples, and data interpretation. Many of them readily provided their reprints, preprints and other

Preface

ix

unpublished material. The cooperation of the crew of the various ships, sometimes under excruciating conditions, is acknowledged. India’s deep-sea mineral campaign owes a great deal to the vision and untiring efforts of the late Dr Hassan Nasiem Siddiquie, the former director of the National Institute of Oceanography. We respectfully remember his initiation of the manganese nodule programme for India. Dr. S. Z. Qasim, as former Secretary of the then Department of Ocean Development and ex- member, Planning Commission, Government of India, has always been a source of inspiration to the project. The earlier versions of the chapters received critical reviews and comments from A. L. Paropkari, G. V. Rajamanickam, J. N. Pattan, K. S. Krishna, M. Shyam Prasad, M. V. Ramana, N. H. Khadge, R. K. Drolia, R. P. Das, R. V. Karanth, and V. Ramaswamy. We thank them all. Besides, we had rewarding discussions with many of our colleagues at NIO and at the University of Calcutta. Technical help from A. K. Saran, A. Y. Mahale, R. Uchil and S. Akerkar is acknowledged. We thank Elsevier for publishing this book and also Martin Hale, Series Editor, for patiently going through the manuscript and for writing the Foreword. We acknowledge the support (and patience) of Particia Massar, Pauline Riebeek and Tirza Van Daalen during preproduction stage, to Conny Krainz for supervising the production of the Title and Prasenjit Bakshi and his dedicated team for printing the book. Most of the materials presented here are in public domain. Many of the figures were reprinted from various books and journals. We thank the authors and publishers for kindly permitting us to reproduce these figures. Last but not the least, we express a deep sense of appreciation to the sacrifice of our wives Sumita, Jayanti and Kamakshi, respectively, and to our children. It is a pleasure to place on record our gratitude to our colleagues both from within and outside India who helped with literature support and liberal comments to help characterise the resources and to better understand the formational regimes of ferromanganese deposit in the IONF. However, any discrepancy regarding the views and opinions expressed in this book is solely the responsibility of the authors. 31 January 2007 Ranadhir Mukhopadhyay National Institute of Oceanography, Goa [email protected] Anil K. Ghosh University of Calcutta, Kolkata [email protected] Sridhar D. Iyer National Institute of Oceanography, Goa [email protected]

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SERIES EDITOR’S FOREWORD

On 13 March 1874, in a part of the Pacific Ocean between Hawaii and Tahiti, the crew of HMS Challenger raised a bucket of ocean-floor sediment from a depth of 4727 m and discovered that it contained potato-shaped nodules. This event marks the starting point of the ocean-floor research that has yielded a wealth of scientific information, including that presented in this latest volume of the Handbook of Exploration and Environmental Geochemistry. During the period 1872–1876, HMS Challenger made a 127,500-km circumnavigation of the globe, taking deep-sea soundings, bottom dredges, open water trawls and serial water temperature observations in the Atlantic, Indian, Pacific and Antarctic oceans. It was conducting the world’s first oceanographic expedition which, in the words of the supervising scientist of the expedition report, led to ‘the greatest advance in the knowledge of our planet since the celebrated discoveries of the fifteenth and sixteenth centuries.’ Although HMS Challenger, formerly a frigate of the British Royal Navy, had a team of scientists on board and had been fitted out with laboratories, the potato-shaped nodules remained little more than a scientific curiosity for almost 20 years. Then, in 1891, they were analysed in Britain and found to comprise mainly oxides of manganese and iron and to contain unusually high concentrations of nickel, copper and cobalt. Further, oceanographic observations over the next century gradually established that such nodules were, to various extents, abundant in the low-latitude zones of the deep oceans. The Indian Ocean floor began to receive particular attention from the early 1970s, when the French research ship Marion Dufresne recovered 450 ocean-floor sediment cores. At about the same time, ocean-floor nodules were recognised as mineral resources of potential economic value. Commercial interest soared and entrepreneurial ocean-floor mining consortia made substantial investments in innovative technologies for nodule exploration, mining and mineral processing—even though a legal framework for the exploitation of the ocean floors was lacking until the United Nations finally concluded its Law of the Sea Treaty in 1982. Subsequently, the UN International Seabed Authority allotted a 1.5 million km2 site of the floor of the central Indian Ocean to India. The country’s National Institute of Oceanography then set out to map, characterise and evaluate the nodule resources of this region, the best-endowed portion of which has become known as the Indian Ocean Nodule Field (IONF). xi

xii

Series Editor’s Foreword

In this volume, Ranadhir Mukhopadhyay and Sridhar D. Iyer, senior scientists at the National Institute of Oceanography, and Anil K. Ghosh, an eminent academician in the field of mineral resources at University of Calcutta, bring together the extensive but hitherto scattered scientific results of the exploration of the IONF since the 1980s. Their first chapter explains the general physiography, geology, biology, physics and chemistry of the seafloor occupied by the IONF and some characteristics of the overlying water column. The subjects of their next two chapters are the structural, tectonic and volcanic features of the IONF. In Chapter 4, they describe the bottom sediment that hosts the ferromanganese nodules and crusts and go on, in Chapter 5, to explain the processes of the formation of nodules and crusts in the light of variable source material, tectonic activity and mid-plate volcanism. The mining, environment, metallurgy, legal and economic aspects of the IONF resources are described in Chapter 6. Throughout, recent concepts, hypotheses and critical appreciation of the state-of-the-art knowledge on nodule formation and resource management are incorporated. The wealth of information that the authors have brought together here marks out this volume as the definitive work on the IONF. Martin Hale The Netherlands March 2007

ABOUT THE AUTHORS

Ranadhir Mukhopadhyay, 49, was the Director of the Mauritius Oceanography Institute (2002–2004) and currently is a senior scientist with the National Institute of Oceanography, Goa. He is the recipient of the Asiatic Society Medal, Raman Research Fellowship and M. S. Krishnan Gold Medal. He has co-authored a book and has about 40 research papers and articles to his credit. He has worked on plate boundary geomorphology, seamount-tectonics and marine minerals. Anil K. Ghosh, 67, is an UGC Emeritus Fellow with the University of Calcutta and received several awards including the Universtiy Gold Medal, Asiatic Society Medal and Coggin Brown Gold Medal. He specialises in mineral resources and has been teaching at the postgraduate level for more than four decades in two faculties (Geology and Marine Sciences). He has written about 40 research papers, co-authored a book on marine mineral resources and edited another on land-based resources. Sridhar D. Iyer, 48, is a keen petrologist, and is currently a senior scientist with the National Institute of Oceanography, Goa. He is the recipient of Young Scientist award and Raman Research Fellowship. He has worked extensively on volcanics, occurring both on the younger and on older oceanic crusts of the Indian Ocean and has authored about 60 research papers and articles.

xiii

C H A P T E R

O N E

The Indian Ocean Nodule Field

Contents 5 6 8 11 13 17 22 24 25 25 27 28 31 32 35

1. Physiography and Geology 1.1. India–Australia plate boundary 1.2. Australia–Capricorn plate boundary 1.3. The Chagos–Laccadive Ridge 1.4. The Ninetyeast Ridge 1.5. The Carlsberg Ridge 1.6. The Central Indian Ridge 1.7. The Southeast Indian Ridge 2. Physics, Chemistry and Biology 2.1. Physical characteristics 2.2. Chemical characteristics 2.3. Biological characteristics 3. Evolution of the Indian Ocean Nodule Field 3.1. Break-Up of Gondwanaland 3.2. Formation of the Indian Ocean Ridge system

The Indian Ocean constitutes about one-seventh of the earth’s surface and is the world’s third largest water body. This ocean covers an area of 73.6 million km2, and is separated from the Atlantic and the Pacific oceans by roughly 20 E and 147 E, respectively (Fig. 1.1). This ocean has no extension towards the North Pole unlike the Atlantic and the Pacific oceans. The detachment from the Arctic polar water system and the presence of mountainous Asian landmass are responsible for several interesting phenomena in this ocean, such as atypical circulation of water mass, extensive upwelling of nutrients along the east African coast and the formation of summer and winter monsoon in Asia, among others. Scientific exploration in the Indian Ocean started long after the same was initiated in the Pacific and the Atlantic oceans. HMS Challenger from Great Britain was the first ship to start research in this ocean in a big way during its famous transworld voyages between 1872 and 1876. The John Murray Expedition, onboard Mabahiss in 1933–1934, was the next major one. A few other oceanographic expeditions, namely, Dana (1928–1930), Snellius (1929–1930), Albatross (1950– 1952) and Ob (1955–1957), were also made with specific objectives. Later, two major events paved way for a full-fledged scientific research in this ocean. These were (1) International Geophysical Year (IGY, 1957–1958) and (2) International Indian Ocean Expedition (IIOE, 1963–1966). Handbook of Exploration and Environmental Geochemistry, Volume 10 ISSN 1874-2734, DOI: 10.1016/S1874-2734(07)10001-2

#

2008 Elsevier B.V. All rights reserved.

1

2

Mukhopadhyay, Ghosh and Iyer

20⬚E

40⬚

60⬚

80⬚

100⬚ 20⬚ N

0⬚

20⬚ S

40⬚

Figure 1.1 General physiography, essential tectonic elements and bathymetry of the Indian Ocean and neighbouring seas (base map from NGDC). The Indian Ocean Nodule Field (IONF) within the Central Indian Ocean Basin (CIOB) is shown in square box. Note inverted Y-shaped profusely fractured mid-ocean ridge system, the long north-south-trending Chagos^Laccadive Ridge and Ninetyeast Ridge on either side of the IONF, and the Andaman^Sumatra subduction zone to the east.

The magnitude of IIOE, in terms of involvement of several countries, made it the greatest oceanographic endeavour to date. There were in all 36 ships from 13 countries, which took part in the expedition lasting for 4 years. After IIOE, the next major work was the drilling of the oceanic crust by Glomar Challenger, under the Deep Sea Drilling Project (DSDP, later known as the Ocean Drilling Programme, ODP). The drilling programme was further complemented by the Joint Oceanographic Institutions for Deep Earth Sampling ( JOIDES) and, in recent times, by Integrated Ocean Drilling Programme (IODP). In the 1980s, the studies of the Carlsberg and Central Indian ridges started in an organised manner. The German vessel FS Sonne deciphered two sites of significant hydrothermal activity along this ridge under an Indo–German collaborative programme, Gemino. The French, American, British and Indian efforts are presently concentrated along these ridges in search of hydrothermal sulphide mineralisation and to understand the interactions among asthenosphere, lithosphere, hydrosphere, biosphere and atmosphere. Some of the important characteristics of oceanic crust and mantle are furnished in Table 1.1. Among the various basins in the Indian Ocean (Fig. 1.1), the Central Indian Ocean Basin (CIOB) was particularly studied for its mineral potential. During the last two decades, the scientific community explored the CIOB both extensively and intensively. Since 1982, more than 50 cruises have been undertaken in this basin, which ultimately led to the identification of a manganese nodule-bearing area in the CIOB. Such resource is also known as polymetallic nodules and ferromanganese

3


Table 1.1 Physical characteristics of oceanic crust and mantle

Layer

Lithology

1 2A

Sediments Basaltic sheet and pillow lavas Basaltic dykes Gabbros Layered gabbros Peridotites

2B 3A 3B 4

P-wave velocity

Thickness (km)

Avg. density (g/cm3)

1.7–2.0 2.0–4.1 (3.6)

50 mWm2 above regional background and non-linear thermal gradient suggest circulation of hydrothermal fluid along fault planes in the deformation zone. The high heat flow (100–200 mWm2) and three large free-air negative gravity anomalies at 4 S/80 E (40 mGal), at 0 450 N/ 82 550 E (30 mGal) and at 1 N/83 450 E (70 mGal), with corresponding geoidal

8


undulations, favour a uniform crustal folding or buckling within the Moho (cf. Neprochnov et al., 1998). The gravity model study suggests that the average thickness of the crust in the IAPB is about 5 km, increasing up to 8 km, particularly towards the east at around 80 E due to interaction with intensely folded crustal blocks (elastic thickness, Te ¼ 12–15 km). The deformation in the IAPB is a subject of debate. This deformation appears to have influenced drainage and depositional dynamics in the CIOB in conjunction with Himalayan tectonics. The large contribution of clastic materials (including heavy minerals) from the Himalayas is noted in the sediment that tops the crust in the IAPB deformed zone. Such deposition appears to have occurred in two periods— first between 11 and 7.5 Ma (Late Miocene) and the second around 0.9 Ma (Late Pleistocene). These two phases of sedimentation were interrupted by sediment contributions from the Lesser Himalayas and the Indian Subcontinent. It hints at two primary phases of Himalayan uplift during Late Miocene and Pleistocene. The intensification of the Indian monsoon at about 8 Ma resulted in a rapid denudation of clastic sediments. New seismic reflection data reveal (Fig. 1.5) that the first deformation in the IAPB may have occurred during Late Miocene (7.5–8.0 Ma, well evident in the southern part of the IAPB, south of 7 S), followed by another during Early Pliocene (4–5 Ma, well developed in the northern part, north of 7 S), and the last deformation during Late Pleistocene (0.8 Ma, overlapping the Miocene and Pliocene deformed areas). The sedimentation history also supports this seismic finding. For example, the syn-deformation sediments are separated from predeformation sediments by an unconformity dating 7 Ma (Late Miocene), probably indicating the onset of plate boundary deformation. The second unconformity dates to Upper Pleistocene (800 ka) and separates glacial from non-glacial strata.

1.2. Australia–Capricorn plate boundary The data on seismicity and plate motion beyond the southern borders of the IONF are inconsistent, keeping in view the accepted plate tectonic theory that plates are rigid, least deformed in their interior and should have narrow boundaries. However, Two-way travel time (S)

0

10

20

30

40

50

60

70

80

90

100 110 km

Upper Upper miocene u/c pleistocene u/c

6.4 6.6 6.8 7.0 7.2 7.4 7.6

6.4 6.6 6.8 7.0 7.2 7.4 7.6

Deformed sediment Basement rise 04⬚34.86⬘S 86⬚52.47⬘E

Faulted basement

AS10-05

03⬚31.38⬘S 86⬚53.85⬘E

Figure 1.5 Crumpled oceanic basement and sediment in the form of tight folds and high-angle faults located in the Indian^Australian deformation zone (IAPB), just north of the Indian Nodule Ocean Field (IONF).These were formed during various geological periods as marked by unconformities (Krishna et al.,1998).

9


several discrepancies to this concept were encountered in the Australian Plate (southeast of 16 300 S) in terms of (1) uncharacteristic NW-SE shortening, (2) change in the orientation of axis of shortening from NS to NW-SE from the west to east of 86 E, (3) occurrence of an anomalous plate boundary type, for example, off-axis normal faulting earthquakes near 80 E on the Southeast Indian Ridge (SEIR) and (4) reactivation of extinct fracture zones in the Wharton Basin. The reason for such inconsistencies was investigated on an 11-Ma-old seafloor flanking the South West Indian Ridge (SWIR) and SEIR (¼ anomaly 5, Royer and Gordon, 1997), to see whether such a discrepancy could be justified by any deformation within the Australian Plate. The findings suggested that these discrepancies could be suitably rationalised if one recognises the occurrence of a new Capricorn Plate (Figs. 1.6, 1.7). The N40E-oriented gravity undulations in the ACPB(200-km wavelength and 10-mGal amplitude) differ in trend from the IAPB by about 40 and in intensity by a factor ranging between 2 and 3. Incidentally, undulations in the ACPB are oriented along the absolute motion direction of the Indian and the Australian plates in the hotspot reference frame, and occur largely on the fast spreading Australian Plate and not on the slow spreading Antarctica Plate. The plate reconstruction, based on magnetic anomaly data and fracture zone across the SEIR, suggests convergence

69⬚E

Gravity (mgal)

68⬚E

50

70⬚E

71⬚E

72⬚E

73⬚E

74⬚E

CIR 5⬚S

0 −50 6⬚S

CB −100

20⬚N

IND

7⬚S DG

0⬚

8⬚S 20⬚S

CAP

80⬚E

M = 5.0 AUS

100⬚E

Figure 1.6 Plate configuration in the Central Indian Ocean Basin (CIOB) from satellite-derived gravity field (Sandwell and Smith, 1997). IND ¼ Indian, CAP ¼ Capricorn, AUS ¼ Australian plates, CB ¼ Chagos Bank, DG ¼ Diego Garcia, CIR ¼Central Indian Ridge. Earthquake epicentre data (courtesy: ANSS) in open and black circles in the main and inset figures. Gray in inset represents plate boundary zones, while horizontal lines represent extension of India^Capricorn plates (Henstock and Minshull, 2004).

10


60⬚

EURASIA 50⬚

40⬚ 8 30⬚ Bhuj

7

ARABIA

20⬚

6 INDIA

10⬚

5 0⬚ 4 SOMALIA

−10⬚

CAPRICORN 3

−20⬚

M

−30⬚

AUSTRALIA km 0

ANTARCTICA

−40⬚ 40⬚

50⬚

60⬚

70⬚

80⬚

90⬚

20 mm/yr

1000 100⬚

110⬚

120⬚

Figure 1.7 Seismicity recorded between the years 1900 and 2001 on the Indian, Capricorn and Australian plates shown in circles with radius proportional to intensity. Arrows show direction of spreading (Stein et al., 2002).

of 23 2.6 km since 11 Ma between the Capricorn Plate at 10.3 S, 83.5 E and the Australian Plate at 17 S, 105 E. This corresponds to a rate of 2.4 2.1 mm/year. This rate of convergence is however much slower than the global average rate of 70 mm/year, and even less than the slowest one known so far (20 mm/year, DeMets et al., 1990). The reconstruction also indicates that since 11 Ma, a point now at 17 S, 105 E on the Australian Plate has moved 27 km along N45W relative to the Capricorn Plate. Large earthquakes reflect the convergence and large-scale folding of the lithosphere, represented by undulations in northeast striking gravity, occurring on both sides of the NER (Royer and Gordon, 1997). The reconstruction also indicates a divergence of 13 2.4 km since 11 Ma between the Capricorn Plate at a point 26 S, 74 E and the Australian Plate at 41 S, 90 E. The convergence corresponds to a rate of 1.2 2.2 mm/year, which is much slower than global values (average 40 mm/year, least 10 mm/year; Chu and Gordon, 1998). The average rate of rotation during the last 11 Ma between the Indian and the Capricorn plates has been estimated as 0.24 0.02 per million years,


11

while that between the Capricorn and the Australian plates has been 0.07 0.03 per million years (Royer and Gordon, 1997). A recent study suggests extension of 10–12 km of crust over a zone of about 50 km in Chagos area as evidenced by crustal thinning and concentration of seismicity. The localisation of seismic activities beneath Chagos Bank is due to a weak rheology, probably caused by thick crust. The area of deformation related to the composite Indian, Capricorn and Australian plates (IAPB þ ACPB) exceeds the size of several individual plates, and may have influenced the development of an increased number of closely spaced transform faults across the northern CIR.

1.3. The Chagos–Laccadive Ridge Aseismic ridges are a group of non-spreading, topographically elevated chained features on the seafloor formed either by volcanic eruption along a leaky transform fault, or hotspot eruption, or through uplift and rifting associated with tectonic activities. The hotspots are long lived and relatively fixed, mantle thermal manifestations that produce magma chemically distinct from that erupting at mid-ocean ridges (Table 1.2). The hotspot-generated aseismic ridges of the Indian Ocean suggest that rapid, northward migration of the Indian Plate over several hotspots during Cretaceous and Early Cenozoic times produced the majority of the chained volcanic traces (Fig. 1.8). Of these, the two aseismic ridges, which act as western and eastern borders of the IONF, are (1) the CLR (a part of Reunion Island–Mauritius Islands– Southern Mascarene Plateau–Chagos Ridge–Maldives Ridge–Laccadive Ridge– Deccan Traps chain believed to have been formed by the eruption of Reunion hotspot) and (2) the NER(a part of the Kerguelen Islands–Ninetyeast Ridge–Broken Ridge–Rajmahal Traps chain probably caused by Kerguelen mantle plume). Extending from 14 N to 9 S on the Indian and the Australian plates, the CLR is an 2550-km long and 200-km wide aseismic volcano-topographic feature created by the eruption of the Reunion hotspot approximately along 73 E meridian (Fig. 1.8). This ridge, slightly convexed towards the east, forms the western boundary of the IONF. A considerable length of the crest of the CLR is composed of shoals, banks, coral reefs and atolls at depths 44my

94⬚E FZ

13

31?

92⬚E FZ

33

253

32B 33

25

21?

29?

28 29

28

24B

?

10⬚

26

26

25

23B 23? 757 B

26 25

23

20 21? 23A

27? 27?

25

23A 22 20

25

22

96⬚E FZ

30

24

21

88⬚E FZ

215

31 29

24

5⬚

22

32

30

25

22

21

28 31

E

N

32

33

ON

>65my

31

25

29

216

TI

84⬚E FZ 84.5⬚E FZ

30

32

34

32

5⬚ N

BD

83⬚E (Indira) FZ

31

>80my

SU

34

DA

34

32

758 33

SU N

34?

SRI LANKA

25⬚ S

756 >38my

85⬚E

90⬚

95⬚

Figure 1.10 Tectonic disposition of the Ninetyeast Ridge (closed north-south contour) with a decrease in age from north to south. Location of ODP sites (756, 757, 758) and DSDP sites (214, 215, 216, 253) are shown by small solid dots. Also shown magnetic lineation, abandoned spreading centres (solid rectangles) and fracture zones (dashed lines) (Krishna et al.,1995).

16


0 A

50 km

2

4

6 S

2.5

3.0 S B TWT

5 km

88⬚E

2

4 S

0

0

50 km

2

N

S

3.0

3.5 S

4 TWT

TWT

88⬚E

6

0

5 km

8

TWT = Two-way travel time (s)

Figure 1.11 Seismic reflection sections across the Ninetyeast Ridge. Note the thick pelagic sediment cover at the crest, more than 400 m at places (cf. Ramana et al., 2001).

resulted from hotspot volcanism over a spreading axis and the portion south of 17 S was formed again on the edge of the Indian Plate (Fig. 1.11). Emplacement of the NER, particularly between 14 S and 17 S, has been unique as during the period 60 and 54 Ma excess crust was generated because of the location of the Kerguelen hotspot below the Wharton Spreading Ridge. This resulted in thickening of the crust (20 1.3 km, Krishna and Rao, 2000; Sinha et al., 1981). The magnetic and seismic conducted studies in the vicinity of the NER reveal the presence of a chain of abandoned spreading centres (ASC) of different ages. These ASC ceased spreading at about 42 Ma and are interpreted to be the western extension of the east-west-trending Wharton Spreading Ridge. This ridge jumped southward between 65 and 42 Ma, and in the process a considerable amount of the oceanic crust formed originally on the Antarctic Plate between anomalies 30 and 32 n.2 was transferred to the Indian Plate (Fig. 1.12; Krishna and Rao, 2000). Petrologically, the NER is composed of tholeiitic basalt, ranging from low MgO and aphyric olivine in the south to strongly plagioclase-phyric, high MgO towards the north. The rocks are extensively altered (H2Oþ > 1%) in a low-temperature environment under both reducing and oxidising conditions (ODP 115, DSDP legs 22, 26; Weis et al., 1992). Major element and isotopic composition of basalts along

17

86⬚E FZ

84.5⬚E FZ

217 (85 Ma)

A30-A34

4 A3 0-

A3

Spreading center ceased at 65 Ma

A28-A34 A25-A34 A29-A34

A

30

-A

3

. 2n

2

Portion of ANT plate transferred to IND plate

A3

4

A3

756 (82 Ma)

216 (78 Ma)

A3 34 0-A

4

A3

0-

A3

Ridge jump between A32n.2-A33

Ninetyeast ridge

A3

216 (78 Ma)

0-

A3 4

756 (82 Ma)

Indian plate

Indira FZ

A3 4

217 (85 Ma)

0-

Indira FZ

84.5⬚E FZ

Indian plate

Ninetyeast ridge

86⬚E FZ


4 A3 0A3

A3

0-A

34

A3

0-

215 (61 Ma) A26-A29

Kerguelen hotspot

Antarctica plate

Indian plate

65 Ma (after anomaly 30)

Present (only north of 108S)

Figure 1.12 Left panel shows a spreading ridge (slanting brick) separating the Antarctic Plate (dotted) from the Indian Plate (small squares). The spreading ridge jumped towards south between chron A32 (72 Ma) and A33 (78 Ma). In the process, it transferred (right panel) the oceanic crust that accreted originally at the Antarctic (ANT) Plate between anomalies 30 and 32 n.2 to the Indian (IND) Plate (Krishna et al.,1998).

the NER suggest the presence of two compositionally distinct parental magmas. The compositional variations in the rocks along the NER may be due to fractional crystallisation and variable extent of melting and magma mixing. It appears that the lavas for both CLR and NER erupted close to the spreading axis, consistent with the plate tectonic reconstruction model of the eastern and the western Indian Ocean basins (Bhattacharya and Chaubey, 2001; Ramana et al., 2001). As both the Reunion and Kerguelen hotspots were crossed by one spreading axis each (CIR and Wharton Ridge, respectively), the mixing of the plume with ridge-axis melt (N-MORB) depletes the melt both in trace element and in isotopic components. Early products of the Reunion and the Kerguelen hotspots have more depleted isotopic signatures than the later products, reflecting an increase in the mixing of hotspot and ridge-axis melts in the past.

1.5. The Carlsberg Ridge The MOR system in the Indian Ocean (known as the IORS) differs from that in the Pacific Ocean in terms of rate of spreading (Table 1.3), element chemistry (Table 1.4) and isotopic composition (Table 1.5). The entire MOR system forms a 70,000-km long serpentinous mountain chain and rises to about 2–3 km above the surrounding seafloor. This submarine chain encompasses about 33% of the total area of the ocean floor and covers 35.9% of the Pacific, 31.2% of the Atlantic and 30.2%

18


Table 1.3

Characteristics of accreting mid-ocean ridges

Slow-spreading ridge

Fast-spreading ridge

Spreading rate < 40 mm/year (full rate).

Spreading rate 80–160 mm/year (full rate). Absence of accumulated stress and hence very less prone to earthquakes. Smooth seafloor morphology.

Presence of deep-seated earthquakes and major normal faults. Rough to very rough seafloor morphology. Median valley present.

A rise develops because of absence or poorly formed median valley. Symmetrical and elevated volcanic edifices are present. Magmatic activity dominates over tectonic processes.

Highly segmented and asymmetric rifted depressions are conspicuous. Nature and scale of segmentation is predominantly controlled by tectonic activity. Neo-volcanic zone is wider (2–12 km) and Neo-volcanic zone is narrow has small seamounts. (100–200 m) and seamounts are virtually absent. Near-absence of off-axis seamounts. Frequent presence of off-axis seamounts. Magmatic discontinuities and unfocussed Magmatic continuity and relatively magmatism. focussed magmatism. A low rate of magma outpouring prevails. The magma outflow rate is higher. Volcanic eruptions are larger. Volcanic eruptions are smaller. Dykes are less but if present are large in Dykes are common but are of smaller dimension. dimension. Pillow lavas abundant. Sheet flow dominates. A narrow range of relatively A wider range of generally more undifferentiated lavas is produced. differentiated lavas is produced. The compositional variations are The lavas show more complex chemical dominated by relatively simple lowtrends ascribed to polybaric fractional pressure fractional crystallisation crystallisation and/or phenocryst trends. reaction associated with widespread accumulation of calcic plagioclase. Small and/or intermittent, long-lived and Large, short-lived and steady-state non-steady-state magma chambers. magma chambers. Magma mixing is most evident. Crystal fractionation largely overrules magma mixing. Mafic and ultramafic rocks commonly Rare occurrence of mantle rocks. occur. Sources: Iyer and Ray (2003) and references therein.

of the Indian Ocean floors. Along with the Central Indian and Southeast Indian ridges, the Carlsberg Ridge (CR) forms the northwestern boundary of the IONF and appears to have considerable bearing on the evolution of the mineral deposits within the field.

19


Table 1.4

Major and trace element chemistry of the three oceanic ridges IORS

MAR

EPR

Oxides SiO2 TiO2 Al2O3 FeOt MgO CaO Na2O K2O P2O5

50.93 1.19 15.15 10.32 7.69 11.84 2.32 0.14 0.10

50.68 1.49 15.60 9.85 7.69 11.44 2.66 0.17 0.12

50.19 1.77 14.86 11.33 7.10 11.44 2.66 0.16 0.14

Trace elements Rb Sr Ni Cr Co Cu Y Zr Nb V

2.54 141 106 320 42 81 35 112 4.0 243

2.24 116 121 300 49 77 25 46 1.6 289

1.4 120 100 318 47 78 39 88 2.0 318

Sources: Melson et al. (1976), Banerjee and Iyer (1991), Subbarao et al. (1977) and Sun et al. (1979). Note: IORS ¼ Indian Ocean Ridge system, MAR ¼ Mid-Atlantic Ridge, EPR ¼ East Pacific Rise. Oxides in wt%, elements in ppm, t ¼ total.

The CR runs northwest-southeast across the middle of the Arabian Sea between the Owen Fracture Zone in the north and the equator in the south. The CR accounts for the generation of new oceanic crust between Iran–Afghanistan in the north and Madagascar in the south, separating, as a result, the Indian Plate from the African Plate. Along the Owen Fracture Zone, the CR is offset right laterally by about 300 km against the northern Sheeba Ridge. Geophysical investigations have shown that the CR was possibly formed during Late Paleocene (63–61 Ma) by a WNW-trending spreading centre. It displays a rough topography, typical of slowspreading ridge, having half-spreading rate of 13–15 mm/year since 30 Ma. The ridge is displaced by major transform faults particularly in the southern part. The gravity data suggest the median valley of the CR to presently be in a nonisostatic equilibrium. Supporting magnetic, bathymetric and gravity evidences indicate that prior to the Himalayan collision, the CR was spreading at a faster rate and that during the hard collision of India with Eurasia in the Eocene, the spreading rate slowed down considerably. These phenomena, represented by two asymmetric, discordant systems of linear magnetic anomalies, attest to the fact that the structure and spreading of the CR was non-stationary. In contrast, the mantle Bouguer

20


Table 1.5

87

Isotopic composition of basalts from the three spreading ridge axes

Sr/86Sr

206

Pb/204Pb

207

Pb/204Pb

208

Pb/204Pb

143

Nd/144Nd

IORS

MAR

EPR

0.7032–0.7035 0.70261–0.70278 0.70291–0.70329 0.70273–0.70356 0.70257–0.70289 18.43–18.75 17.307–18.019 17.898–18.839 18.000–18.157 15.43–15.49 15.505–15.563 15.437–15.604 15.443–15.472 38.01–38.19 37.214–37.838 37.820–39.077 37.772–38.001 0.51302–0.51305 0.51304–0.51309 0.51293–0.51316 0.51305–0.51309

0.70230–0.70321 0.70215–0.70287

0.70253–0.70275 0.7023–0.7028

17.84–19.28 17.829–19.437

18.43–18.56 17.721–18.726 18.0–18.8

15.46–15.59 15.439–15.586

15.44–15.49 15.309–15.514 15.3–15.6

37.33–38.83 37.418–38.913

37.64–38.07 37.029–38.805 37.0–38.8

0.51309–0.51329

0.51302–0.51313 0.51306–0.51318

Sources: Subbarao and Hedge (1973), Cohen et al. (1980), Price et al. (1986), Mahoney et al. (1989), Rehkamper and Hofmann (1997), Hamlein et al. (1984). Note: IORS ¼ Indian Ocean Ridge System, MAR ¼ Mid-Atlantic Ridge, EPR ¼ East Pacific Rise.

anomaly (MBA) values suggest a more continuous magmatic source beneath this chaotic morphology. The along-axis variations in MBA values and the swath bathymetry map of a part of the CR (Fig. 1.13) reveal differences in magmatic and tectonic processes. Tiny wiggles, that is, second-order magnetic signals, reveal two eastwardpropagating rifts between anomalies 26 and 25, and seven westward-propagating rifts between anomalies 24 and 20, leading to a regional assymetricity (Dyment, 1998). For example, about 56% of the crust formed between anomalies 26 and 25 was accreted to the African Plate, while during anomalies 24 to 20, more than 75% of the generated crust was added to the Indian Plate. The backscatter amplitude data confirm the presence of a 2200–4500-m deep rift valley filled with sediment, and of ultramafics such as lherzolite and serpentinite, indicating the shallow depth of mantle melting in variable thermo-tectonic regimes (Mudholkar et al., 2002). Broken rock fragments of neighbouring continents are also found in many places along the CR. Well-preserved pteropods in the proximity of the CR may suggest the presence of hydrothermal activities (Bhattacharya, 1996). The whole rock and mineral chemistry of the CR suggest a greater influence of tectonic setting (transform fault effect) on the formation of crust through fractional

21


A

9⬚N

6⬚

3⬚

0⬚

−160

−100

−120

58⬚E

−50

−60

−20 −40

60⬚

10 0

62⬚

40 20

80 60

64⬚

110

3⬚S 66⬚

68⬚E

B 1.4 1.3

S, west flank (Cm/y)

1.2 1.4

S, east flank

2

1.3 1.2 1.4

1

3

S, mean

1.3 1.2

Figure 1.13 (A) Mantle Bouguergravity anomaly map of apart of the Carlsberg Ridge (CR) based on satellite altimetry data (Sandwell and Smith, 1997).The NW-SE-trending thick black line is the ridge axis. (B) Spreading rate distribution for anomalyA5 along the CR for the west (top), east (middle) flanks and average spreading rate (bottom).The solid dots are rates as measured, thin lines are running average, while the thick lines show theoretical rate calculated using anomaly A5 and CR pole finite rotation (Merkourievand Sotchevanova, 2003).

crystallisation and magma mixing (Iyer and Banerjee, 1993). The presence of plagioclase phenocrysts with anorthite-rich core and corroded margin, coexistence of calcic plagioclase (An86–89) and olivine (Fo83–91), presence of twinned sodic plagioclase and/or reverse zoned plagioclase, low CaO/Al2O3 ratio, absence of clinopyroxene, hyperbolic relation of the ratio–ratio plots of trace elements and variable range of TiO2 at a given FeO/MgO in the bulk rocks, when considered together, suggest short interval of crystallisation, differential extent of partial melting at shallow depth and magma mixing.

22


1.6. The Central Indian Ridge The CIR with a trend of NNW-SSE and a length of about 2775 km separates the Indian, Australian and Capricorn plates from the African Plate. The CIR joins the CR in the north around the equator and the Indian Ocean Triple Junction (IOTJ) at 25 S. The ridge shows a linear progressive increase in spreading rate from 34 mm/year (spreading direction N142E) at the equator to 55 mm/year (spreading direction N152E) at the IOTJ. It is characterised by en echelon displacement by numerous NE-SW-trending fracture zones and transform faults. The southern part of the ridge differs in topography, structure and age from the northern part. Magnetic and multibeam bathymetric data from the intersection between the Vema Fracture Zone and the CIR reveal asymmetric spreading up to chron A5 (10 Ma; Drolia et al., 2000) and variable along-axis crustal thickness (Kamesh Raju et al., 1997). The deepest point (>6200 m) over the CIR is encountered along the Vema Fracture Zone (also known as Vema Trench). Multibeam swath bathymetry profile along and across the trench is shown in Fig. 1.14. Except for the offsets along the Alula-Fartak FZ and the fracture zone at 3 S, all other offsets of the CIR are left lateral. The length and nature of these offsets appear to have influenced the variable extent of melting of the source rock and the depth of magma generation (Mukhopadhyay and Iyer, 1993). The comparatively slow-spreading northern CIR and the intermediatespreading southern CIR meet at the Wide Boundary Zone (WBZ, i.e. the extended part of the deformation zone occurring at the boundary of the Indian and the Australian plates, Fig. 1.5) on the ridge axis between 6 S and 9 S, suggesting initiation of a new triple junction among African, Indian and Australian plates (DeMets et al., 1994; Drolia et al., 2000). Recent gravity model indicates that the northern CIR in the equatorial region has a low elastic thickness of about 5–10 km (average 7 2 km) and a crustal thickness of 18 km. The former may have been caused by the higher than average temperatures beneath this portion of the ridge. This in turn may suggest probable perturbation of the thermal structure by the effect of the nascent triple junction, which may be the contact of the CIR with the WBZ (Drolia et al., 2003). Rao et al. (1996) proposed different ages of evolution for the northern and the southern parts of the CIR at about 30 Ma (anomaly 10) and 68–66 Ma (anomaly 28), respectively. The presence of a transform fault at 7 450 S displacing the CIR axis in a right lateral sense has also been reported (Kamesh Raju et al., 1997). Recent detailed multidisciplinary studies carried out under the InRidge programme from the northern CIR revealed variable and atypical petro-tectonic characteristics at the transform fault (fracture zone), ridge axis, ridge-transform fault intersection and at near-axis seamount (Mukhopadhyay et al., 1998). The presence of pillow and columnar basalts in a neo-volcanic zone at some transform faults suggests a complex relation between magmatism and tectonics (Drolia et al., 2003). At 19 S, the interaction between mantle plumes (Reunion–Rodrigues–CIR hotspot track) and spreading centres (CIR) is manifested in terms of significant variations in axial topography, non-transform offset morphology, segmentation characteristics, crustal thickness and basalt chemistry. Rocks along the CIR are

23


NW 4

Depth in m

−4500

Distance in km 6 8 10 12

SE 14

Across transform

−5000

PTDZ

−5500 −6000

−9.00⬚ −4000 −4500 −5000 −5500

Depth in m

Along transform axis

−6000

−9.50⬚ 67.00⬚ 2000

67.50⬚

3000

4000

5000

50 6000

NW

100 150 200 Distance in km SE

Depth in m −7.00⬚

−7.50⬚

−8.00⬚

−8.50⬚ 67.50⬚

68.00⬚

2000 3000

4000

68.50⬚ 5000 6000

Depth in m

Figure 1.14 Fifty-nine beam swath bathymetry mosaic across and along theVema Fracture Zone cutting across the Central Indian Ridge (Drolia and DeMets, 2005). Note the deepest point (>6200 m) along the Vema Fracture Zone (arrow) (Courtesy: InRidge, project CLP 0886, cruise SK165).

24


mostly normal Mid-Oceanic Ridge Basalt (N-MORB) type, but transitional and enriched basalts (T/E-MORB) are also found at places. In addition, ultramafics and rocks containing hydrothermal indications also occur (Mukhopadhyay and Iyer, 1993). Rocks along the CR–CIR stretch indicate that the magma is very primitive and less differentiated than that of the Mid-Atlantic Ridge (MAR) and the East Pacific Rise (EPR). The melt beneath the IORS is isotopically distinct from those in the Atlantic (MAR) and the Pacific (EPR) oceans. For example, the IORS is characterised by a higher 87Sr/86Sr and a lower 206Pb/204Pb and 143Nd/144Nd than the EPR and MAR (Table 1.5). This suggests either mixing of Kerguelen, Marion, Reunion and Crozet mantle plumes with the ridge-axis melt (Dupal anomaly) or mixing of earlier consumed continental crust with ridge-axis melt (Iyer and Ray, 2003; Mahoney et al., 1989; Rehkamper and Hofmann, 1997; Subbarao and Hedge, 1973). Along the 4200-km long CR–CIR stretch, variations in mineralogy, texture, trace element geochemistry and isotope chemistry of rocks in response to ridge segmentations were observed, and seem to have played a significant role in facilitating polybaric fractional crystallisation.

1.7. The Southeast Indian Ridge The SEIR separates the Antarctic and the Australian plates and connects the IORS with the EPR at the Macquarie triple junction (Fig. 1.1). Except for a much younger part between Broken and Kerguelen–Gaussberg ridges (age 44 Ma), the other stretches of the SEIR hold a spreading history since 95 Ma. The SEIR has a spreading rate varying from 58 mm/year in the west near IOTJ to 76 mm/year in the east around 120 E (Fig. 1.3). The SEIR is characterised by remarkable variations in ridge-crest morphology, suggestive of excess magma supply caused by hotspots. This ridge is divided into two super-segments, one of which extends from 70 E (IOTJ) to 90 E and is highly segmented, bordered by several transform faults and influenced along part of its length by the Amsterdam/St. Paul and Kerguelen hotspots (Klein et al., 1991). Some of the prominent fracture zones extend up to the Bay of Bengal. The other super-segment extends from 90 E to the Tasman Sea in the east and includes Australia–Antarctica Discordance (AAD). The satellite-derived gravity and limited bathymetry data suggest that the SEIR exhibits transition in axial morphology, separated by two transform faults. For example, the SEIR west of the 102 450 E transform fault is characterised by axial rift volcanic ridge (EPR type) of 400-m height and 10-km width, while on the east of 114 E transform fault, the SEIR shows a rift valley of about 1-km depth and 20-km width. The SEIR between these two faults is represented by an axial valley of moderate depth and width (600 m, 10 km). The absence of any large variation in mantle temperature (25 C–50 C) is manifested by near-uniform crustal thickness below the SEIR, and this may be explained by the presence of a steady-state axial magma chamber and a nearly constant spreading rate (Shah and Sempere, 1998). However, the transition in axial morphology appears to coincide with variations in the geochemistry of axial lavas (Sempere and Klein, 1995).


25

The AAD, located between 119 E and 126 E, is an anomalously deep portion and marks a boundary between the isotopic provinces of the Indian and the Pacific oceans. The rocks from AAD are highly enriched with incompatible elements as well as with Si, Fe and Na compared to N-MORB rocks on its either side. The lowest solidus pressure and least extent of melting, supported by the bathymetric, gravity and seismic evidences from beneath the AAD, suggest the presence of a cooler temperature in the mantle (Klein et al., 1991). It is well known that the Indian Ocean MORB has isotope systematics that are different from those of the Atlantic and the Pacific MORB (Table 1.5). Also, the SEIR basalts are different from those occurring at the IOTJ, and their source may involve variable proportion of a component tentatively assigned to recycled ancient hydrothermal and abyssal sediments. The Sr, Pb and Nd isotope data suggest that the SEIR basalts may also be accounted for by binary mixing of an Indian MORB with ocean-island basalt melt. Hence, the SEIR is an ideal site to study the effects of varying mantle temperatures on crustal accretion, the thermo-mechanical structure and the magmatic interaction between MORB and mantle plumes (Amsterdam/St. Paul and Kerguelen hotspots).

2. Physics, Chemistry and Biology The physical, chemical and biological characteristics of the sediment and water column in the IONF appear to have influenced the formation and growth of its mineral resources. A gist of these parameters is outlined below.

2.1. Physical characteristics Much of the physical oceanographic data from the IONF were systematically collected and analysed under the Environmental Impact Assessment (INDEX) programme of India’s deep-sea mineral venture (see Chapter 6 for details). It has been found that the physical parameters (temperature, salinity and potential density) vary down the water column in the IONF, but the variations are largely limited to the top 3500 m, and suggest a restricted basin-scale deep circulation. The topography at 5000 m in the central part of the IONF displays the abyssal circulation and is generally characterised by a southwestward weak flow around 10 S. This flow regime is flanked by cyclonic and anti-cyclonic eddies to its right and left, respectively (Ramesh Babu et al., 2001), and could be linked to the entry of Antarctic bottom water into the CIOB through a saddle across the NER at 10 S (Warren, 1982). The annual mean temperature and salinity distribution at surface, and at depths of 100 and 1000 m in the IONF, have also been studied (Fig. 1.15). Geostrophic current circulation patterns were estimated at a few locations in the IONF. At 79 E, an anticlockwise cell is clearly discernible in the upper 300 m, with moderately strong westerly flow (0.7 m/sec) in sectors A and B and in the northern half of sector C. In sector D and the southern part of sector C, the flow becomes weak (0.3 m/sec), easterly and clockwise. At 71 E, the circulation cell in the upper

26


34.60

34.80 S (psu)

34.70

0.0 0

1.0

2.0

3.0

PO4-P (mM)

1000

Depth (m)

2000 Salinity

3000 Phosphate

DO 4000

Nitrate

pH 5000 0 0 7.6

100 10

20 7.8

300 DO (mM)

200 30 8.0

40 NO3-N (mM) 8.2 pH

Figure 1.15 Vertical profile in the water column in the Central Indian Ocean Basin (CIOB) showing annual mean salinity, pH, dissolved oxygen, nitrate and phosphate (De Sousa et al., 2001).

300 m shows clockwise rotation and a weak easterly flow (0.1 m/sec) in sectors A and B and in the northern part of sector C. This trend gets reversed in the south, with strong westerly flow (0.4 m/sec). At 10 S, this closed circulation pattern becomes mixed with South Equatorial Current (SEC) and sets up an alternating strong northward and southward flow (1 m/sec). Above 200 m, the northward flow is dominant, while at deeper levels the southward flow becomes prominent (Ramesh Babu et al., 2001). Seasonal variability of currents in the IONF was determined from four depth zones: subsurface (450–670 m), intermediate (1150–1370 m), deeper (3450–3670 m) and near-bottom (4270–5100 m) at eight current metre mooring stations (Murty et al., 2001). Of these, three stations were from the northern part of the IONF (10 S, sector A) and the remaining five stations from the southern IONF (15 S, sector D). In the northern IONF, seasonal variations in currents were found at depths of 500 and 1200 m, caused probably by seasonal north-south shift in the westward-flowing SEC. In the southern IONF, a dominant eastward flow at 500 m suggests the occurrence of current shear below the SEC. Low-frequency oscillation of 30–60 days superimposed on high-frequency inertial, diurnal and semi-diurnal,

27


meso-scale fluctuations was noticed at all depths. In the northern IONF, the total kinetic energy (KE) recorded high values during spring–summer and low values during fall–winter. Predominance of zonal flows over meridional flows, intensification of currents and decrease in KE through the water column were some of the interesting features in the IONF. At deeper zones, synoptic-scale oscillation (12–15 days) was responsible for higher energy currents in the northern IONF, while mesoscale oscillation influenced the southern IONF (Table 1.6).

2.2. Chemical characteristics The chemical properties of the IONF water column (Table 1.7) reveal the existence of three north-moving distinct water masses: (1) the water mass between 125- and 200-m depth characterised by high salinity (34.74–34.77 psu) and oxygen minima, associated with weak load of nutrients, (2) the deep oxygen maxima in the depth range of 250–750 m associated with minima in nutrients and relatively high pH and (3) the salinity minima (34.714–34.718 psu) at depths between 800 and 1200 m. The third water mass, in the density range of 27.2–27.5, corresponds to the Antarctic Intermediate Water (AAIW) (De Sousa et al., 2001). Detailed study suggests that changes in water masses are probably the result of mixing of various layers of waters, and the variable oxidation of organic detritus en route. The oxygen maxima water mass occurring between 250 and 750 m shows the least changes in these properties and subsequently moves faster. The variable concentration of dissolved organic carbon (DOC) with depth is related to the biochemical activity as well as to the chemical features of the water column. DOC is least in the oxygen minimum zone and maximum in subsurface waters. The lack of significant correlation between DOC and apparent oxygen utilisation may suggest

Table 1.6 Velocity, direction and kinetic energy of currents in the IONF

Location

Sector A (Sept.–Jan.)

Sector D (Sept.–Jan.)

Water column depth Velocity Direction Total (m) (cm/s) ( )

Kinetic energy (cm2/sec2)

Mean

Eddy

Syno

Meso

0500 1200 3500 4900 0600 1300 3600 4400

13.68 04.30 00.82 01.47 02.44 00.76 00.00 –

159.04 45.85 05.00 07.36 12.76 04.57 00.43 –

145.27 40.42 03.14 05.92 05.91 00.87 00.06 –

13.77 05.43 01.86 01.43 06.85 03.70 00.37 –

5.23 2.93 1.28 1.72 2.21 1.23 0.07 0.11

Source: Murty et al. (2001). Note: IONF ¼ Indian Ocean Nodule Field.

329 269 271 240 076 117 132 277

172.72 50.16 05.82 08.03 15.20 05.33 0.44 –

28 Table 1.7


Average chemical characteristics of water column in the IONF

Water depth (m)

pH

Dissolved oxygen

Nitrate

Phosphate

Salinity

Surface 1000 2000 3000 4000 5000

8.2 7.6 7.66 7.70 7.70 7.70

220 90 125 185 185 185

00 36 35.9 35.5 35.5 35.5

0.10 2.45 2.45 2.40 2.25 1.90

34.64 34.715 34.72 34.72 34.72 34.72

13 S

–

97.5 8.9

36.1 2

2.59 0.24

34.714 0.001

Sources: Modified from De Sousa et al., (2001) and Sardessai and De Sousa (2001). Note: Units for dissolved oxygen, nitrate and phosphate in micrometer, salinity in psu. Sector C is the most promising resource area in the Indian Ocean Nodule Field (IONF), hence shown specially. Water depth at 13 S (Sector C) is 4960 m.

simultaneous consumption of oxygen by other species in the water column (Sardessai and De Sousa, 2001).

2.3. Biological characteristics The benthic community consists of two major members: macrofauna (size >500 mm) and meiofauna (size 3 cm) lies in the fact that they live in highly stressful environment under extreme hydrological conditions, such as high pressure, low temperature, low oxygen, low salinity, minimum water current, absence of light and a low rate of sedimentation. Because

Table 1.8

Distribution of organisms in sediment layer in the IONF

A. Major macrobenthos (%) Annelida Crustacea B. Major macrofauna (%) Polychaeta Harpacticoida Nematoda Tanaidacea C. Bacterial and organic nutrients

Grab # 191 195 196

Total bacteria (1010/g) 10.11 17.02 39.29

Living biomass (mg/g) 0.059 0.027 0.020

Carbohydrates (mg/g) 2.4 2.4 2.1

0–2 cm 44.5 14.3 0–2 cm 31.6 15.8 38.1 7.90

Protein (mg/g) 0.98 0.87 0.66

Lipids (mg/g) 0.39 0.23 0.01

2–5 cm 7.1 3.3 2–5 cm 30.83 23.05 38.33 7.78

5–10 cm 9.5 6.2 5–10 cm 35.08 24.95 24.95 15.00

10–20 cm 3.9 8.0 10–20 cm 11.67 5.95 17.62 41.19

Labile organic matter (mg/g) 3.787 3.514 2.754

Total organic carbon (mg/g) 5.2 7.3 7.3

Total organic matter (mg/g) 9.4 13.2 13.2

Sources: Ingole et al. (2001) and Raghukumar et al. (2001a,b). Note: Grab locations: 191 (10 10.260 S, 76 00.240 E), 195 (10 05.410 S, 76 05.380 E) and 196 (10 02.450 S, 76 02.520 E). IONF ¼ Indian Ocean Nodule Field.

30


of difficulties in observing and sampling deep-sea organisms, megafauna are normally studied from underwater photographs. However, in an in situ study (Rodrigues et al., 2001), 11 groups of invertebrates—xenophyophores, sea cucumber, holothurians, sponges, sea anemones, sea pens, black corals, shrimps, starfish, brittle star, sea flowers, sea urchins—and one group of vertebrate, namely, fish (Typhlonus nasus), were found on the floor of the IONF. It was recorded that xenophyophores and sea cucumber constitute about 41% and 30%, respectively, of the total megafauna present in the IONF. In addition, a few protuberant molluscs, polychate worms, sea fans and squids were also found.

(b) 0−2

0 2 4 6 8 10 12 14

(a) 0−2

0 2 4 6 8 10 12 14

A

2−4

2−4

4−6

4−6

0−2

15−20

2−4 197

(c) 0−2

20−25 4−6

(d) 8−10

0.12

0.08

0.06

0.04

6−8 0−2

10−15

2−4

194

25−30 0.00

0.12

0.10

0.08

0.06

0.04

0 2 4 6 8 10 12 14

0.00 0.02

15−20

0.02

10−15

10−15

0 2 4 6 8 10 12 14

8−10

8−10

0.10

(e)

0 2 4 6 8 10 12 14

6−8 6−8

2−4 193

15−20

8−10

LOM Biomass

10−15

0.12

0.10

0.08

0.06

0.04

6−8

0.02

4−6 0.00

4−6

6−8 8−10 10−15

TOC

190

188 0.10

0.08

0.06

0.04

0.00 0.02

25−30

Vertical scale refers to depth in sediment core in cm. Top horizontal scale represents concentration of LOM and TOC (mg/g of dry sediment) Bottom horizontal scale measuring concentration of biomass (mg C/g of dry sediment)

Figure 1.16

(Continued)

0.12

0.10

0.08

0.06

0.04

0.00

20−25

0.02

15−20

15−20

31


B (a)

0.0 0.2 0.4 0.6 0.8 1.0 1.2 1.4

0.0 0.2 0.4 0.6 0.8 1.0 1.2 1.4

(b) 0−2

0−2

2−4 2−4 1.2

4−6

4−6 0.0 0.2 0.4 0.6 0.8 1.0

(e) 0−2

6−8 8−10

2−4

10−15

4−6

6−8 8−10 10−15 15−20 20−25

197

0−2

6−8

0.0 0.5 1.0 1.5 2.0 2.5 3.0 3.5 8−10

(d)

0−2

10−15 193

2−4

2−4 0.0 0.5 1.0 1.5 2.0 2.5 3.0

15−20

4−6 6−8

4−6 6−8

8−10 10−15

Protein

15−20

CHO Lipids

8−10 10−15 190

188 15−20 0.0 0.5 1.0 1.5 2.0 2.5 3.0 3.5

25−30

0.0 0.5 1.0 1.5 2.0 2.5 3.0 3.5

20−25

194

25−30 0.0 0.2 0.4 0.6 0.8 1.0 1.2 1.4

0.0 0.2 0.4 0.6 0.8 1.0 1.2 1.4

(c)

2.5 3.0 3.5

0.0 0.5 1.0 1.5 2.0

15−20

Vertical scale refers to depth in sediment core in cm. Top horizontal scale represents concentration of lipids and protein (mg/g of dry sediment) Bottom horizontal scale refers to concentration of carbohydrates (mg/g of dry sediment)

Figure 1.16 Biology of the Central Indian Ocean Basin (CIOB) (Raghukumar et al., 2001a,b), showing (A) concentration of total organic carbon (TOC), labile organic matter (LOM) and biomass-C, and (B) lipids, proteins and carbohydrates in five sediment cores (a ¼ 197, b ¼ 194, c ¼ 188, d ¼ 190, e ¼193).

3. Evolution of the Indian Ocean Nodule Field India has been a central element to any tectonic reconstruction of the world for its unique history of cruising for several thousands of kilometres to join Laurasia after being detached from Gondwanaland. And expectedly there have been schools of theories over the exact nature of dismemberment of Gondwanaland based on geological, geophysical and isotopic signatures, and those established through the distribution of tetrapods, Glossopteris flora and Permo-Carboniferous glacial strata.

32


Since the break-up of Gondwanaland and the evolution of the IORS have influenced the volcanic and tectonic environments of the IONF, some discussion on the evolution of the IONF seems necessary.

3.1. Break-Up of Gondwanaland The reconstruction of Gondwanaland is primarily based on geophysical data, seafloor-spreading magnetic anomalies, transform faults, fracture zones, satellitebased sea surface altimetry and paleomagnetism. Following several studies (e.g. McKenzie and Sclater, 1971; Norton and Sclater, 1979; Reeves and de Wit, 2000; Royer et al., 1992; Schlich, 1982), the major tectonic episodes are recognised as a prelude to the break-up of Gondwanaland and formation of the Indian Ocean (Table 1.9; Fig. 1.17). The break-up of Gondwanaland is believed to have commenced during the Jurassic era (180–200 Ma), with eastern Gondwanaland, comprising Madagascar, Seychelles, India, Antarctica and Australia, separating from western Gondwanaland, made up of South America and Africa. Eastern Gondwanaland broke further in the Late Mesozoic (about 133 Ma), with Madagascar–Seychelles–India (Greater India) rifting away from Antarctica–Australia along the old Eastern Ghat trend of eastern India through a clockwise rotation of the latter by about 20 (Muller et al., 1993). Uplift associated with this rifting tilted peninsular India to the west, resulting in the Table 1.9

Dismemberment of Gondwanaland and evolution of the Indian Ocean

Phase 1 Early part (160–130 Ma): Gondwanaland divides into two groups—Group A comprising South America [SA], Africa [AF], Somalia [SM] and Arabia [AR] and Group B consisting of Madagascar [MD], India [IN], Seychelles [SE], Sri Lanka [SL], Australia [AS] and East Antarctica [EA]. Late part (130–95 Ma): Group A divides further into two sub-groups [SA] and [AF þ SM þ AR], and the Group B also into two sub-groups [MD þ SE þ IN þ SL] and [AS þ EA] Phase 2 Early part (96–65 Ma): [MD] of Group B joins [AF þ SM þ AR] of Group A, while [AS] and [EA] separate from each other. Late part (65–43 Ma): [SE] of Group B separates from [SL þ IN] and joins the [AF þ SM þ AR þ MD] of Group A. Phase 3 Early part (43–10 Ma): [SA] remains separated from [AF þ SM þ AR þ MD þ SE]. [AS] leaves EA to join [IN þ SL]. Late part (10–00 Ma): [AR] separates from [AF þ SM þ AR þ MD þ SE]. [AS] leaves [IN þ SL]. Source: Royer et al. (1992).

33


D

A

EURASIA

ARB

PA

Pa le

o-

A

Panthalassic Ocean

NG

E A

SAM

d u c ti o n z o n e

CI Ne

MM

Te t

AFR

hy s

RH

ER

IA

o-

AFR

Te t

MAD

NAM

S ub

hy s

AUS

IND

spreading centre

AUS ANT

ANT

B

E CIMMERIA ARB Ne

Te t

hy s

AFR

SEY

ARB

AFR

sp ce read ntr in g e

G O N DWA

SAM

o-

RH AUS

NA

L

on

ti uc bd Su ne zo

IND

A

ND

AUS

ANT

C

F EURASIA

ARB ARB AFR

CR

CIR

AFR

IND

AUS

AUS

SW

IR

MAD

RH

SE

IR

Figure 1.17 Dismemberment of Gondwanaland blocks and transhipment of India during geological periods (cf. Bhattacharya and Chaubey, 2001). (A) at >200 Ma, (B) at 150 Ma, (C ) at 84 Ma, (D) at 50 Ma, (E) at 30 Ma, (F) presentday.

34


flow of the major rivers from east to west (Cox and Hart, 1986). Continuous movement across the ridge between Greater India and Australia–Antarctica caused the evolution of the Indian Ocean. Around 118 Ma, ocean-floor spreading between Greater India-Madagascar and Africa stopped. The Middle Cretaceous (95 6 Ma) is widely recognised as the first era of major plate reorganisation in the Indian Ocean, with Australia breaking away from Antarctica and India separating from Madagascar. The slight clockwise rotation of India away from Madagascar amounts to trans-tensional rifting represented then by about 200-km wide wedge of sediment-filled extended crust, which is presently preserved as the western continental shelf of India (Reeves and de Wit, 2000). The separation between Greater India and Madagascar was caused by two phenomena occurring at two different periods. First, the 1000-km southward movement of India away from Madagascar before 90 Ma was caused by prolonged dextral trans-tension along the line of weakness between the east coast of Madagascar and the west coast of India (megashear of Reeves and Leven, 2000). Later, the rapid movement of India towards the north was facilitated by widespread eruption of basalts and rhyolites from the Marion hotspot at about 88–89 Ma. The volcanism was so voluminous that the entire 1500-km long rifted eastern margin (megashear?) of Madagascar was constructed within 6 million years (Storey et al., 1995). The ridge, about which India got separated from Madagascar, is presently believed to exist in the form of abandoned spreading centres on the conjugate crust to the east of Madagascar on the African Plate and in the Laxmi Basin in the eastern Arabian Sea on the Indian Plate (Bhattacharya et al., 1994; see Fig. 1.1). The development of NNW-SSE-trending magnetic lineation, grabens, ridges and basins on the marginal shelf and slope of south-western India between 84 and 65 Ma has been interpreted to represent a two-limbed spreading sequence in the Laxmi Basin. Despite the difference in free-air gravity and isostatic signatures, some workers suggested Laxmi Ridge (having a negative anomaly) as the continental continuation (slivers) of the CLR (showing a positive anomaly) formed by rifting along the margin (Kolla and Coumes, 1990; Naini and Talwani, 1982). However, Todal and Edholm (1998) characterised this enigmatic ridge as a complex marginal high of both continental and oceanic crust, which might have experienced later magmatic and/or tectonic deformation also. Around 84 Ma began the formation of the CIOB (including the IONF). The opening of the Arabian Sea commenced during the Late Cretaceous–Early Tertiary when the ridge system in the Mascarene Basin ceased spreading and jumped to the north to form the CR at about 68–66 Ma, resulting in the separation of Seychelles from India and its transfer to the African Plate (McKenzie and Sclater, 1971). Around this time, one of the largest continental flood basalts, the Deccan Traps, erupted and covered vast areas of the central and western India and the Seychelles. As India was moving towards Eurasia, the Reunion hotspot perforated the crust to give rise to the Chagos–Laccadive–Mascarene Ridge. Between 60 and 50 Ma, the rate of northward movement of India varied from fast (95 mm/year, half rate) to slow (26 mm/year). The slow spreading rate might have been caused by the collision of India with Eurasia. However, the exact timing of this collision is controversial. It is suggested that the first collision between India


35

and Eurasia occurred around 50 Ma (Middle Eocene) with noticeable decrease in spreading rates across the CR, CIR and SEIR (Windley, 1996). A critical appreciation of the seafloor lineaments and the deformation caused by lithospheric compression in a north-south direction suggest an initial touch between India and Eurasia during 58 Ma, and starting of collision during 51 Ma (Mukhopadhyay et al., 1997). However, radiometric dating of garnet and zircon from Himalayan rocks tentatively shows that the huge mountain-building activity of the Himalayas, which resulted from the collision of India with Eurasia, in fact, dates to more than 450 Ma, that is, 9 times older than the previously suggested age (Gehrels et al., 2003). It is proposed that sometime after the collision, India pulled back. Then, around 50 Ma it ploughed into Asia once again, causing episodic mountain-building activity. The second major plate reorganisation event in the Indian Ocean occurred between 45 and 38 Ma. The event saw the spreading between Australia and Antarctica, which started at 45 Ma, resulting in the merging of the India–Antarctic and India–Australia ridges to form the SEIR. Meanwhile, the Chagos-Maldives and Mascarene plateaus broke up around 35.5 Ma, the former becoming a part of the Indian Plate and the latter stationed on the African Plate. During this period spreading across the Wharton Ridge ceased, and later jumped to the south (Liu et al., 1983). The younger tectonic events in the Indian Ocean included opening of the Gulf of Aden at about 10 Ma (Laughton et al., 1970) and onset of deformation between the Indian and the Australian plates at 7–8 Ma in the CIOB, just north of the IONF (Krishna et al., 1998; Weissel et al., 1980).

3.2. Formation of the Indian Ocean Ridge system The evolutionary history of the IORS has been understood largely from the research findings of the last three decades (Table 1.10; Fig. 1.17). Between anomalies 33 and 30, the Indian Ocean Triple Junction (IOTJ) was lying in the eastern Mascarene Basin, about 300 km west of the Vishnu FZ. At anomaly 29, the IOTJ rapidly migrated eastward and was located between the Vishnu FZ and Indrani FZ. The CIR and SEIR were being progressively offset at a rate of 1.4 mm/year, mainly because of their different spreading rates. When the offset was large, the IOTJ jumped eastward and a new CIR segment was created,resulting in an overall constant length for the SEIR, while the CIR and SWIR lengthened. A jump of every 1 Ma created ten 70-kmspaced SWIR fracture zones, including Gallieni and Atlantis II fracture zones. The close association of numerous fracture zones across the SWIR and IOTJ jumps seems to suggest that the physiography of the SWIR records the history of the IOTJ evolution. The Indian Ocean has examples of many ridge jumps. For instance, the India– Madagascar Ridge, earlier existing between India and Madagascar, ceased its activity at about 68 Ma, and took a major jump towards India before opening of the Arabian Sea (Naini and Talwani, 1982). Because of this ridge-jump process, the northern part of the present CIR came into existence about 30 Ma ago, and transferred the Reunion hotspot from the Indian Plate to the African Plate. Similarly, a major ridge jump northward between chrons 31 and 25 transferred the Crozet hotspot from the Indian Plate to the Antarctica Plate (Muller et al., 1993). The CIR has experienced few minor ridge jumps since about 3 Ma (Mitchell, 1991).

36


Table 1.10 The evolving IOTJ Position of IOTJ Period

Age (Ma)

Anomaly

latitude

longitude

Spreading rate (cm/year)

Present Pleistocene Upper Pliocene Middle Pliocene Lower Pliocene Early Oligocene Early Eocene Palaeocene

00 02 04 06 08 38 53 65

00 01 02 03 04 16 22 28

25.7 S 25.7 S 25.7 S 25.7 S 25.7 S 25 S 26.50 S 31 S

70 E 70 E 69.02 E 68.7 E 68 E 63.52 E 60 E 54.6 E

1.84 1.62 2.70 3.78 – 2.88 6.75 –

Sources: Norton and Sclater (1979), Tapascott et al. (1980) and Murthy and Rao (1992). Note: IOTJ ¼ Indian Ocean Triple Junction.

Young oceanic crust forming at the IORS normally consists of tholeiitic basalt (N-MORB) characterised by low concentrations of incompatible elements such as K, Rb, Cs, Ba and the light rare earth elements (LREE). The basalts appear to have been derived from a mantle depleted in these elements. Three major rock types are found in the IORS, each with characteristic phenocrystic assemblage reflecting a distinctive magmatic lineage of differentiation. The rocks have low, moderate and high abundances of Na2O and TiO2 with sequentially lower CaO/Al2O3. The Indian MORB, however, is distinct from the Atlantic and the Pacific MORB in terms of Sr, Nd and Pb isotopic composition (Table 1.5). Isotopic ratios of Sr, Nd and Hf in N-MORB further demonstrate that the incompatible elementdepleted melt is inherited from a portion of the asthenosphere, while other relatively enriched (less depleted), the so-called enriched or transitional, MORB are associated with mantle plumes. The differences in isotopic composition among the rocks from the three oceans may have been possible by any one or a combination of two or more of the following processes: (1) recycling of the ancient oceanic crust (Dupre and Allegre, 1983), (2) contamination of ridge melt by the Kerguelen hotspot [plume] melt (Storey et al., 1989), (3) mixing of sub-Gondwanaland continental lithosphere (Mahoney et al., 1989) and (4) contamination of ridge melt from several hotspot plumes that formed oceanic islands (Mahoney et al., 1993).

C H A P T E R

T W O

Tectonics and Geomorphology

Contents 40 41 43 44 44 45 47 48 49 52 54 56 59 60 61 63

1. Ridge-Normal Lineaments 1.1. The Vishnu Fracture Zone 1.2. The Indrani Fracture Zone 1.3. The Indira Fracture Zone 1.4. Trace of the triple junction 2. Ridge Parallel Lineations and Anomalies 2.1. Seafloor crenulations 2.2. Seafloor faults 2.3. Tectonic, thermal and geoidal anomalies 2.4. India–Eurasia collision event revisited 3. Seamounts 3.1. Spreading rate and seamount distribution 3.2. Structural lineaments and seamount abundance 3.3. Structural lineaments and seamount morphology 3.4. Seamount petrology 3.5. Growth conditions of seamounts

As mentioned in the previous chapter, the Indian Ocean Nodule Field (IONF) contains the world’s second largest and second richest manganese nodule resources after the Equatorial North Pacific nodule belt. The structural boundaries of the IONF, which appear to have an influence on resource characteristics, are also described in the preceding chapter. In this chapter, we focus on tectonic and geomorphological characteristics of this nodule field. The detailed bathymetric information of the IONF was acquired through single (narrow and wide) beam echosounder and multi-beam (59 beams) swath bathymetry systems (Fig. 2.1). These data acquired since the early 1980s helped improve our understanding of seafloor bathymetry compared to that gained through General Bathymetric Chart of the Oceans (GEBCO) maps and other available information. The IONF (area 690 103 km2) interestingly encompasses three types of crust that were generated from the Indian Ocean Ridge System (IORS) during Palaeocene–Eocene time between 60 and 49 Ma, and formed during pre-, syn- and post-collision periods between India and Eurasia. These periods respectively correspond to fast, intermediate and slow rates of generation of crust from the ridge axis. Consequent to such tectonic conditions involving tensional and compressive stresses Handbook of Exploration and Environmental Geochemistry, Volume 10 ISSN 1874-2734, DOI: 10.1016/S1874-2734(07)10002-4

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37

38


caused by the fast- and slow-spreading regimes respectively, the floor of the IONF is riddled with the following features.

Broadly NNE-SSW-oriented ridge-normal ancient fracture zones (FZ, major structural lineaments). Folding, stretching and sometimes faulting of the seafloor along an approximately ESE-WNW direction, which trends parallel to the then mid-ocean ridge (MOR) axis. Numerous volcanic features in the form of seamounts and abyssal hills. However, only a few integrated studies have been made to unravel the relations among the N-S to NNE-SSW-trending regional lineaments, ESE-WNW-oriented local tectonic activities and seamount volcanism (Mukherjee and Iyer, 1999; Mukhopadhyay et al., 2002). These studies brought out new fundamental information regarding basinal dynamics, which probably occurred at two stages: first, the tectonic and volcanic activities associated with the formation of the oceanic crust at the ridge crest between 60 and 49 Ma; and second, mid-plate activities on the said crust during its journey from the ridge crest to the present location in the abyssal region. These works offered possible answers to quite a few outstanding problems, such as how the spreading of the crust and the eruption kinematics altered in the geological past, how the crustal accretion and the presence of fracture zones influence the origin, distribution and morphology of near-axis seamount production and how the seamounts grew in response to mid-plate tectonic activities and local volcanism. A 72E

74

76

78

80 27

Chain E

24

26 Indrani FZ

Chain G

25

26

Chain B

12

Vishn u FZ

25

Chain C

26

Chain A

10 S

Chain H

26 27

25 25

TJT-In

23 24

24

14

16

Figure 2.1 (Continued)

Chain F

Chain D

23

39


B

1000 S

1030

1100

76E

7620

1230

1300

1330

1130

1400

1200

1430

7600E

7620E

Figure 2.1 (A) Essential tectonic elements in the Indian Ocean Nodule Field (IONF) with seamounts (stars) arranged in eight prominent chains (A to H), trace of the triple junction on the Indian Plate (TJT-In, dashed line), and nearly east-west-running magnetic anomalies dislocated at places by NNE-SSW-running fracture zones (Das et al., 2005).The crust of the IONF was generated at the ridge crest between 60 and 45 Ma. (B) A simplified bathymetry map of a part of the IONF prepared from multi-beam swath survey input shows seamounts (closed contours with numbers) of various dimensions. Seafloor crumpling (folding and faulting) of variable intensities is marked with nearly east-west-oriented bold and thin short lines with tick marks indicating dip direction. Note more crumpling at the southern latitudes, and occurrence of both north-facing normal and south-facing reverse faults (Mukhopadhyay and Batiza,1994).

40


1. Ridge-Normal Lineaments The ridge-normal structural lineaments in the IONF are largely oriented N-S to NNE-SSW, and lay perpendicular to the present-day Southeast Indian Ridge (SEIR). Among the ridge-normal lineaments, the Vishnu FZ (along 73 E), Indrani FZ (along 79 E) and Indira FZ (along 83 E) are important (Figs. 1.2, 2.1). A magmatically pronounced approximately NNE-SSW-oriented trace of the ancient plate boundary [i.e. trace of the triple junction on the Indian Plate (TJT-In)] along 76 300 E meridian has been another important feature. Although located outside the IONF, the influence of Indira FZ on dynamics and resource potential of the field has been considerable and hence included here for discussion. All these structural lineaments have disturbed the seemingly smooth gradient (1:1000 to 1:7000) of the IONF seafloor. The various characteristics of these structural features are tabulated in Table 2.1. The formation algorithm of a fracture zone at the MOR axis is quite interesting in constraining the movement of the plate in response to variable spreading rate and consequent stress regime. Transform faults offset the ridge axes as oceanic plates slip sideways past each other. As the seafloor spreads, the imprint of the transform fault develops on crust on either side of the ridge. Such traces on older crust are known as fracture zones, which are useful in reconstruction of ancient plate motions. The fracture zones normally lie within a single oceanic plate, do not register any relative lateral motion across it and exhibit no major seismic activity. They can be demarcated by offset in magnetic lineation. The orientation and nature of the fracture zone hold fair and accurate evidence of change in spreading direction in the geological past, as these normally offset the magnetic lineation right-laterally. In the CIOB, the progressive northward shift of the oceanic crust interspersed with the fracture zone may have been caused because of ridge jumps on the SEIR, the magnitude of which decreased from east to west (Royer and Schlich, 1988). For example, the Indrani and Indira FZ appear to have undergone a change in trend from N-S to NE-SW as a consequence of the Eocene Table 2.1

Some characteristics of fracture zones in the Indian Ocean Nodule Field

Fracture zone (FZ) Orientation

Width (km)

Central portion

Offset (km)

Vishnu 73 E

13–17

Horst

70

Indrani 79 E

Small

Trough

—

83 E

Small

Graben

160

Indira

Contact with surrounding seafloor Movement West

East

Right lateral Right lateral Right lateral

Sharp

Smooth

Smooth

Uneven

Sharp

Subsided

Sources: Tapscott et al. (1980), Kodagali (1992), Kamesh Raju et al. (1993), Mukhopadhyay et al. (1994) and Kessarkar (1998).


41

reorientation of the SEIR, which changed direction from E-W to NW-SE during this era (between anomaly 18 and 20) marked by global plate reorganisation between 40 and 44 Ma (Kamesh Raju, 1990). Co-incidentally, the obliquity between the SEIR and the CIR increased from 6 to 12 between magnetic anomaly A26 and A24; decreased to 3 at A23; and again increased to 7 at A22 (Dyment, 1993). The ridge normal lineaments are described below, very briefly.

1.1. The Vishnu Fracture Zone The Vishnu FZ is located on the Indian Plate running parallel to the meridian 73 E (trending approximately N15E in the IONF), while its conjugate part can be traced on the African Plate (Dyment et al., 1999). A detailed multi-beam swath bathymetric study of a part of this fracture zone reveals that this fracture zone is essentially a much wider one compared with other known fracture zones in the world’s oceans. The Vishnu FZ runs for hundreds of kilometres and varies in width from about 13 km in the north to about 17 km in the south. The fracture zone shows very sharply delineated contacts with the abyssal plain along the eastern and western margins (Fig. 2.2). The average throw along the western contact margin ranges between 400 and 500 m, while that along the eastern contact is about 200 m, making the western contact more distinct and pronounced. Between the eastern and western margins of the fracture zone, the crust was domed by about 300 m along its length. This elevated central portion hosts many small abyssal hills with ESE-WNW orientation. The fine-scale bathymetric variations reveal that the average amplitude and wavelength of the seafloor crenulations in the Vishnu FZ has been 233 m (range 40–530 m) and 22.5 km (range 8.1–38.2 km), respectively, and is 5230-m deep in the north. The disposition of magnetic anomalies and nature of ridge-parallel lineation in the surrounding areas suggest formation of the Vishnu FZ on the crust spreading at variable rates, 80 mm/year between 57.9 and 51.3 Ma and 36 mm/year between 51.3 and 46.3 Ma (henceforth spreading rates correspond to half-rates). This fracture zone has offset the magnetic lineation of the same age right-laterally by about 70 km; the offset is larger in the eastern side than in the western side (Kamesh Raju, 1990). The northern and the central parts of the fracture zone trend N15E (Fig. 2.2), while its southern part, beyond the southern limit of the IONF, shows a striking change in trend from N15E to N45E. This change in trend, as mentioned earlier, may have been caused by the global plate reorganisation event at about 42 Ma. Magmatically, the Vishnu FZ appears to be dormant for quite long periods. The rocks dredged from this fracture zone are either largely weathered and/or have thick coatings of ferromanganese oxides—both suggesting lack of recent volcanism. The fracture zone reveals wide zones of brittle deformation with highly fractured, permeable rocks—the cracks possibly penetrating deep down into the crust (Mukhopadhyay et al., 1994). Such opening facilitates mixing of hot circulating seawater within the crust with percolating cool seawater and results in reducing the build-up of pressure within the crust (or magma chamber), inhibiting magma ascension and eruption.

42


7330E

7345

740 948 S

A

Vishnu

Fractu r

e Zone

1000

1015

Vish

nu F ra

cture

Zone

B

11S

1153 73E

74

Figure 2.2 (A) A part of the NNE-SSW-oriented Vishnu Fracture Zone(FZ) along the 73 E. Note the sharp western and comparatively diffused eastern contacts with the seafloor and the presence of several large seamounts (height >500 m). The structural lineations on the seafloor are manifested by east-west-closed contours. Note its bend contacts with the fracture zone, suggesting a much younger age for the Vishnu FZ (Das et al., 2005). (B) Simplified bathymetry of a

43


1.2. The Indrani Fracture Zone The Indrani FZ lies on either side of the SEIR, and has its southern segment extending northward from the eastern edge of Crozet Island on the Antarctic Plate. This fracture zone runs parallel approximately to 79 E and reveals ridge-trough topography with an elevation difference of about 300 m (Fig. 2.3). The topography is associated with steep gradients in the south and is gentler in the north. The seafloor along the western flank is shallower than that of the eastern flank. Seamounts near this fracture zone are mostly located in the southern part. 79E

7923 14 S

12 S

1417

14

7810E

Indrani Fracture Zon

e

13

79

7925

Figure 2.3 Tectonics and physiography of NNE-SSW-oriented Indrani Fracture Zone along the 79 E (Kamesh Raju et al.,1993). Disposition of a group of seamounts occurring close to the fracture zone is shown in inset (Das et al., 2005).

part of the13^17-km wideVishnu FZ. Contour interval 100 m, with tick marked contour at 500 m. Seafloor lineation of at least three orientations (NW-SE at the south-western part, NNE-SSW at the north-western part and WNW-ESE trends at the eastern part) make the area tectonically complex (Kessarkar,1996).

44


Several east-west-trending ridge-parallel bathymetric lineations curve towards the south at the contact with this fracture zone. This bending of lineations occurs particularly on the crust formed at a fast-spreading rate between 60 and 50 Ma ( anomaly A26–A23) in the Indian Ocean (80 mm/year, spreading half-rate), suggesting a fast slipping transform fault environment during this time. The causative factor for such bending could have been the strong shear coupling at the contact of the fracture zone with the spreading ridge. The free-air gravity values along this fracture zone vary from 0 to 45 mGal. Lithospheric thickness of this fracture zone, deduced from the gravitational edge effect, is estimated to be around 100 km. The elastic limit of the lithosphere underneath, deduced from thermal structure and horizontal heat conduction, has been estimated to be 23 km (Te; Kamesh Raju et al., 1993).

1.3. The Indira Fracture Zone This fracture zone runs in a north-south direction roughly along 83 E between 3 N and 22 S, and offsets the magnetic lineation from A20 to A34 right-laterally by about 160 km. The single channel seismic profile of the southern part of the Indira FZ indicates that the seafloor topography manifests as a discrete U-shaped graben with differing flank relief. The depth to this graben gradually increases from the surrounding flanks to about 1700 m. The axial part of the graben consists of micro-troughs and rises. The western flank of this fracture zone is shallower by 750 m and contains steplike structures. Because of its older age, the eastern flank is deeper than the western one. Faults are found to exist mostly at the subsurface level, though some, reaching to a greater depth, probably developed later than the main fracture zone. The sediments within this fracture zone are distributed heterogeneously and show complex acoustic character and structure. While a thin layer of pelagic sediment carpets the fracture zone without changing the graben’s identity in the south, the northern part is filled and buried by thick sediments. At about 1 S, the fracture zone appears as an anticline and remains fully buried under the Bengal fan sediments. The likely reasons for such structural change could be the inconsistent shape of the fracture zone, which has changed with time depending on the evolutionary conditions such as spreading rates, volcanism near the ridge-transform fault intersection (RTI) and offset between the ridge segments (Krishna, 1996).

1.4. Trace of the triple junction The Indian Ocean Triple Junction (IOTJ) is essentially the meeting point of three spreading ridges, the CIR, the SEIR and the SWIR, and is a magmatically pronounced plate boundary. Although the IOTJ at present is located at 25 S, 70 E, it moved over the geological past and ran approximately parallel to 76 300 E, the course now represented by a linear trace (Fig. 2.4). The northern half of the TJT-In is located between the Vishnu FZ and the Indrani FZ on the Indo-Australian Plate, with its western and southern conjugate parts lying on both sides of the SWIR on the African and the Antarctic plates (Dyment, 1993; Patriat and Segoufin, 1988). The nature of evolution of this triple junction has been unique. For example, the recent data obtained favours a ridge–ridge–ridge (RRR) type configuration for the IOTJ corresponding to magmatic mode, and ridge–ridge–fault (RRF) configuration

45


75E

78

FZ

12S

FZ

15

>300 nt

200−300 nt

100−200 nt

0−100 nt

0− −100 nt

< −100 nt

Figure 2.4 Magnetic anomalies of a part of the Indian Ocean Nodule Field (IONF), often disturbed by a fracture zone (FZ).The fracture zone was later verified as trace of the triple junction, that is, the meeting point of three ridges on the Indian Plate (Kamesh Raju and Ramprasad,1989; Dyment,1993).

relating to tectonic mode. Because of the slow-spreading rate of the CIR and the intermediate-spreading rate of the SEIR, these ridges are being progressively offset. The CIR changed its orientation with time, for example, from approximately E-W at 57.7 Ma to ESE-WNW at 52.8 Ma, and reverted again to E-W between 52.8 and 47.1 Ma. The SEIR–CIR obliquity was least at 51.3 Ma (3 ) between two elevated zeniths (6 –12 , 7 ) just before and after this period. Propagating rifts appear to have played an important role in the evolution of the triple junction, particularly between the period 63.6 and 46.3 Ma. The volcanic and magmatic potential of this IOTJ trace can be gauged from the occurrence of several volcanic forms and seamounts in its proximity (discussed in Section 2.3).

2. Ridge Parallel Lineations and Anomalies The average depth of the IONF increases from west to east, and the nodule field can be broadly divided longitudinally into three bathymetric areas. The western part (71 –74 E) is extremely rugged with great variations in relief

46


(depth range 2900–5000 m). The eastern part of the field between 79 E and 82 E is a medium relief area, with a narrow depth variation (5000–5500 m), while the central part between the two uneven regimes (74 E to 79 E) is almost a plain area (4900–5100 m). In general, about 90% of the IONF area has a slope angle 58 90

10.26 – 10.96 58–56 55

56–61 95

Beyond 13.76 100m throw Abundance of normal faults (%) Abundance of reverse faults (%)

19.60 21 23 None ND None

27.88 18 25 33.75 31.32 36.44

22.41 20 23 22.11 40.66 22.03

30.10 08 26 44.12 28.02 41.52

Sources: Mukhopadhyay et al. (1997) and Kessarkar (1996, 1998). Note: IONF ¼ Indian Ocean Nodule Field; ND ¼ Not determinable due to sediment cover.

variations may reflect the low and high degrees of stress, respectively, to which the crust was subjected. It appears that a spectrum of distortion of the seafloor exists, that is, least distorted, older, thicker and stronger crust in the north (sector A) to highly crumpled, relatively thinner, younger and weaker crust in the south (sector D). The relationship between seafloor roughness and spreading rate in the world’s oceans is best expressed by the empirical orthogonal functions (EOF) analysis and by the root-mean square (RMS) inverse equation. The EOF analysis offers a quantitative distinction between the deterministic and the stochastic components of mid-ocean ridge topography, whereas the RMS is a conventional and well-proven procedure. The rapidly accreted areas in the IONF (north and south-central, sectors A and C, respectively) are theoretically calibrated to represent the RMS roughness of the order of 65 and 90 m, respectively (Table 2.2). Similarly, for the intermediate (northcentral, sector B) and slow (southern, sector D) accreted crusts, the RMS roughness was estimated as 194 and 154 m, respectively. However, the seafloor roughness as determined from real-time multi-beam data shows a lower degree of crenulations. For example, the rapidly (fast)-accreted crust of sectors A and C recorded an amplitude of the order of 62–78 m, while for the slowly accreted crust of sector D the amplitude of crenulations is 118 m (Table 2.2; Mukhopadhyay et al., 1997). Thus, the real-time seafloor roughness values of the crusts in the IONF are lower than the calibrated theoretical RMS values (Malinverno, 1991). These lower values may have been caused by various factors, such as the ancient oceanic crust tends to get smoothened with time because of subsidence, loss in thermal budget, sediment overburden, erosion and mass wasting. In summary, the roughness of bathymetry in the IONF differs in values, but not in trend from that of the world’s oceans.

2.2. Seafloor faults In the IONF, a number of ridge parallel east-west trending faults are found, each of which appears to be composed of several smaller faults. Two types of faults are encountered in the IONF—those having their planes facing northward away from the spreading ridge axis from which these have been generated, and those having


49

their planes facing southwards towards the ridge axis. The north-facing faults are normal faults. In contrast, the south-facing faults , that is, towards the ridge axis, are considered as ‘reverse faults’. The south-facing reverse faults (generally caused by crustal compression) are probably transformed from original north-facing normal faults (formed by crustal tension). The throw direction of some of these normal faults underwent change owing to a shift in the pattern and magnitude of stress, from tension to compression. In general, the density and the total population (number) of faults increases from north to south. The faults are oriented normally in a direction between N80E and N110E. The length of these faults ranges between a little 50 Ma) IONF seamounts appears similar in terms of distribution, petrology and origin to the younger seamounts formed at the fast-spreading EPR of recent times (age 1000 m). Among these, the major seamount (also the largest one in the entire IONF) is located at 11 090 S and 73 040 E, with its summit height reaching to 1550 m, and has used up about 115 km3 of lava. This left the rest (15 km3) of the lava to form 30 minor and 33 small edifices. This enormous draw of magma by the lone major seamount probably explains the absence of any medium- and large-sized seamounts in the neighbourhood of this fracture zone. A low abundance of seamounts along the Vishnu FZ is, however, inconsistent with the concept that an FZ positively influences seamount formation. It is possible that cool seawater might easily enter through this wide FZ (13- to 17-km) causing a reduction in the pressure and temperature over the magma chamber, thereby diminishing the chances of magma eruption. In addition, the occurrence of thick ferromanganese oxide coatings (6–9-mm) around ancient, depleted, normal-MOR basalt (N-MORB) fragments recovered from this fracture zone would indicate lack of volcanism in the recent past (Mukhopadhyay et al., 1994). Along the Indrani FZ, seamounts are present in the southern part, the summit height rising to 4000-m water depths, while the northern part is conspicuously devoid of seamounts (Kodagali et al., 1992). Again, the seamounts in the western part show a linear trend parallel to the Indrani FZ and tend to be circular, conical and single peaked. Those occurring to the east of this FZ, in contrast, are irregular in shape and multi-peaked. Such variations in the distribution of seamounts across the Indrani FZ may be related to an anomalous thermal input from a nearby source.

60


Abundant seamounts of variable sizes are found close to the NNE-SSW-trending TJT-In (100 20 km on either side of the lineament). In fact, about 70–75% of the IONF seamounts are present in the vicinity of this magmatically active zone (MAZ, Fig. 2.1). Of the total seamounts present in the IONF, about 87% of small and minor seamounts (height 50–499 m), 73% of large (height 500–999 m), and 57% of major (height > 1000 m) seamounts occur within this MAZ. The high abundance and density of seamounts in this zone could be attributed to the excess eruption of magma through faulted, fissured and thin lithosphere that normally occurs in the vicinity of a triple junction. The abundance and distribution of seamounts may thus suggest the possibility of a thermally elevated source at the TJT-In. The seafloor located between the Vishnu FZ and the TJT-In lineament, on the other hand, contains about 133 minor, 90 small and 4 large seamounts that used 244 km3 of lava. This suggests that seamounts can form even in the absence of any structural lineaments (Table 2.5). All these findings prove that the mere presence or absence of a FZ need not necessarily influence the formation and abundance of seamounts and that the magma budget along the FZ/lineament is very important.

3.3. Structural lineaments and seamount morphology Several parameters influence the morphology of seamounts. They include the local and regional tectonic settings, thermal properties of the underlying lithosphere, conduit geometry, magma composition, sediment cover, flow rate of the melt, gravity pull on the ascending magma and magma viscosity. Variations in overall morphology of the IONF seamounts, when seen in this context, reveal that seamounts occurring within 20 km of fracture zone show moderate-to-high basal area, slope angle and volume. These values are, however, less than those occurring between 20 and 50 km across the fracture zone, which generally are the tallest, largest and most conical. Seamounts located between 50 and 100 km from the fracture zone do not show any definite trend but record wide variations in height (100–1201 m), volume (0.4–101 km3), basal area (6–236 km2) and slope angle (2.5 –17.8 , Table 2.5). Although in plan most of these volcanoes are circular to sub-circular, a few show subtle to prominent elongation of the base and the summit. For smaller seamounts, the elongation trend is NW-SE and NE-SW, while for larger edifices the trends are generally N-S and E-W. Flatness (ratio of summit area to basal area) appears to be a pivotal morphological parameter, and it is found that seamounts located away from the fracture zones are more pointed or conical than those occurring close to it. Also, mean flatness of seamounts decreases with increasing height. The taller the seamount, the lower the flatness. Small seamounts located in the MAZ (near TJT-In) particularly show very high flatness values and have slope angles of 8 . Conicality of seamount seems to depend on the amount of erupted lava, conduit geometry, physical state of magma and rate of eruption. For small seamounts, the amount of erupted lava is naturally less, conduits are shallow and narrow and, therefore, a flattened summit is formed. A comparison of height and flatness data between the IONF seamounts and those of the Pacific Ocean (Fig. 2.9) reveals that the bias in seamount distribution is probably related to conduit geometry and eruption style.

61


Eruption of magma through narrow openings is likely to produce tall seamounts with a pointed summit, while small, flattened volcanic swells are formed by eruption through elongated fissures. Principal component analysis of various morphological parameters (both independent and dependent variables) of the IONF seamounts suggests that summit height is the prime responsible factor (first eigenvector) that accounts for 54% of the total morphological variance. This is followed by flatness as the second eigenvector. Among the factors that control overall seamount morphology in the IONF have been the availability of magma and style of eruption (Mukhopadhyay and Khadge, 1992). Similar studies from the North Atlantic seafloor suggest that volume of plume material and conduit geometry may have controlled the formation of seamounts.

3.4. Seamount petrology The rocks dredged from the IONF seamounts are pillow basalts with phenocrysts largely of plagioclase and a few of olivine. These basalts can be broadly divided into three types. The first type shows fresh glass with microlites of plagioclase and a small amount of olivine set in it. Towards the interior, the rock is holocrystalline with abundant plagioclase, forming typical flow and glomeroporphyritic texture.

B

0.1

Volume (km3)

Flatness

A

0.01

100 10 1

0.001 0 200

600 1000 Height (m)

1400

Conical

600 1000 Height (m)

1400

D 0.80

Old edifice, MORB type, Pacific Recent edifice, K2O rich, Pacific Near-axis edifice, MORB type, IONF

Recent intraplate seamount

0.00 0

1000 2000 Height (m)

3000

103 Volume (km3)

Off-axis seamount

Flatness

Truncate

C

0 200

Off-axis seamount, Pacific

102 10 1 10−1 102

Intraplate seamount, Pacific Near-axis seamount, IONF

103 Height (m)

104

Figure 2.9 Growth characteristics of seamounts: (A) and (B) relations among height, flatness and volume in semi-log term for the Indian Ocean Nodule Field (IONF) seamounts, (C) and (D) similar relations in the IONF seamounts are compared to seamounts in the Pacific Ocean to decipher their origin (Mukhopadhyay and Khadge,1990).

62


At places olivine is replaced by iddingsite, while phenocrysts of plagioclase exhibit twinning and zoning. The second type of basalt is partly altered with plagioclase as the dominant mineral, together with olivine and clinopyroxene. The rocks show intergranular, subophitic and intersertal textures. The third type is holohyaline to holocrystalline showing a trachytic texture formed by plagioclase laths. Compositionally, the samples from the IONF seamounts are silica-saturated, show depletion in incompatible elements and appear akin to N-MORB. FeO and TiO2 exhibit an inverse relationship with MgO and Al2O3, while K2O and P2O5 are generally low, and the rocks show high Mg # (62.3; high Mg # is a measure of melt pristinity and consequently represent low degree of melt fractionation). The Mg # is calculated as 100 MgO/Mg þ Feþ2, assuming that Fe2O3/FeO ¼ 0.15. The average composition (in wt%) of these rocks is SiO2 50.6, Al2O3 15, TiO2 1.52, FeO 10.7, MgO 7.34 and CaO 12 (Table 2.6). This primitive nature of seamount lava in the IONF is comparable to the depleted tholeiitic lava from the Pacific and the Atlantic ocean seamounts. It is suggested that the melt that formed the large majority of the IONF seamounts was ancient in nature and probably derived near the ridge crest from a heterogeneous upper mantle silica-saturated source, depleted in large ion lithophile elements. The other circumstantial evidences for the near-axis origin of the IONF seamounts come from the fact that many of the rocks dredged from these seamounts overlie normally polarised oceanic crust, and have a thick coating of ferromanganese oxides accumulated over a long period. In fact, volcanic evolution of any seamount could be ascertained by the type of lava it erupts, with pillow and sheet lava forming the two end members of the evolutionary lineage. In the IONF seamounts, pillow lavas were predominant, indicating that the magma was less evolved, and surfaced Table 2.6 Chemistry of primary and secondary seamount materials from the IONF and other Oceans

SiO2 TiO2 Al2O3 FeO MnO MgO CaO Na2O K2O P2 O 5

a

b

c

d

e

f

50.53 1.27 15.44 9.81 0.19 7.73 12.30 2.80 0.08 0.11

50.60 1.52 15.03 10.70 na 7.34 12.00 na na na

50.10 1.25 16.73 9.20 na 8.13 11.45 2.79 0.06 0.12

49.89 1.36 15.52 9.42 na 7.76 12.40 3.58 0.10 0.13

50.74 2.27 13.42 13.28 0.21 6.20 10.86 2.74 0.09 0.28

48.37 1.63 16.18 8.44 0.16 6.20 10.08 3.04 0.85 0.29

Sources: a, b, e, f ¼ Mukhopadhyay et al. (2002), c ¼ Batiza et al. (1989) and d ¼ Batiza et al. (1990). Note: a ¼ ancient near-axis seamount, IONF, b ¼ ancient near-axis seamount, IONF; c ¼ young near-axis seamount, Mid-Atlantic Ridge; d ¼ young near-axis seamount, East Pacific Rise; e ¼ secondary eruption at the enlarged portion of the seamount, in ‘a’, IONF, f ¼ alkali basalt from lower flank of a seamount, IONF.


63

with a low rate of effusion. The petrography and chemistry of these seamounts also suggest eruption of less fractionated near-pristine magma at or near the then moderate- to fast-spreading IORS. In contrast, seamounts occurring in isolation in the IONF were the products of separate batches of magma rising from small and shallow magma chambers.

3.5. Growth conditions of seamounts The mode and type of emplacement of seamounts through the geological ages are generally deduced from the underlying magnetic anomalies, style of disposition, morphology and chemistry. Magnetisation of a seamount, normal or reversed, may yield constraints on the timing of seamount volcanism, particularly if a seamount is magnetised in the opposite direction to the underlying crust. For instance, if a seamount and the underlying crust display the same magnetisation, the seamount may be considered to have formed in the same magnetic epoch as the underlying crust, and if they are oppositely magnetised, the seamount must have formed during a later magnetic epoch. Normally magnetised seamounts show negative anomalies to the north of the seamounts and positive anomalies to their south. Reversibly magnetised seamounts, in contrast, display positive anomalies to the north of the seamounts and negative anomalies to their south. The rule of thumb has been that if a seamount is oppositely magnetised from the underlying crust and is situated more than 25 km from the nearest younger isochron, the seamount may be considered an off-axis edifice (Barr, 1974). Using this argument, it is found that a majority of the IONF seamounts were ancient in origin, and formed at or near the ridge crest between 60 and 49 Ma (on-axis or near-axis origin during Palaeocene–Eocene). As suggested earlier, the growth patterns and evolution of the IONF seamounts are influenced by the regional tectonic activities and volcanic regime that prevailed during their formation. A majority of these seamounts show different trends of basal elongation, consequent probably to the speed at which the underlying crust was generated. For instance, basal elongation of seamounts generated on the fastaccreting crust in sectors A and C (rate of accretion 90–95 mm/year) strikes 050–230, that is, nearly normal to the then IORS trend. This trend of basal elongation appears to be reconciled to the unhindered and fast northward movement of the Indian Plate, and the resultant overall tensional stress conditions prevailing between 60 and 51 Ma. However, the stress regime changed to compressive type once the collision occurred between India and Eurasia at about 51 Ma. As a consequence of this change in nature of stress, basal elongation of seamounts present in sector D and to its south altered to a nearly ridge-parallel direction 090–270 and 135–315. Some striking similarities between the IONF seamounts and those in the tectonic regime of the Pacific and the Atlantic oceans have been recorded. For example, a close similarity exists in the style of distribution between the small, isolated ancient seamounts occurring near the Vishnu FZ in the IONF and those formed at the slowspreading MAR. Such similarity in distribution style was also maintained between the chains of ancient seamounts occurring close to TJT-In in the IONF and those generated presently at the fast-spreading EPR. Further, the density of the seamounts in the IONF also closely resembles the present near-axis volcanoes in the Pacific

64


Ocean. For example, the average seamount density on the ancient IONF crust (age 60–50 Ma) is 5 to 9 seamounts/103 km2 and that of the young Pacific crust (age tabular. The extrusion rate of lava into the seawater may also determine the morphology of the plagioclase and olivine. For instance, sheaf-like radial clusters could have been caused because of rapid spherulitic growth during quenching of the lava (Natland, 1991) and/or due to disequilibria because of resorption with the melt (Best and Bothner, 1971, Fig. 3.2D). The predominance of plagioclase (olivine) and the plagioclase phyric nature of the IONF basalts imply a process of fractional crystallisation of a plagioclaserich magma, an observation that concurs with the other reported MORB (BVSP, 1981). The paucity of zoned crystals and variation in anorthite content (An34–52) in plagioclase in the IONF basalts may suggest a continued slow growth of plagioclase within the melt during its sluggish ascent to the seafloor (BVSP, 1981; Iyer and Karisiddaiah, 1990). Fresh basaltic glass from a few IONF seamounts shows depletion in incompatible elements and a moderate-to-high Mg # (61.7–63.5). FeO and TiO2 are inversely related to MgO, while K2O and P2O5 are generally low. These compositions are typical of N-MORB and comparable to those from the MAR and the EPR seamounts and from other sites in the Indian Ocean. As in the EPR seamounts, the melts that formed the IONF basalts were derived from a heterogeneous source. Since basalts from the ancient IONF seamounts are nearly indistinguishable from the younger, near-axis originated seamounts of the MAR and the EPR (Mukhopadhyay et al., 1995), it can be reasonably presumed that other seamounts in the IONF also have a similar origin since most of them occur in ridge-normal chains (i.e. normal to the present SEIR). Hence, it is conceivable that most of the seamounts may have tapped a common magma chamber.

1.2. Ferrobasalts Ferrobasalts or FeTi basalts (i.e. enriched in Fe and Ti) are significant because of their paragenetic relation with the intensity of magnetic values of the seafloor, which formed the basis for the concept of magnetic telechemistry (Vogt and Johnson, 1973). Ferrobasalts occur either at propagating rifts, high-amplitude magnetic zones (HAM), fast-spreading ridges or in a combination of all these. Ferrobasalts from the IONF, recovered near topographic highs and HAM zones, possibly formed on the then moderate-to-fast-spreading crust (Iyer et al., 1999a). In the IONF, ferrobasalts are found in two clusters in sector C: cluster I (in the north) and cluster II (in the south) (Fig. 3.3, Table 3.2). The two clusters also show differences in terms of geological settings and petrographic features of the ferrobasalts (Table 3.3). The former ferrobasalts overlie the prominent magnetic anomaly A25 and the crust that was generated at a fast rate (190 mm/yr) during the Late Paleocene– Early Eocene time (56 Ma). Also, multi-beam mapping in this area reveals a prominent east-west-trending zone of elongated topographic elevations, narrow faults and fissures. Rocks were dredged from the lower slope and foothill regions of a large seamount (11 -11 120 S and 77 490 -78 E, summit height 1100 m), which has a wide semicircular base and a gentle slope. Another seamount of moderate dimension (height 500 m) is located about 80 km south-southeast from this cluster.

73

Volcanics

74⬚

76⬚

78⬚

80⬚

82⬚

84⬚

Z Indrani F

11⬚

TJT-In

Vishnu

FZ

72⬚E 9⬚S

13⬚

15⬚

Figure 3.3 Distribution of different types of volcanics in the Indian Ocean Nodule Field (IONF). Solid circles ¼ seamounts; crosses ¼ normal-mid-ocean ridge basalt (N-MORB), altered basalts and ferrobasalts; solid squares ¼ zeolitites; thick solid rectangle ¼ spilites. Dashed outer boundary approximately limits the spread of pumice. Note overlaps of various volcanics at many sites.

Table 3.2

Ferrobasalts in the IONF: Tectonic regime Cluster I

Ridge-dependent characters 1. Magnetic anomaly (A) 25 2. Crustal age (Ma) 56.5 3. Spreading rate (mm/yr) 110 4. Magnetic intensity (nT) 400 to þ500 5. Stress regime Compression 6. Collision phase Feeble stage of phase I Flexuring of surrounding seafloor (m) 7. RMS flexuring 194 8. Flexuring amplitude 67 9. Wavelength of folds 5754

Cluster II

24 53–52 190 200 to þ400 Tension Inter-collision phase 90 78 6454

Source: Iyer et al. (1999a). Note: IONF ¼ Indian Ocean Nodule Field. RMS (root mean square) flexuring or roughness of the seafloor was calculated as roughness ¼ 1296 spreading rate0.539 (Malinverno, 1991); collision phase represents collision between India and Eurasia.

The ferrobasalts occurring in cluster II encompass anomalies A24a, A24b and A25. Three samples in this area were recovered from the underlying crust bounded by magnetic anomalies A24a and A24b (53– 52 Ma), with two occurring on the trace of the triple junction on the Indian Plate (TJT-In; Dyment, 1993), that is, at

74


Table 3.3

Ferrobasalts in the IONF: Geological setting and petrography

Sample no. Location ( S/ E)

Water depth (m)

Cluster I 16

11 000 , 77 550

5250

28

11 000 , 78 000

5300

30

11 120 , 77 490

5400

40

11 050 , 77 500

5300

Cluster II 02

12 330 , 77 120

5210

Seamount cluster (100–800 m height)

06

13 200 , 77 300

5040–5500

Abyssal plain

17

13 000 , 76 300

5400

Abyssal plain þ seamounts (200–1200 m)

Geological setting

Petrographic features

Plagioclase laths dominant, less olivine and opaque minerals; sheaf-like texture, altered Slightly altered Near anomaly plagioclase, A25 at foothill palagonite; of a seamount sheaf-like texture Abyssal plain, near Glass unaltered, plagioclase, A25 at a pyroxene; seamount base intersertal texture Near A25, abyssal Plagioclase, palagonite, plain, away FeMn coating; from sheaf-like seamounts texture, slightly altered Foothill and slope of a large seamount

Plagioclase dominant, olivine rare, opaque minerals, texture hyalopilitic and trachytic Plagioclase, palagonite, opaque minerals, FeMn; medium grained Plagioclase, altered olivine, hematite; sheaf-like texture

x

75

Volcanics

Table 3.3 (continued) Sample no. Location ( S/ E)

Water depth (m)

Geological setting

26

12 330 , 77 120

5200

Seamount cluster (100–800 m)

37

12 590 , 76 300

4900–5050

Seamounts, EW lineaments, near TJT-In

50

12 000 , 76 300

5400

Abyssal plain, 17–35 km from nearest seamounts

Petrographic features

Plagioclase, few olivine, FeMn, trachytic texture Plagioclase dominant, opaque minerals, palagonite; fine grained altered Plagioclase dominant, olivine, opaque minerals; intergranular texture, altered

Source: Iyer et al. (1999a). Note: A25 ¼ Magnetic anomaly 25; TJT-In ¼ Trace of the Triple Junction on the Indian Plate. Fe-Mn oxides ¼ Ferromanganese oxides; IONF ¼ Indian Ocean Nodule Field.

the boundary between the crust generated by the CIR and the SEIR, and one sample from a prominent east-west trending lineament. The other ferrobasalts were from the crust generated between anomalies A24 and A25 (i.e., between 53 and 56 Ma), one of which is from the northern slope of an irregular seamount (height 525 m, basal area 22 km2, volume 4.7 km3, slope angle 13.6 ). Magnetic anomaly data suggest that ferrobasalts in clusters I and II were probably emplaced at 56 Ma and between 52 and 56 Ma, respectively. The two clusters also differ in other ways. For instance, the magnetic anomaly data, when compared with synthetic regional magnetic values, indicate large-scale variations in the intensity of magnetisation of the oceanic crust. The synthetic value of magnetic intensity in A25 (cluster I) is between –400 and þ500 nT while in anomalies A24a and A24b (cluster II) it is –200 to þ100 nT and –150 to þ400 nT (Mukhopadhyay et al., 1997). Mineralogically, plagioclase occurs predominantly as microlites and laths in the ferrobasalts while a few phenocrysts and microphenocrysts of plagioclase are set in a holo- to hypo-hyaline groundmass. Olivine, being rare, is anhedral and sometimes altered to iddingsite. Small euhedral magnetite and reddish patchy hematite mainly represent opaque minerals. The rocks commonly show trachytic texture and, in some cases, the plagioclase laths form sheaf-like aggregates with intervening glass. In few cases, the rocks have a thin top veneer of unaltered glass and/or Fe-Mn oxides.

76


In terms of their petrography (Table 3.3) and chemistry (Table 3.4), the ferrobasalts are noticeably dissimilar to N-MORB but comparable with other oceanic ferrobasalts. The generally low content of MgO (3.24 and 6.14%) reflects the scarcity of olivine while the high MnO (1.99) in sample # 40 is due to the presence of Fe-Mn oxides. Ferrobasalts of cluster II are slightly enriched in SiO2 (45.87–51.67%) and depleted in TiO2 (1.80–2.12%) and FeO* (i.e. total iron, 11.92–17.09%). The lower MgO (1.76–5.71%), as compared to that in cluster I samples, suggests a greater paucity of olivine. In the FeO*/MgO versus TiO2 plot (Fig. 3.4A), two groups are distinct, with most of the IONF and GSC (Galapagos Spreading Centre) samples in one group and a few IONF samples and the ferrobasalts from global oceanic domains in the other group. Note that the K2O contents in the ferrobasalts at the DSDP site 216, Spiess and Chain ridges and Iceland have values similar to those of IONF samples (Table 3.5). It has also been suggested that the high K2O content may not be due only to alteration but may as well be a characteristic of the ferrobasalts (le Roex et al., 1982; Thompson et al., 1978). Considering the trace elements (Table 3.4), the IONF samples have identical Ba but enhanced Y and Zr contents compared to those of the DSDP sites but values lower than those of the Spiess and Chain samples. Vanadium (V) content in these samples is comparable to the DSDP ones while in the Spiess and Chain ridges’ samples it is less by an order of magnitude. A positive correlation between TiO2 and V suggests incorporation of V in TiO2 phases. The incompatible elements (Ba and Y) show systematic increase with differentiation when plotted against Zr, with Ba (Fig. 3.4B) being slightly scattered than Y (Fig. 3.4C) because of the apparent immobility of Y. The low Cr content indicates the role of crystallisation of clinopyroxene and/or titano-magnetite during the evolution of the lava while decrease in Ni could be either due to these phases and/or olivine. The low Ba and high TiO2 (53– 68 ppm and 2–3%, respectively) and enhanced Fe, K, Ti and P probably indicate the alkaline nature of the IONF ferrobasalts akin to the Iceland samples (Thy, 1989). The IONF ferrobasalts are therefore broadly similar to their global counterparts, as further attested by the paucity of olivine, profuse plagioclase and opaque minerals and enrichment in FeO*, TiO2 and trace elements. Several magmatic processes can be conceived to decipher the formation of these rocks (1) crystal fractionation at the shallow level of a MORB melt, depleted in the LREE (Christie and Sinton, 1981; Vogt and Byerly, 1976); (2) a parent magma inherently enriched in Fe and Ti (Klein et al., 1991); (3) a mixed magma (Wilson et al., 1988); (4) a large molten magma chamber, for example, the moderately fractionated ferrobasalts along the fastest-spreading segment of the EPR (75 mm/ year between 13 S and 23 S; Sinton and Detrick, 1992); (5) variable degrees of shallow crystallisation; and (6) differences in the depths of magma generation (Scheidegger, 1973). In most of the processes mentioned above, it is presumed that the commonly evolved magmatic compositions remain essentially constant over millions of years in a near-steady state, maintained by repeated increments of mixing and fractionation. However, the basic questions that arise are, how and under what conditions can magma be retained at shallow depths for long duration while undergoing continuous differentiation? This could be explained by a mechanism of ‘neutral buoyancy

Table 3.4

Major and trace element contents of ferrobasalts in the IONF Cluster I

Cluster II

Sample no.

16

28

30

40

2

6

17

26

37

50

SiO2 TiO2 Al2O3 FeO* MgO CaO Na2O K2O MnO P2 O 5 Total FeO*/TiO2 FeO*/MgO Ba Co Cr Cu Ni Y V Zn Zr

45.49 2.99 15.01 16.22 4.17 10.20 2.67 0.47 0.39 0.58 98.19 5.42 3.89 54.9 77.4 96.8 143.5 122.1 64.4 471.3 111.6 201.0

48.47 2.54 15.77 15.58 3.44 9.13 2.98 0.84 0.25 0.30 99.30 6.13 4.52 – – – – – – – – –

47.68 2.54 14.98 15.11 6.14 9.25 1.94 0.97 0.24 0.18 99.03 5.95 2.46 – – – – – – – – –

43.37 2.76 14.37 19.13 3.50 7.84 3.97 0.96 1.97 nd 97.87 6.93 5.46 – – – – – – – – –

47.08 2.12 17.33 17.09 1.76 8.27 3.60 0.51 0.15 0.19 98.10 8.06 9.71 – – – – – – – – –

45.87 1.80 16.89 13.95 5.71 10.14 3.03 0.78 0.27 nd 98.44 7.75 2.44 53.1 65.7 138.3 69.4 201.9 59.6 342.1 270.7 192.2

49.43 2.00 17.84 13.40 3.02 8.84 3.19 1.07 0.24 0.23 99.26 6.70 4.44 – – – – – – – – –

49.64 1.93 16.34 12.92 3.41 8.50 3.42 1.64 0.31 0.45 95.56 6.69 3.79 68.4 67.5 86.7 140.4 63.5 56.9 352.1 204.5 190.9

51.67 2.10 16.49 11.92 2.84 9.05 3.51 0.79 0.27 0.30 98.67 5.68 4.20 61.8 81.1 151.8 104.0 78.1 59.6 317.4 168.6 181.3

48.83 2.01 17.09 13.45 3.55 9.50 3.50 0.88 0.20 0.24 99.25 6.69 3.79 – – – – – – – – –

Sources: Iyer et al. (1999a) and Mukhopadhyay et al. (2002). Note: FeO* represents total iron. nd ¼ not determined; IONF ¼ Indian Ocean Nodule Field. Major elements in wt%, trace elements in ppm.

78


A

4

TiO2

3 x 2 1

1

3 4 5 FeO*/MgO

170 160

C 70

140

60

120

50

100

Y

Ba

B

2

30

60

20

140 160 180 Zr

200 220

7

40

80

40 100 120

6

10 100

120 140 160 180 Zr

200 220

Figure 3.4 (A) TiO2(%)vs. FeO*/MgO of the Indian Ocean Nodule Field (IONF) ferrobasalts together with those from other global occurrences for comparison. Solid circles and solid triangles ¼ IONF ferrobasalts from clusters I and II, respectively, squares ¼ Galapagos Spreading Centre (GSC); circles with dot ¼ Deep Sea Drilling Project (DSDP); circles with cross ¼ East Pacific Rise (EPR); open triangle ¼ Spiess and Chain ridges; inverted triangle ¼ Iceland. FeO* ¼ total iron; (B) Zr vs. Ba and Zr vs. Y(all in ppm) to show the extent of differentiation in the IONF ferrobasalts. Circles ¼ Spiess and Chain ridges; othersymbols are same as in 3.4A.

zonation’ structure of magma reservoirs, together with the role of fractional crystallisation. The horizon of neutral buoyancy (HNB), defined as that ‘depth interval within which the magma density and the aggregate country rock density are equal’ occurs sublithospherically and has a narrow vertical and a wide lateral extent. Beneath the HNB region, magma ascends because of positive buoyancy and is stabilised at a shallow depth (2–4 km), while above this region the magma descends by negative buoyancy. The HNB of tholeiitic melts thus provides congenial conditions for the long-term stability (over millions of years) of magma reservoirs beneath the crust (Ryan, 1994). The IONF ferrobasalts may have originated when some amount of ascending N-MORB magma was trapped at a shallow depth in the HNB, continuously fractionated and was enriched in Fe and Ti. The recurrent intrusion of magma at the same shallow HNB could possibly have maintained the necessary temperature

Table 3.5

Major and trace element concentration in basalts from the world oceans GSC A

Major elements (wt%) SiO2 50.12 TiO2 3.39 Al2O3 12.09 FeO* 17.01 MgO 4.35 CaO 9.33 Na2O 2.62 K2O 0.17 MnO nd P2O5 0.30 Total 99.38 FeO*/MgO 3.91 FeO*/TiO2 5.02 Trace elements (ppm) Ba Co Cr Cu Ni V Y Zn Zr

– – – – – – – – –

GSC B

49.94 2.09 12.95 13.36 5.84 10.59 2.59 0.15 0.20 0.20 94.82 2.29 6.39 – – – – – – – – –

DSDP 214

48.10 2.35 14.90 14.60 6.45 9.04 2.75 0.37 nd 0.19 98.75 2.26 6.21 45 65 38 na 50 525 26 – 120

DSDP 216

49.50 2.75 13.50 13.80 6.57 8.79 2.57 0.90 nd 0.22 98.60 2.10 5.02 140 53 45 na 44 445 31 – 159

DSDP 254

47.66 2.12 15.37 12.90 8.92 8.53 2.69 0.32 nd 0.24 98.75 1.45 6.08 56 47 469 na 194 342 50 – 156

DSDP 256

50.36 2.36 13.34 13.23 6.58 10.20 2.76 0.25 0.21 0.25 99.54 2.01 5.60 39 36 108 155 93 451 45 – 159

EPR 8oN

Iceland

Spiess Ridge

49.69 2.45 13.19 12.57 6.72 10.22 2.65 0.07 0.20 0.26 98.02 1.87 5.13

46.61 3.99 13.93 14.65 5.55 9.88 3.00 0.74 0.23 0.57 99.15 2.17 3.67

50.74 2.46 14.44 11.53 5.32 9.80 3.71 0.74 0.20 0.39 99.33 1.94 4.67

– – – – – – – – –

– – – – – – – – –

115 38.3 46.17 66.5 32.08 233.5 40.3 114.3 218.7

Chain Ridge

51.89 2.52 14.66 10.58 5.46 9.81 3.12 0.80 0.18 na 99.02 2.64 4.20 162 45 70 60 40 250.5 37 141 221.5

Sources: Anderson et al. (1980), Byerly et al. (1976) and Thompson et al. (1978). Note: Location and other details of the GSC (Galapagos Spreading Centre) and DSDP (Deep Sea Drilling Project) sites: GSC A/B- 00 710 N/85 500 W, depth 2523m, at ridgefault intersection, aphanitic plagioclase, pyroxene; DSDP 214—11 200 N/88 430 E (Ninetyeast Ridge), plagioclase, pyroxene, magnetite, trachytic, olivine absent, vesicular amygdaloidal; DSDP 216—01 280 N/90 120 E (Ninetyeast Ridge), same as DSDP 214; DSDP 254—30 580 S/87 540 E (South tip of the Ninetyeast Ridge), pyroxene, plagioclase, rare olivine, amygdaloidal; DSDP 256—23 270 S/100 460 E (South Wharton Basin), plagioclase, pyroxene, magnetite, strain shadow. EPR ¼ East Pacific Rise. nd ¼ not determined, na ¼ not available.

80


and in turn promoted the long residence time. Again, most of the ferrobasalts are associated with topographic highs (and possibly a thick oceanic crust) and consequently individual magma batches would have had an opportunity to experience a greater degree of differentiation during their ascent. The eruption of the ferrobasalts during chrons A25 and A24 could have been facilitated by intense intra-plate tectonic activities caused by the frequent shifting and propagation of the spreading ridges at the intermediate rate of spreading, particularly during the Indo-Eurasia collision event (56–51 Ma).

1.3. Spilites Spilites are not uncommon on the ocean floor but are less abundant worldwide when compared with tholeiitic basalts. In the IONF, spilites occur near the Indrani FZ (Fig. 3.3), which has a throw of 300 m towards the east (Kamesh Raju et al., 1993) and was probably reactivated in the past (Mukhopadhyay and Khadge, 1992). Fresh as well as highly altered spilitic rocks, sometimes exceeding 25 cm in length and with sparse glass coverage, were recovered from this area. The rocks are fine to medium grained with albitic plagioclase (few as phenocrysts), clinopyroxene and olivine, while chlorite, epidote, hematite, other opaque minerals and smectite are minor components. Typical microlitic intersertal texture of spilites, defined by chlorite, and also porphyritic, intersertal and glomeroporphyritic textures are observed. Some spilites are highly altered, as revealed by very low CaO and high K2O, while the others have lower CaO than N-MORB but are analogous to the Sao Paulo basalts, Atlantic Ocean. On binary plots, the sub-alkaline spilitic affinity is noted (Fig. 3.5). The relatively higher than normal K2O content in the IONF spilites (Table 3.6) could be due to seawater alteration subsequent to spilitisation. Given the fact that spilites were mostly found in the neighbourhood of the reactivated Indrani FZ and some major seamounts, a fault-controlled spilitisation of pillows under the influence of hydrothermal solutions cannot be ruled out.

A

B

ALKALINE

8

Spilitic Field

Mildly alkaline

4

Sp ec t

ru m

Sub-alkaline

us

5

2

eo

THOLEIITIC

Ig n

Na2O + K2O

10 6

0 40

0 50 SiO2

60

0

10 K2O K2O + Na2O

20 ⫻100

Figure 3.5 (A) SiO2 vs. total alkalies (after Schwarzer and Rogers, 1974) to depict the subalkaline affinity of the Indian Ocean Nodule Field (IONF) spilites; (B) spilites of the IONF plotted on the Hughes (1972) diagram.

Table 3.6

Composition (wt%) of spilites in the IONF

Sample no.

1

2

3

4

5

6

7

8

9

10

11

SiO2 Al2O3 TiO2 MnO FeO Fe2O3 MgO CaO Na2O K2O P2 O 5 LOI Total

51.05 14.54 2.55 1.40 nd 14.98 2.10 0.29 4.19 2.39 0.31 5.68 99.48

50.40 14.48 1.95 0.36 1.94 15.46 2.59 5.00 4.31 0.82 0.33 1.84 99.48

46.37 14.75 1.90 0.31 nd 18.10 3.79 7.07 5.08 1.00 0.33 1.22 99.92

46.99 18.67 2.05 0.13 3.08 13.68 2.97 5.89 3.58 0.74 0.32 1.48 99.58

46.38 16.30 1.85 0.20 3.22 12.31 3.46 7.07 6.26 0.84 0.33 1.08 99.30

45.41 16.41 2.05 0.27 nd 17.37 4.48 6.48 4.03 0.25 0.35 2.65 99.75

44.41 16.96 2.05 0.27 nd 17.52 4.59 6.77 4.03 0.49 0.35 2.65 100.09

53.70 14.35 1.65 3.50 nd 12.81 2.54 0.59 3.71 1.94 0.44 4.60 99.83

41.98 13.37 2.00 0.64 nd 21.50 1.56 1.77 6.48 2.34 0.29 8.08 100.01

47.96 17.49 1.85 0.28 3.72 12.18 3.92 5.63 3.21 1.76 0.30 1.50 99.80

50.67 13.69 2.18 0.19 13.05 nd 4.81 5.38 4.61 0.80 0.30 nd 95.68

Sources: Karisiddaiah and Iyer (1992); Fodor et al. (1980). Note: nd ¼ not detected; LOI ¼ Loss on ignition; IONF ¼ Indian Ocean Nodule Field. Samples 1–10 from IONF; Sample 11 ¼ Sao Paolo, MAR.

82


1.4. Pumice Pumice is widely distributed in the oceanic sediments. In the Indian Ocean, drift pumice has been recorded to occur near Sri Lanka, Madagascar and Reunion. In the IONF, however, a large field of pumice covering an area of about 600,000 km2 between 9 S to 20 S and 72 E to 84 E has been identified (Fig. 3.3). Of the 4083 pumice pieces examined from 1925 locations in the IONF, 94% are between 0.1 and 4 cm with only about 1% being larger than 8 cm in size (largest 36-cm long). The pumice is buff, grey, black and brown in colour and oval, lineated, subrounded, rounded and irregular in shape (Fig. 3.6). Many of the clasts form substrate and are coated either fully or partially by ferromanganese oxides. Pumice shows silicic glassy webs with a few phenocrysts of plagioclase and/or pyroxene, and vesicles that are mainly lineated, with a few rounded. In some instances, quartz, radiolarians, phytoliths, diatoms and ferromanganese micronodules occur within the glassy interstices. The presence of such materials could make the pumice heavy and cause it to sink to the ocean floor. Compositionally, the IONF pumice is of two kinds: (1) trachyandesite (type 1, SiO2 60%) and (2) rhyodacite (type 2, SiO2 70%). The former shows low K2O and nearly constant Na2O as compared to the latter (Table 3.7, Fig. 3.7). The presence of pumice in the IONF may be due to: (1) transportation from the 1883 eruption of Krakatoa volcano (Martin-Barajas and Lallier-Verges, 1993; Mudholkar and Fujii, 1995); (2) in situ eruption on the seafloor (Iyer and Sudhakar, 1993a; Svalnov, 1981); or (3) a combination of (1) and (2) (Mukhopadhyay et al., 2002). Multi-beam maps display a few of the IONF seamounts with summit craters and their occurrence proximal to the spatially distributed pumice may suggest a possible relationship between explosive silicic eruption and the formation of pumice. Experimental studies have shown that the hot pumice sinks faster and nearer to the source as compared to colder pumice (Whitham and Sparks, 1986). The hot pumice absorbs water, becomes heavier and sinks. In the light of these observations, it can be argued that much of the larger pumice clasts in the IONF might have been derived as hot clasts that rapidly sank near their eruptive sources. Although pumice has a tendency to drift, the frequent occurrence of pumice in the IONF and the absence of any favourable surface current circulation pattern that can transport pumice from the IVA or Toba volcanoes seem to favour an in situ origin of the IONF pumice. Evidences that pumice-forming eruptions can occur at great water depth (1500 m or more) come from all oceans, for example, pumice at the plate boundaries (near Tonga Trench, Okinawa Trough, South of Japan) and at intra-plate areas (over Atlantic seamounts in the Atlantic Ocean and the Ninetyeast Ridge), and trachytic pumice flow from intra-plate volcanoes in the Society and Austral hot spot regions. Collectively, all these occurrences suggest that widely distributed pumice, in fact, is a product of recent submarine eruptions. In the light of these findings, an in situ origin of the IONF pumice at abyssal depth could be a possibility. It is also pertinent to note that pumice forms substrate and nucleus for many manganese crusts. Since the oxides accrete at a rate of 1 mm/Ma, it would take about 10 Ma to form a 1-cm thick layer around a nucleus of pumice. This deduction indirectly testifies the coated pumice to be older than the volcanic eruptions of Krakatoa in 1883 and that of Toba 74 ka years ago (Mukherjee and Iyer, 1999).

83

Volcanics

A

0 1 2 3 4 5 cm

B

0

1

2

3

4

5

cm C

0

1

2

3

4

5

cm

Figure 3.6 Hand specimen of pumice from the Indian Ocean Nodule Field (IONF). (A) pumice of different shapes and sizes; (B) pumice coated either partially or fully by ferromanganese oxide; (C) an example of pumice specimens from one sampled site showing recovery of coated and non-coated types.

It has been found that fractional crystallisation of abyssal tholeiites could lead to layered or stratified magma chambers with alkali- and volatile-enriched silicic melts (Hawkins, 1985). A similar situation can be visualised for the IONF pumice since the basinal area is dominated by tholeiitic basalts, which in the course of time may have fractionated to a more silicic melt. The presence of basaltic and silicic volcanics

Table 3.7

Composition (wt%) of the CIOB pumice

Sample

1

2

3

4

5

6

7

Avg.

8

9

10

11

Avg.

SiO2 TiO2 Al2O3 FeO* MnO MgO CaO Na2O K2O P2 O 5 LOI

59.48 0.51 13.18 7.73 0.20 2.37 5.79 7.05 3.56 0.10 6.23

56.71 0.73 13.25 9.53 0.17 3.37 6.49 6.23 3.33 0.13 12.50

59.58 0.66 13.29 8.19 0.28 2.38 4.46 5.89 5.14 0.14 15.10

61.42 0.46 12.46 5.91 1.34 2.17 5.57 5.87 4.71 0.10 11.20

60.30 0.54 12.44 8.33 0.64 2.55 5.16 5.81 4.10 0.14 15.00

60.59 0.63 13.22 8.71 0.58 2.51 6.12 3.95 3.57 0.12 3.00

63.91 0.97 13.78 6.75 0.21 2.02 2.34 6.63 3.09 0.30 3.03

60.28 0.64 13.09 7.88 0.49 2.48 5.13 5.92 3.93 0.15 9.43

68.57 0.17 10.97 3.95 0.28 1.11 3.31 5.10 6.50 0.05 12.30

73.75 0.28 13.05 1.74 0.13 0.46 2.33 3.96 4.23 0.08 5.61

69.20 0.97 14.03 3.76 0.41 1.33 2.52 4.84 2.75 0.17 12.73

69.52 0.77 14.24 5.51 0.35 1.97 0.56 3.81 3.04 0.21 6.50

70.26 0.55 13.09 3.74 0.29 1.22 2.18 4.43 4.13 0.13 9.28

Source: Iyer (1995). Note: LOI ¼ Loss on Ignition; Avg. ¼ Average; FeO* ¼ Total iron as ferrous iron; CIOB ¼ Central Indian Ocean Basin. For groups see text. Samples 1–7 ¼ Trachyandesite; Samples 8–11 ¼ Rhyodacite.

85

Volcanics

8

Na2O

4

4

0

0

0.8

8

K2O

TiO2

FeO

8

0.4

4

0

0 22

8

MgO

Al2O3

18

14

4

0

10 80

72

CaO

SiO2

8

64

56

4

0 4

8

12 SI

16

4

8

12

16

SI

Figure 3.7 Variation diagrams for the CIOB pumice. SI = solidification index (100MgO/ MgOþNa2OþK2OþFeO*); filled circles ¼ trachyandesite pumice; open circles ¼ rhyodacitic pumice; open triangles ¼ data from Frick and Kent (1984); solid triangles ¼ data from Martin-Barajas and Lallier-Verges (1993).

86


may point to a process of simple fractionation of the basaltic magmas to form the pumice. Hence, it is plausible that the IONF pumice and glass shards may have jointly been contributed from the cataclysmic Krakatoa eruption of 1883, Toba eruption at 74 ka and from in situ volcanism (i.e., local mid-plate basinal eruptions).

2. Minor Volcanics 2.1. Hydrothermal materials Volcanic-hydrothermal materials (vhm), discovered in the IONF (Iyer et al., 1997a,b, 1999b), have important implications on the structural and volcanic histories of the field. Coarse fractions from surface and sediment column from various sedimentary regimes were investigated for vhm. Two occurrences of vhm have been reported, one from the base of a seamount, about 45 km from the TJT-In in the siliceous sediment domain; the other scooped up from a red clay sediment domain (Fig. 3.8). The seamount, located at a water depth of 4440 m and overlying magnetic anomaly A 23b (¼age 50.81– 50.64 Ma), is 800-m high, covers an area of 37.56 km2 (length ¼ 5.77 km, basal width ¼ 6.51 km) and has a summit width of 0.18 km. The coarse fraction contains 2200–2500 particles of vhm per gram. The vhm from the red clay sediment domain, however, contains only 25 spherules per gram. This location is about 53 km from a major seamount. Individual entities of native Al spherules and particles occur associated with the vhm in the vicinity of seamounts (Iyer et al., 2007a). Radiolarians and diatoms were identified from the various subsamples of the vhm-hosted sediments so as to determine a possible age. Two radiolarian zones, Buccinosphaera invaginata and Collosphaera tuberosa, are present with a distinct boundary between them at 30 cm down core. This defines the first appearance datum (FAD) level of B. invaginata (Gupta, 1988) and are paleomagnetically dated to be synchronous at 180 20 ka in nearby cores (RC 14–22 and VM 34–53, Caulet et al., 1993). Factor analysis of percentage data of 47 species in the IONF helped to recognise three conspicuous radiolarian assemblages, which were related to the overlying sea surface temperature (Gupta, 1996). The assemblages were quantified, and mutual ratios were plotted (Fig. 3.9) to identify datum within the latest Quaternary period. In the vhm from a siliceous sedimentary domain, the ratios of transitional to warm and transitional to cold fauna show three prominent peaks, which, on the Pleistocene climatic scale of Martinson et al. (1987), correspond to 130, 70 and 10 ka ages for the 28, 16 and 6 cm (dominant vhm) depth, respectively. Cacsinodiscus nodulifer diatoms are generally absent in most of the upper sections of the core but occur abundantly between 20 and 28 cm and at 36 cm core depths (Fig. 3.9). Burckle and McLaughlin (1977) found that this species (>120 mm) proliferate in warmer climates, suggesting the signature of the last interglacial age (130 ka) between 20 and 28 cm in this area. In contrast, the vhm from red clay domain show absence of the index fossil B. invaginata (FAD level at 180 ka; Gupta, 1988), but presence of a number of Stylatractus universus (extinction ¼ Last Appearance Datum level at 425 ka;

87

Volcanics

30⬚ N

5⬚ S

F/98/3

10⬚ N

F/99/3

10⬚ S

IONF

40⬚E

F/101/0

F/154/0

S/120/1 S/128/1

Indrani FZ

TJT-In

Z S/106/0 S/124/2

80⬚

100⬚

F/156/4

F/155/1 F/152/0 F/153/0

Vishnu F

10⬚

60⬚

S/139/0

S/129/0 F/81/2

S/120/2 S/126/2 F/56/1 15⬚

S/231/1

* S/89/2200 S/657

S/206/1 S/210/5

S/241/2

S/94 20⬚ 70⬚E

75⬚

80⬚

85⬚

Figure 3.8 Distribution of volcanogenic sediments in the Indian Ocean Nodule Field (IONF). Open stars ¼ seamounts; solid dots ¼ sediments examined for volcanogenic hydrothermal materials (vhm); solid triangle ¼ S657 and asterik ¼ S89 near a large seamount; circle with dot ¼ S94, the southernmost sample with vhm.The last numerals after station number (e.g. S/124 and S/ 210/5) represent the number of magnetic particles present in 1 g of the sediment coarse fraction.

Hayes and Shackleton, 1976) and Collosphaera Orthoconus (FAD level at 650 ka; Johnson et al., 1989). From these two dates, the age of the vhm in the IONF seems to range between 425 and 650 ka. Older sediment exposures at this site, as confirmed by the presence of Australasian microtektites of 770 ka age (Prasad, 1994), suggest intensive scouring by local bottom currents. This is also evident by the presence of 1–2 cm-sized spherical manganese nodules that testify to periods of no sedimentation or erosion (Iyer, 2005). Generally, three types of vhm occur in the IONF: (1) ochrous metalliferous sediment (2) magnetite spherule and (3) glass shard (Iyer, 2005).

88


A 1

Ratios 2

3

0 Sample depth in core (cm)

4

10 ka

8 12 70 ka

16 20 24

130 ka

28

180 ka

32 36

Transitional v/s warm fauna

FAD of B. invaginta Transitional v/s cold fauna

B 0

Glass shards in 1000/slide 2.5 5.0

0

5

Magnetics ( ⫻ 100) 10 15 20 25 30

35

40

0 4 8 12 16 20 24 28 32 36 0

50 100 No. of diatoms/slide

Figure 3.9 Down-core variation in sediment core S657 to show (A) the occurrence of radiolarian species and ratios of transitional to warm water fauna (solid triangle) and transitional to cold fauna (open circle); FAD = First Appearance Datum and (B) the occurrence of glass shards per slide (hatched portion) and diatoms per slide (dotted portion). The panel on the right shows the number of magnetic particles.

2.1.1. Ochrous metalliferous sediment The magnetic fractions in the vhm consist of highly porous to indurated material, interpreted as ‘ochrous precipitates’ of hydrothermal origin, that displays moss-like pyroclastic texture (Fig. 3.10A). The smallest precipitate is only a few microns in diameter and the largest, 3.35 mm 1.45 mm, is oval shaped, weighs 30 mg, and has a depressed narrow end and overturned rims that formed possibly because of melt flow. The precipitate also hosts a bunch of spherules and has glass shards

89

Volcanics

A

C

B

D

Figure 3.10 Scanning electron microscope (SEM) images of the ochrous sediment particles and spherulesç(A) similar to moss-like morphology of Heiken and Wohletz (1985); scale bar = 10 mm: (B) condensed particle and spherules similar to the experimentally produced ones of Wohletz and McQueen (1984); scale bar = 100 mm; (C) nontronite showing striations recovered in core S94; scale bar = 100 mm; and (D) magnetite [spherules] and titanium [irregular] particles in core S94; scale bar = 10 mm.

protruding from it. The ochrous precipitates exhibit various properties and hues, from highly porous, light yellow and orange to more condensed variety (Fig. 3.10B). The porous ochrous precipitates do not have any distinctive mineralogy vis-a`-vis the condensed particles with magnetite and maghemite as the main phases. Compositionally, the metalliferous sediments are a mixture of SiO2 (16–30%) and FeO (54–73%) with variable contents of other oxides (Table 3.8). The inclusions of Ni and S in the metalliferous sediment are generally between 0.02 and 1.40% and 0.21 and 4.65%, respectively, with the highest Ni (up to 2.24%) and S contents (up to 9.82%) detected as inclusions in one place (Table 3.9). The four possible modes of formation of metalliferous sediments are (1) derivation from hydrothermal emanations (Zelenov, 1965), (2) reaction of seawater with hot lava (Corliss, 1971; Honnorez, 1981), (3) hydrothermal precipitation (Alt et al., 1987; Hekinian et al., 1993) and (4) contributions due to bacterial processes (Alt, 1988). Interestingly, the IONF samples resemble Fe-Si oxyhydroxides of the Pacific Ocean (Si/Fe ratio 0.14–0.33), and are also comparable to Fe-oxide mud of inactive deposits of the Red Seamount in the Pacific Ocean (Table 3.8). It therefore appears that

90 Table 3.8

Microprobe analyses (wt%) of metalliferous sediments in the IONF

Serial no.

Samples

SiO2

TiO2

Al2O3

FeO

1 2 3 4

89 89 89 89

16.31 16.18 29.99 25.87

0.33 0.46 0.14 0.17

3.02 3.23 6.62 4.06

71.01 69.32 53.87 58.79

1.35 1.19 0.37 0.13

0.65 0.63 0.05 0.19

0.11 0.30 0.30 0.03

0.28 0.13 – –

97.30 94.82 91.80 90.93

5

89

25.75

0.04

5.07

60.49

0.40 0.46 1.21 0.25 0.20

0.09

–

93.96

6

89

26.11

0.10

4.70

58.56

0.18 1.53 2.52 0.51 0.73

0.01

0.34 95.29

7

89

25.02

0.02

6.43

58.59

0.27 1.49 1.57 0.40 0.79

0.29

0.35 95.22

8 9 10 11

657 657 657 EPR

20.77 22.86 21.93 31.67

0.10 0.05 – –

0.34 2.25 0.23 –

71.06 59.26 73.50 41.20

0.08 – 0.29 0.18

0.19 1.38 – –

0.11 0.20 0.07 –

MnO MgO CaO Na2O K2O

1.66 1.63 0.28 0.46

1.13 6.15 1.12 2.45

0.92 1.01 0.14 0.98

1.59 6.29 1.18 2.37

1.66 0.74 0.04 0.25

0.05 0.28 0.16 1.23

0.26 0.23 0.20 0.52

Cr2O3 BaO Total

95.68 98.95 99.31 79.62

Remarks

T-1, Si/Fe ¼ 0.14 T-1, Si/Fe ¼ 0.14 T-3, Si/Fe ¼ 0.33 T-1/3, Si/Fe ¼ 0.26 T-1/3, Si/Fe ¼ 0.26 T-1/3, Si/Fe ¼ 0.27 T-1/3, Si/Fe ¼ 0.26 T-1, Si/Fe ¼ 0.18 T-1, Si/Fe ¼ 0.23 T-1, Si/Fe ¼ 0.18 Fe hydroxide from seamounts, T-3, Si/Fe ¼ 0.46

12

Society

17.24

–

6.84

56.22

0.03

–

–

–

–

0.02

–

80.35

13

EPR

41.82

–

1.06

26.63

0.65

–

–

–

–

–

–

70.16

14

Red 8.25 Seamount

0.04

0.50

56.06

0.01 0.68 0.86 0.67 0.26

0.002 0.005 67.34

15

Red 17.27 Seamount

0.05

0.85

56.45

0.18 0.90 1.04 0.86 0.39

0.002

16

Galapagos

47.06

–

0.18

32.72

0.31 2.44 0.74 1.54 1.78

–

17

Fukujin

53.91

0.41

0.71

26.65

1.04 1.88 0.84 2.02 1.50

–

18

Red 47.50 Seamount

0.04

0.56

32.16

0.02 1.92 0.09 3.16 1.90

0.03 78.02

–

86.77

0.04 89.00

0.001 0.005 87.36

Fe hydroxide. T-1, S ¼ 0.3%; Si/Fe ¼ 0.18 Fe-Si hydroxide. T-3, S ¼ 0.4%; Si/Fe ¼ 0.95 Fe-oxide mud of active vent, Pacific Ocean (3),Si/Fe ¼ 0.09 Fe-oxide mud of inactive deposit, Pacific Ocean (3), Si/Fe ¼ 0.19 Nontronite (3), Si/Fe ¼ 0.86 Nontronite, Si/Fe ¼ 1.22 Nontronite (2), Si/Fe ¼ 0.89

Note: EPR ¼ East Pacific Rise; IONF ¼ Indian Ocean Nodule Field; 1–10, S89 and S657; Iyer et al. (1997a); 11, Alt et al. (1987); 12,13 and type T-1, T-3, T-1/3, Hekinian et al. (1993); 14,15 and 18, Alt (1988); 16, Corliss et al. (1978); 17, McMurty et al. (1983). Serial number 11 to 15 represents Fe-Si oxyhydroxides and 16–18 Pacific Ocean nontronite. The number of analyes is in parenthesis.

91

92


Table 3.9 Microprobe analyses (wt%) for Nickel and Sulphur Inclusions in Volcanic Magnetite Spherules and Metalliferous Sediments in the IONF Serial no.

Sample no.

Nickel

Inclusions in volcanic magnetite spherules 1 89 0.13 2 89 0.02 3 89 0.11 4 89 0.14 5 89 – 6 89 – 7 657 0.11 Inclusions in metalliferous sediments 8 89 0.02 9 89 0.06 10 89 0.03 11 89 0.08 12 657 1.40 13 657 1.09

Sulphur

Remarks

0.08 0.04 0.06 – 0.01 1.55 0.54

Metallic matrix Oval-shaped spherule Tear drop spherule (3) Matrix Matrix Edge of matrix Matrix (2)

4.65 0.53 0.23 0.21 0.37 0.43

Metallic (3) Metallic (2) Metallic Metallic (2) Metallic (2) Metallic (4)

Source: Iyer et al. (1997a). Note: Serial number 8 and 13 show the highest S (9.82%) and Ni (2.24%) contents at a single probe spot, respectively. The number of analyses is in parentheses. IONF ¼ Indian Ocean Nodule Field.

favourable conditions exist in the IONF for hydrothermal solutions to percolate through the basaltic crust, leach the rocks and mix with the seawater for subsequent cooling and oxidation, resulting in FeOOH precipitates. It is possible that during the discharge of hydrothermal solutions from the igneous crust into the sediments, metal deposits formed either within the sediment column or at the sediment–water interface, resulting in the formation of intra-sedimentary hydrothermal deposits (Bonatti, 1983). The Fe-rich metalliferous sediments in the IONF are sometimes associated with nontronite-like particles (Fig. 3.10C). Nontronite, a typical hydrothermal mineral, shows similar Si and Fe enrichment but has higher than normal Al2O3 and MgO content as compared to other oceanic occurrences. As hydrothermal solution percolates through the sediment, it cools slowly under reducing conditions to form Fe and Si precipitates that later react with biogenic silica to form nontronite (Murnane and Clague, 1983). The composition of the IONF metalliferous sediments and their geologic setting affirm their derivation as in situ precipitates from hydrothermal exhalations (cf. Harrison and Peterson, 1965). Restricted occurrence of these sediments near seamounts and absence of any strong bottom currents in the IONF, as corroborated by several hundreds of underwater photographs, also support their in situ origin. 2.1.2. Magnetite Spherules The volcanic spherules found in the metalliferous sediments are fresh and display metallic luster (Fig. 3.10D). They range in size from a few microns to 475 mm and are mostly spherical, but dumb-bell, tear-drop and oval-shaped spherules also occur, albeit in lower proportions. Many of the spherules have a reddish-brown ochrous

93

Volcanics

coating (Fig. 3.11A). Electron microscopy of the spherules shows the predominant presence of magnetite that forms an array of surface textures, such as brickwork, interwoven, corkscrew type and dendritic (Fig. 3.11B–D). Some spherules have a smooth surface with no apparent crystallinity, while others are polytextured. The textures are caused by quenching, that is, melting followed by immediate supercooling, a process commonly experienced by cosmic dust. In contrast to cosmic spherules, the crystalline structure of the IONF spherules is restricted to the outer zone. In a few spherules, Si-rich blobs either protrude from an A

B

C

D

E

Figure 3.11 Photomicrograph of volcanic magnetite spherules from core S657: (A) adhering ochrous sediments on a few of the spherules; scale bar = 50 mm; (B) well-developed large euhedral magnetite crystals on a spherule; scale bar = 10 mm; (C) quench texture on a magnetite spherule; scale bar = 100 mm; (D) a broken magnetite spherule with linearly arranged crystal edge; scale bar = 10 mm; and (E) polished sections of spherules showing a central Si-rich capsule with a thick wall and enclosing magnetite crystals; scale bar = 100 mm.

94


otherwise Fe-rich matrix or appear as inclusions (Fig. 3.11E). Vacuoles present within the spherules are located concentrically or eccentrically (Fig. 3.11C) and may have formed because of blowout made by escaping volatiles. Magnetite is ubiquitous in these spherules, together with minor proportions of ilmenite, hematite and maghemite (Table 3.10). In core S89, Ti-rich spherules and grains were also recovered (Fig. 3.10D). The spherules are non-welded, probably due to their low population per unit area or due to the rapid formation of a thin, rigid skin on the surface (cf. Heiken and Lofgren, 1971). Except for their sizes, the IONF spherules are similar to the archetypal terrestrial volcanic spherules, and to the magnetite spherules described from the 45 N MAR. The reported size of terrestrial volcanic spherules from different volcanic areas is between 1 and 300 mm (Iyer et al., 1997b) vis-a`-vis the maximum diameter of IONF spherules (475 mm). Microprobe analysis of spherules (Table 3.11) shows high FeO (87–99%; average 95%, Fe 74%) and minor amounts of Ti and Mn, while sectioned spherules show inclusions enhanced in Mg and Al. One of these is hyalosideritic olivine with no crystallinity, while others show orthopyroxene-like composition (bronzite-enstatite). Besides these, the inclusions enriched in Si, Al, Fe, Ca and Na indicate the presence of feldspar but the Fe content is too high for a pure plagioclase. A few inclusions in the spherules are substantially enriched in Ti (52–60%), Ni (0.02–0.14%) and S (0.01– 1.55%, Tables 3.11 and 3.12). The spherules from a red clay domain display Fe content of 65 to 78% (average 76%), while Ni is between 0.03 and 4.46%, and S between 0.02 and 0.17%. Besides these, Ti-rich spherules and inclusions also occur that could represent either ilmenite or titaniferous-magnetite. Different sources can be ascribed for the formation of magnetite spherules (Iyer et al., 1997b and references therein) such as (1) industrial, (2) diagenetic, (3) biological, (4) extraterrestrial and (5) volcanic. An assessment of the composition of the IONF spherules rules out the first three options as sources, as does their likeness to droplets of basaltic glass of Kilauea lki, Hawaii, basaltic hydromagmatic ash, basaltic microspherules from the Eastern Pacific Ocean, and microlapilli occurring near the Vityaz FZ, Indian Ocean. Moreover, an extraterrestrial source is unlikely since the IONF spherules lack a Ni-rich core, a common feature of extraterrestrial spherules (Blanchard et al., 1980). On the contrary, the dominance of magnetite associated with silicate fragments and the presence of magmaphile and diagnostic elements like Ti and Mn in the IONF spherules suggest that they are volcanogenic (Iyer et al., 1997a,b). To understand the mechanism of the formation of volcanic spherules the ‘fuelcoolant interaction (FCI)’ process, in which a rapid vaporisation of water occurs when in contact with hot molten materials, was advocated (e.g. Peckover et al., 1973; Wohletz et al., 1995). When magma contacts water, it is essential that the two mix thoroughly to create vapour explosions and the resulting products fragment before explosion. Fragmentation increases the contact area of melt and water causing heat exchange at explosive rates. To study the formation of spherules, Wohletz and McQueen (1984) made laboratory-scale experiments using thermite melt as an equivalent to basalt. Ignition of a mixture of thermite (Al þ Fe3O4), quartzofeldspathic materials and water in steel containers resulted in the formation of

95

Volcanics

Table 3.10

Gandolifi X-ray diffractometry of the magnetite particles in the IONF

Sample

Size (mm)

S89

275

S89

S657

410

800

Description

Black, metallic spherule

Black, metallic spherule with orange coloured adhered material

Black, metallic botryoidal particle with quench droplets

Exposure time (hours) dA

Intensity

Minerals present

32

2.9827

MS

Ilmenite

S MS W MS MS MS MS MS W W

Magnetite

80

2.7300 2.5363 2.2306 2.1204 1.8596 1.7241 1.6313 1.5063 1.4777 3.6792

VS VS MS VW MS S W W W W

Magnetite

80

2.6861 2.5130 2.1905 2.0891 1.8409 1.6926 1.6099 1.4873 1.4554 2.9307

2.6968 2.5050 2.1188 2.0814 1.6988 1.6099 1.5126 1.4821 1.1159 1.0866

VW VS W W W S VW S W MS

Hematite

Hematite

Maghemite

(continued)

96


Table 3.10 (continued)

Sample

Size (mm)

Description

Exposure time (hours) dA

1.0642 0.9685 0.9548 0.9345

Intensity

Minerals present

W W VW W

Source: Iyer et al. (1997b). Note: VS ¼ very strong; S ¼ strong; MS ¼ medium strong; VW ¼ very weak; W ¼ weak; IONF ¼ Indian Ocean Nodule Field.

hydrovolcanic particles of 1–2 mm after weak interaction, and 1–100 mm following strong explosive interaction. The resultant quenched particles were irregular aggregates, spheriods or blocky conforming to hydrovolcanic ash. Iron and aluminum oxides were formed that resemble a basaltic melt (Table 3.13). Although the compositions are still very different from natural silicate magmas, their similar density, liquidus temperature and viscosity were considered to be viable working analogues (Iyer et al., 1997b; Wohletz et al., 1995). Considering congenial conditions represented by abundant siliceous sediment, predominance of tholeiitic basalt and evidence of hydrothermal activities, a natural field set-up in the IONF seems to have prevailed, similar to the above-mentioned laboratory-scale experiments. It is possible that lavas or hydrothermal emanations enriched in iron might have got admixed with the siliceous sediments resulting in localised hydrovolcanic events(s) and the production of the spherules. The locations where volcanic spherules are found in the IONF overlie major magnetic anomalies close to seamounts. As discussed in the section on ferrobasalts, enrichment of Fe and Ti in the IONF spherules can be related to high remnant magnetisations influenced by strong magnetic anomalies. Iyer et al. (1997a,b) explained the chemogenesis (i.e. the relative enrichment and depletion of Fe and Si) of the IONF spherules based on a process of liquid immiscibility coupled with oxygen fugacity (fO2). It was suggested that oxygen controls the fractionation of magma and results in Fe-rich and Fe-depleted liquids, which on cooling precipitate magnetite (Osborn, 1979). As the temperature of the cooling liquid falls, iron oxidises faster and combines with oxygen to form magnetite while silica forms a glassy microlayer for the Fe-rich liquid cooling in the interior of the spherules, where euhedral magnetite crystallises. With a more complete oxidation, unmixing of the phases occurs so that Si is either pushed aside or else is separated from the Fe matrix as blobs (Del Monte et al., 1974). The observed internal structure and the presence of vesicles in the IONF spherules may be better explained by a mechanism of cooling and solidification resulting from an abrupt decrease in the solubility of oxygen in the oxidised particles, leading to oxygen concentration in the inner, still-molten region of the particle.

Table 3.11

#

Microprobe analyses (wt%) of volcanic magnetite spherules, and inclusions within the spherules in the IONF

Sample no.

SiO2

Magnetite spherules 1 89 and 1.91 657

TiO2

Al2O3

FeO

MnO

MgO

CaO

Na2O

K2 O

Cr2O3 BaO

Total

Remarks

Magnetite spherules (24) Ti-rich areas in spherules (5) Ti-rich areas in spherules (2)

0.60

0.68

94.98

0.46

1.04

0.45

0.39

0.08

0.16

0.13

100.88

2

89

15.73

52.12

5.30

12.32

6.41

1.43

2.45

0.50

2.66

0.63

0.76

100.31

3

89

2.81

60.76

4.41

23.92

6.82

0.40

0.65

–

0.09

0.03

1.04

100.93

Inclusions in spherules 4 89 38.68

0.26

1.76

35.48

0.18

24.02

0.40

0.14

0.08

0.18

–

101.18

5

89

57.71

0.20

3.26

14.23

0.07

23.42

0.23

0.10

0.45

0.24

0.11

100.02

6

657

59.32

–

2.00

9.15

–

28.76

0.08

0.03

0.02

0.60

0.03

99.99

7 8 9

89 89 89

43.26 59.85 48.14

– 0.27 0.29

36.76 25.76 25.18

8.88 2.60 21.25

0.38 – 0.28

0.51 0.05 0.50

4.58 3.55 0.74

5.23 6.40 0.48

0.23 0.93 1.53

0.19 – 0.08

– – 0.29

100.02 94.41 98.76

Source: Iyer et al. (1997a). Note: Figures within brackets are the number of grains analysed.

Hyalosideritic olivine Bronziteenstatite Bronziteenstatite Feldspar Feldspar Feldspar

Table 3.12

Microprobe analyses (wt%) of volcanic magnetite spherules and hydrothermal precipitates from the IONF

98

Specimen #

Si

Ti

Al

Fe

Mn

Mg

Ca

Na

K

Cr

Ba

Ni

S

1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 18 19 20 21 22 23 24 25 26

0.34 1.60 0.17 0.20 0.01 0.07 nd 0.09 nd 0.27 0.02 0.07 0.15 0.16 0.13 0.15 0.08 0.25 0.89 13.38 26.37 11.84 15.03 4.90 1.72 1.93

0.25 0.56 0.02 nd nd 0.03 nd 0.04 0.05 0.13 nd 0.03 0.07 nd 0.05 0.09 0.41 nd 0.36 0.26 nd nd 0.07 0.11 0.07 0.06

nd 4.49 nd 0.03 nd nd nd nd nd nd 0.01 nd 0.66 nd 2.10 nd nd nd 0.36 3.81 13.07 0.98 8.37 35.30 8.26 23.23

74.8 64.9 75.9 75.2 77.6 77.1 76.8 77.1 77.8 76.2 77.3 78.6 76.4 77.4 74.4 77.3 76.9 76.7 73.8 30.9 nd 51.6 37.0 11.2 61.4 26.5

0.43 nd 1.32 1.47 0.08 0.33 0.15 0.37 0.03 0.64 0.03 0.18 0.05 nd 0.04 nd nd 0.01 0.36 0.39 0.19 0.09 0.55 0.10 0.07 0.03

nd 1.21 nd nd nd nd nd nd nd nd nd 0.03 nd nd nd nd nd nd 0.63 11.59 0.75 1.24 0.07 0.27 0.15 0.38

nd 0.43 0.13 0.02 0.01 nd 0.04 0.07 nd nd 0.01 nd nd 0.03 nd 0.04 0.05 0.39 0.32 0.31 6.03 1.77 0.18 0.65 0.26 7.40

0.07 0.19 0.02 0.09 nd 0.01 nd 0.04 0.02 nd 0.07 0.01 nd 0.04 nd nd 0.04 0.05 0.29 0.11 6.22 0.48 0.22 0.18 0.07 0.05

0.02 0.18 0.07 0.02 0.04 0.03 0.05 0.07 0.05 0.07 nd 0.07 nd nd nd nd nd nd 0.07 0.14 0.47 0.82 0.07 0.15 0.13 0.12

0.01 0.11 nd nd nd nd 0.16 0.11 0.34 0.06 nd ndd ndd ndd ndd ndd ndd ndd 0.11 1.38 ndd 0.06 1.80 4.21 0.64 4.32

0.01 0.18 nd 0.34 nd 0.16 0.02 0.01 0.10 0.18 nd ndd ndd ndd ndd ndd ndd ndd 0.12 ndd ndd 0.06 0.15 0.17 ndd 0.22

0.03 nd 3.27 nd 0.04 nd 0.04 nd 4.46 0.07 0.09 ndd ndd ndd ndd ndd ndd ndd ndd ndd ndd ndd ndd ndd ndd ndd

0.07 nd 0.17 nd nd nd 0.07 0.05 nd 0.03 0.02 ndd ndd ndd ndd ndd ndd ndd ndd ndd ndd ndd ndd ndd ndd ndd

Source: Iyer et al. (1999b). Note: na ¼ not available, nd ¼ not detected, ndd ¼ not determined; IONF ¼ Indian Ocean Nodule Field. In sediment cores S89 and S657, Ni and S contents range from 0.02 to 0.14% and 0.01 to 1.55%, respectively, in spherules, and 0.02 to 1.4% and 0.21 to 4.65%, respectively, in hydrothermal precipitates. #1–17 ¼ Volcanic magnetite spherules of S94; 18 ¼ Metallic grain of S94; 19 ¼ Average of 24 spherules from S89 and S657 (Iyer et al. 1997a); 20 ¼ Olivine inclusion of S94; 21 ¼ Plagioclase inclusion of S94; 22–26 ¼ Si-Fe-Al-enriched phase of S94.

Table 3.13 #

Composition (wt%) of magnetite spherules, hyalosidertic olivine and Ti-rich particles from the IONF

Si

Ti

Al

Fe

Mn

Mg

Ca

Na

K

Remarks

1 2 3

0.86 18.08 7.35

0.30 0.16 31.25

0.21 0.82 2.80

73.83 27.58 9.58

0.36 0.14 4.96

0.39 14.49 0.86

0.26 0.29 1.75

0.22 0.10 0.37

0.07 0.07 2.21

4 5

1.31 0.23

36.43 tr

2.33 0.10

18.59 72.05

5.28 na

0.24 tr

0.46 tr

na na

0.07 na

6

0.15

tr

0.09

71.91

na

tr

tr

na

na

7

0.07

tr

0.05

72.10

na

tr

tr

na

na

8

0.25

tr

0.16

71.01

0.05

tr

tr

na

na

9 10

na 8.20

na na

na na

71.50 42.90

0.37 0.25

na na

na na

na na

na na

11

23.67

0.65

7.33

10.04

na

3.90

6.85

1.97

0.27

12

23.50

1.27

7.44

10.10

na

4.20

7.54

1.98

0.30

13

21.86

0.70

7.41

7.15

0.15

5.40

9.39

na

0.36

Avg. of 24 IONF spherules Hylosidertic olivine from the IONF Avg. of 5 Ti-rich areas of a IONF spherule Avg. of 2 Ti-rich areas in the inclusion Magnetite spherule, 30–100 mm, Mt. Etna Magnetite spherule, 30–100 mm, Mt. Lipari Magnetite spherules, 30–100 mm, Mt. Vesuvius Magnetite spherules, 30–100 mm, Mt. Bracciano Volcanic spherules from the MAR Irregular Si-rich channels in MAR spherules Avg. of 57 basaltic spherules from DSDP site 32, NE Pacific Ocean Avg. of 14 microlapilli (300–2000 mm) from DSDP site 32, NE Pacific Ocean Avg. of basaltic microlapilli from two cores near Vityaz Fracture Zone, Indian Ocean

14

15.00

7.00

7.00

31.00

tr

na

3.00

na

1.00 (continued)

Table 3.13 (continued) #

Si

Ti

Al

Fe

Mn

Mg

Ca

Na

K

15

23.00

3.00

10.00

20.00

tr

na

6.00

na

1.00

16 17 18 19 20

4.71 5.02 6.70 17.08 8.53

na na 1.35 0.86 1.14

16.63 13.28 6.01 18.21 22.47

45.04 49.89 44.31 13.46 18.10

na na 1.22 0.74 1.03

na na 3.88 2.23 3.52

0.32 na 1.46 1.18 1.36

na na 2.28 1.47 2.37

0.26 na 1.59 1.65 1.54

Remarks

Avg. of 50 volcanic spheroids of avg. 101 mm from 5 Pacific coast volcanoes Avg. of 95 volcanic spheroids of avg. 54 mm from Irazu, Kilauea Iki, Ubinas and Huainaputina Large Fe-Al sphere Small Fe-Al sphere Blocky Fe particle Coating on Fe particle Fe-Al spindle

Source: Iyer et al. (1997b). Note: # ¼ 1–4 ¼ S89 and S657; #5–8 ¼ Del Monte et al. (1975); #9, 10 ¼ Aumento and Mitchell (1975); #11 ¼ Melson et al. (1988); #12 ¼ Vallier et al. (1977); #13 ¼ Nath and Iyer (1989); #14 ¼ Hodge and Wright (1964); #15 ¼ Wright and Hodge (1965); #16–20 ¼ experimental products (Heiken and Wohletz, 1985). avg ¼ average, na ¼ not available, tr ¼ trace; IONF ¼ Indian Ocean Nodule Field. In #1, 2 and 4 the Cr values are 0.08, 0.43 and 0.02% and Ba 0.08, 0.68 and 0.93%, respectively. Cr in #14 is up to 1% and in traces in #6 and 8.

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101

2.1.3. Glass shards Glass shards are common in the IONF sediments (Fig. 3.12). Abundant platy glass shards of 150–250-mm length occur within the ochrous sediments or as individual fragments. Their number gradually decreases down-core from a maximum of 7000 per slide at 1–2 cm to about 2500 per slide between 5- and 6-cm depth and decreasing further down-core. The shards are rhyolitic with SiO2 varying between 74 and 81.5% (average 77.5%; Table 3.14A). The source of the IONF shards is not yet convincingly deciphered. For instance, the supereruption of Toba volcano in northern Sumatra about 74 ka ago (Rose and Chesner, 1987) resulted in widespread fallout of tephra to more than 3000 km in a northwesterly direction (Dehn et al., 1991; Ninkovich et al., 1978). Glass shards (28–470-mm long) from a few sediment cores in the IONF compositionally correspond, in fact, to the Younger Toba Tuff (YTT) (Pattan et al., 1999), which is suggested to have been dispersed up to 14 S. However, the Ti/Al ratio seems to indicate that the IONF shards originated as distal fallout from the IVA (Martin-Barajas and Lallier-Verges, 1993). Sukumaran et al. (1999) suggest that the shards could not have been carried 3000 km from the eruptive source, either by surface or by northward-flowing deepsea currents, and thus rule out Toba as a source. Based on 230Thxs values and radiolarian biostratigraphy, the authors favour the deposition of the shards consequent to suboceanic volcanic activities during periods of elevated glaciation. The major Indonesian volcanic activities are dated to have erupted at 74,000, 450,000 and 840,000 years and at 1.2 Ma (Rose and Chesner, 1987). These dates do not correspond to the 10 ka age of the vhm with which IONF shards coexist. Again, significant dissimilarities appear to exist when certain elemental concentrations and their ratios of the IONF shards and for the Toba shards are compared (Tables 3.14B). Hence, it is yet unclear whether the shards are (1) formed contemporaneously with the metalliferous sediments, or (2) pre-existing entities in the IONF sediments, or

Figure 3.12 Scanning electron microscope (SEM) image of platy glass shards recovered with the magnetite spherules from core S657. Scale bar ¼ 10 mm.

Table 3.14

Elemental abundance (wt%) in volcanic glass shards from the IONF and Toba, Indonesia

A. Composition of glass shards Author

SiO2

TiO2

Al2O3

FeO

MnO

MgO

CaO

Na2O

K2 O

Cr2O3

BaO

Remarks

I

77.48

1.27

9.28

6.87

0.28

0.32

1.08

0.95

2.33

0.27

0.14

S P

77.10 76.81

0.05 0.07

12.60 12.77

0.90 0.92

0.07 0.06

0.05 0.05

0.76 0.79

3.25 3.41

5.04 5.08

na na

na na

Rhyolitic shards (6) YTT (81) IONF cores (91)

B. Representative elemental concentrations (wt%) and ratios

Si Ti Al Ti/Al Si/Al K

1

2

3

4

5

36.22 0.19 4.91 0.04 7.38 1.94

33.75 0.07 6.83 0.01 4.94 4.23

33.56 0.08 6.93 0.01 4.84 4.23

36.59 0.04 6.72 0.01 5.44 3.54

25.39 1.01 9.24 0.11 2.75 2.87

Source: Iyer et al. (1997a). Note: Number of shards analysed are in parentheses. YTT ¼ Youngest Toba Tuff, I and S ¼ Shane et al. (1995), P¼ Pattan et al. (1999). na ¼ not available; IONF ¼ Indian Ocean Nodule Field. Note: 1 ¼ Iyer et al. (1997a); 2, 3 ¼ Toba glass shards (Ninkovich et al., 1978); 4 ¼ Average of 88 Toba glass shards (Rose and Chesner, 1987); 5 ¼ IONF glass shards (MartinBarajas and Lallier-Verges, 1993).

Volcanics

103

(3) from earlier IVA eruptions, or (4) attrition products of the pumice. Geochemical studies by Mascarenhas-Pereira et al. (2006) have provided proof of intra-plate volcanism as a source for the glass shards.

3. Alteration of Volcanics Low-temperature alteration of deep-sea volcanics is a ubiquitous process that produces various authigenic minerals. Such alterations involve a precursor, one or more processes and the end products. The IONF basalts are most commonly altered to palagonite, clay minerals (montmorillonite and smectite), zeolites and iron oxides and hydroxides.

3.1. Alteration of basaltic glass The main processes in the alteration of the oceanic crust are related to the availability of weak zones, composition of seafloor rock, seawater weathering, metamorphism, hydrothermal circulation and deuteric compositional changes brought about by magmatic solidification. The effects of seawater–rock interaction are the result of two processes: (1) low-temperature weathering (or halmyrolysis) and (2) hightemperature alteration. The former occurs at 106 years). The latter process is hydrothermal, occurs in the temperature range of 70–400 C or higher, is generally restricted to sites near accreting plate margins or mid-plate volcanic areas and spans relatively for short intervals (102–104 years; Thompson, 1991). Altered basalts are a source of Si, Mg and Ca and a sink for Ti, K, P, Mn, total Fe and Na. Water and potassium are the most effective and noticeable parameters of alteration. It is generally agreed that the magnitude of changes in the trace elements is always greatest in the altered glass. For instance, Ba is enriched and Zn is quite constant during palagonitisation of basaltic glass. Basaltic lavas erupted on the seafloor are quenched to form a glassy skin that preserves the nearly pristine magma chemistry. One notable alteration product of this glass is palagonite, supposedly formed at high temperatures almost instantaneously during eruption (Bonatti, 1965). However, it is now accepted that the palagonitisation occurs gradually after the lava is emplaced (Moore, 1966). Three stages of palagonitisation of basaltic glass have been recognised (Honnorez, 1981): Initial—coexistence of fresh glass relics with residual altered glass (palagonite), increase in K and Mg and loss of Ca from altered rocks. Mature—alteration of fresh glass and in situ replacement of residual glass by zeolites and smectites. Final—complete replacement of residual glass by authigenic minerals. Dredged basalts recovered from various sites in the IONF (Fig. 3.3) were examined for the presence of glass and its state of alteration (Fig. 3.13). The samples are of various sizes and non-uniformly altered to yellowish-red palagonitic materials.

104


Ferromanganese oxides of variable thickness may overlie the basalts while the inner zone below is partly altered. Petrographic studies revealed the alteration of glass to palagonite (Fig. 3.14), plagioclase to sericite, olivine to iddingsite and appearance of limonite (Table 3.15). X-ray diffraction (XRD) analysis of a few samples shows phillipsite, the formation of which during the halmyrolysis of glass is well documented (e.g. Bonatti, 1965; Honnorez, 1981; Thompson, 1991). The major, minor and trace element concentrations in the altered basalts of the IONF samples (Table 3.16) are comparable to those from other oceans (Table 3.17). Removal or addition of certain elements during the alteration can be deciphered through binary plots. For example, with increasing K2O, CaO þ MgO decreases

Figure 3.13 Holohyaline and glass-flow textures of seamount basalt in the Indian Ocean Nodule Field (IONF). Note the alteration and devitrification of glass. Scale bar = 0.05 mm.

Figure 3.14 Photomicrograph of altered basalt with sub-microscopic sheaves of plagioclase. Small elliptical blebs occur around nucleus of plagioclase microlites. Scale bar = 0.05 mm.

105

Volcanics

Table 3.15

Petrography of altered basalt in the IONF

Location ( S/ E)

Depth (m)

Description

13 000 / 76 290 0 13 21 / 77 300

5346

Altered glass

5270

12 280 / 76 520

na

12 570 / 75 000

5348

12 330 / 77 120

5200

12 300 / 78 120

5280

12 130 / 76 110

5100

0.5 cm glass with rust coloured stains. 3 distinct zones: altered glass (A), glass þ interior (B) and interior (C). Zone B has sub-microscopically intergrown plagioclase and a few pyroxenes. Zone C is partly altered and has plagioclase microphenocrysts. Zeolites, limonite and smectite occur. 2 layers of glass. The upper unaltered glass has olivine and is separated from the lower layer by a band of greenish-yellow palagonite, between the bands of palagonite zeolites occur. Highly altered, 0.5 cm glass with rust coloured stains. Interior is buff to light gray, 2 glassy layers separated by palagonite. Uppermost glass hosts well-developed and fresh plagioclase. Lower glassy layer is oxidised and has some fresh plagioclase. The inner layer contains plagioclase and glass showing intersertal texture. RFA is also present. Glass of 0.5 cm with a veneer of sediment and Fe-Mn oxides. Plagioclase form trachytic texture. A few olivine grains present. Altered glass with patches of Fe-Mn oxides. A few altered grains of olivine, abundant vesicles and sheaves of plagioclase, RFA. Altered glass with patches of Fe-Mn oxides in a matrix of fibrous palagonite. Abundant plagioclase microlites as sheaves and in trachytic forms in glassy groundmass. Occurrence of a few fresh euhedral plagioclase phenocrysts, altered and cracked euhedral plagioclase, and olivine. Sub-ophitic and RFA.

Source: Iyer (1999a). Note: na ¼ not available; RFA ¼ Red feathery alteration texture; IONF, Indian Ocean Nodule Field (cf. Baragar et al., 1977).

reflecting the alteration of plagioclase and olivine and the formation of smectite. Decrease in CaO and MgO and increase in K2O, with increasing loss on ignition (LOI) (Fig. 3.15), could represent a nearly isomolar exchange during the alteration of basaltic glass to palagonite, an observation in accordance with Staudigel and Hart’s (1983) study. Although MgO shows higher concentration in less altered materials for a given LOI content, the variation in MgO is small. The relation also shows K2O to increase linearly with increasing alteration, except in those samples having higher MnO. CaO has a strong negative trend, indicating depletion of lime with increasing alteration (Iyer, 1999a).

106 Table 3.16


Composition of altered basalt in the IONF

Sample no.

81

6

14

24

26

85

87

SiO2 TiO2 Al2O3 Fe2O3 MgO CaO Na2O K2O P2 O 5 MnO LOI Ba Co Cu Cr Ni V Y Zn Zr

43.64 0.68 14.80 7.58 0.97 1.56 5.42 3.74 0.39 3.71 15.79 784 160 1203 32 1519 191 96 342 334

39.59 2.02 14.29 15.39 3.41 3.66 2.58 1.97 0.10 1.71 12.39 190 146 442 151 588 190 33 295 206

41.72 1.09 15.34 13.93 4.81 3.08 2.77 1.91 nd 0.30 12.14 83 90 302 519 217 220 23 270 128

41.18 1.63 17.84 13.35 1.84 1.34 2.59 2.82 nd 0.25 15.41 81 67 321 373 101 100 22 310 186

38.28 1.79 13.16 14.68 1.49 1.53 2.92 3.18 0.10 3.89 17.34 237 486 408 84 900 182 54 375 275

37.28 1.31 13.86 11.71 1.52 1.19 3.18 3.01 nd 6.75 17.98 604 459 1406 92 2470 154 60 420 240

37.10 1.14 14.08 12.44 1.58 1.22 3.60 2.94 nd 7.29 16.97 547 429 1290 83 2224 123 72 390 272

Source: Iyer (1999a). Note: nd ¼ not determined; LOI ¼ Loss on ignition; IONF ¼ Indian Ocean Nodule Field. Major elements in wt%, Trace elements in ppm.

Table 3.17 Composition (wt%) of palagonite, most altered basalt material and red feathery altered (RFA) materials from the world oceans Oxides

1

2

3

4

5

6

7

8

SiO2 TiO2 Al2O3 FeO* MgO CaO

42.05 2.13 11.43 19.31 3.95 1.48

53.42 2.38 15.60 20.58 5.15 1.11

44.19 2.81 15.42 15.72 2.82 0.60

41.80 1.90 12.15 19.96 4.06 0.72

37.88 1.21 15.38 11.97 7.56 5.44

41.96 1.63 16.06 9.39 3.69 1.49

47.11 1.55 18.37 8.91 1.84 2.34

47.65 0.42 19.45 9.80 3.44 11.03

Source: Iyer (1999a). Note: 1¼ DSDP leg 37 (Scarfe and Smith, 1977, average of 4 samples); 2 ¼ DSDP leg 37 (Andrews, 1977, average of 6 samples); 3 ¼ average of 3 samples (Melson, 1973); 4 ¼ average of 6 samples (Baragar et al., 1977); 5 ¼ average of 4 samples (Furnes, 1978); 6, 7 ¼ Swallow Bank, Atlantic Ocean (Matthews, 1971, 1 sample each); 8 ¼ RFA material from DSDP sites 332 and 335, MAR basalts (Baragar et al., 1977, average of 6 samples).

107

Volcanics

B

A

4

81 26

5

3

85

K2O

4 K2O

x 3

x

87

24 2

6 14

2

1

1 0 0

2

4

6

8

0

10 12 14 16

10

8

CaO + MgO C 4.0

14

12

16

18

LOI

81

H

3.5 26

85

K2O

3.0 24

87

2.5

L 6

2.0

14

1.5 2

4

6 8 10 CaO + MgO + MnO

12

D

100

81 H

2

87

6 26 85 87

TiO2

L 14

1

85

60

26

Y

24

H 6 24 L 14

81 20

50

150

250 Zr

350

50

150

250

350

Zr

Figure 3.15 Inter-element relations in the Indian Ocean Nodule Field (IONF) basalts. (A) K2O vs. CaO þ MgO for palagonite and altered basalts. Palagonite: solid dots ¼ Iyer (1999a); cross ¼ Scarfe and Smith (1977); circle with cross ¼ Andrews (1977); open circle ¼ Melson (1973); square and plus sign ¼ Barager et al. (1977); inverted triangle ¼ Matthews (1971). Altered basalts: circle with dot ¼ Melson (1973); triangle ¼ Barager et al. (1977); (B) loss on ignition (%)vs. K2O (%)for the altered basaltic glass of the IONF and other oceanic occurrences. Symbols as in (A); (C) K2O vs. CaO þ MgO þ MnO of the altered basalts. L and H represent samples with 2% MnO content, respectively; (D) Zr vs.TiO2, and Zr vs.Y for the altered basalts. See text for explanation. Zr and Y in ppm, rest in wt%.

108


Barium seems to relate more closely with MnO than with K2O and H2O, indicating its association with the biogenic debris or aluminosilicates. Similarly, Co, Cu, V, Ni and Zn contents are higher in MnO-enriched samples, a feature common in manganese nodules. Chromium correlates with MgO signifying Cr to generally occur in olivine or be supplied from the seawater. Depending on the MnO contents, inter-elemental plots made for Y, Zr and Ti (elements less sensitive to weathering and non-mobile) point to two clusters (Fig. 3.15). Samples with >2% MnO have high Ba, Co, Cu, Ni, Y, Zn, Zr, TiO2, Co and low Cr, (cluster H) while those with lower MnO content show a reverse trend (cluster L). Strong positive correlations between Y and Zr in both the clusters and between Zr and TiO2 in cluster L (Fig. 3.15) suggest that though both Zr and Ti are immobile, some amount of Zr, which probably admixed with the phosphatic phase (fish debris?), is removed during the leaching of P. The composition of palagonite, most altered basalt material and red feathery alteration (RFA) from the world oceans are furnished in Table 3.17. Based on the above pattern of chemical exchanges, the following process is envisaged for alteration, leading to palagonitisation of the IONF basalts. Basaltic eruptions that occurred about 60 and 45 Ma ago, chilled to form pillow basalts with a quenched glassy exterior and holocrystalline interior. Subsequent interaction with seawater caused a gradual breakdown of the homogeneous glass that shows a distinctly wider range of elemental change vis-a`-vis the polyphase crystalline interior. The mineralogical changes noted are the presence of palagonite, clays, iron oxides and phillipsite. Compositionally, Si, Mg and Ca were leached while Na and K were incorporated, with Fe and Ti remaining immobile. These characteristics typify an initial-to-intermediate stage of palagontisation of the IONF basalts under low-temperature oxidising conditions. Later, hydrogenous precipitation of ferromanganese oxides formed a protective veneer over the basalts and arrested the progress of alteration (Iyer, 1999a,b).

3.2. Zeolitites Marine authigenic phases such as phosphorites, manganese nodules, feldspar, metalrich sediments and zeolites are of considerable importance, as they play key roles in the geochemical cycling of different elements, both in seawater and in sediments. Zeolites may constitute up to 80% of oceanic and volcaniclastic sediments and their composition depends, among other factors, on (1) reaction time for devitrification, (2) crystallochemical transformations of glass of different Si/Al ratios at low temperature and pressure, (3) sedimentation rates, (4) sediment type and (5) Eh–pH conditions. A few studies that have been carried out concerning the Indian Ocean zeolites are by Arrhenius (1963) and Bonatti (1963) who reported natrolite occurrence, while Kastner and Stonecipher (1978) and Iijima (1978) detailed the presence of clinoptilolite (59%), phillipsite (22%) and analcime (17%). Kolla and Biscaye (1973) concluded that the alteration of basic and silicic volcanics in the Indian Ocean could have formed phillipsite and clinoptilolite, respectively. Besides these, phillipsite occurs in the IONF as well developed diagenetically formed crystals within manganese micronodules (Banerjee and Iyer, 1991) and in one case, as large crystal (21 10 8 mm) forming the nucleus of a manganese nodule (Ghosh and

109

Volcanics

Mukhopadhyay, 1995). Interestingly, there are indurated zeolitic slabs in the IONF, as in the Pacific Ocean (Morgenstein, 1967), which are described later. From the IONF, a few unusual porous and semi- to highly-indurated specimens were recovered, mostly from the siliceous sediment though, a few were from red clay domain (Table 3.18). The samples are between 3- and 20-cm long and may occur proximal to seamounts and fracture zones (Fig. 3.3). The samples are ‘fresh’, light pink to yellowish in colour and unaltered, though weathered ones are common. A few among these occur as the nucleus of nodules and at times are bioturbated (Fig. 3.16) (Banerjee, 2000; Iyer and Sudhakar, 1993b). Binocular microscopy (Fig. 3.17A) of the samples displays irregularly shaped pores, cavities, burrows and tubes that may be lined with ferromanganese oxides. Some of the samples show initiation of nodule growth-cusp around a zeolitic nucleus. Some others show a distinct entrapment of pumice in association with fine needles of zeolite. Electron microscopy (Fig. 3.17B) reveals platelets of clay minerals and phillipsite crystals. Although plagioclase, sometimes as an impurity, obscures the XRD peaks (Sheppard and Gude, 1983) the prominent XRD peaks (d ¼ 3.15 A˚) are similar to those for harmotome (Fig. 3.18A) and to the typical ‘Indian Ocean phillipsite’ (dA ¼ 3.19; location—23 160 S/75 590 E, Gottardi and Galli, 1985). This was confirmed by differential scanning calorimetry (DSC) profiles that Table 3.18

Geological setting of a few zeolitites from the IONF

Sample

Description (size in cm)

1

25 21.5, palagonite and Fe-Mn oxides 7.5 4.5, hard, compact

2 3 4 5 6 7 8 9 10 11

4.5 2.5, flat, ‘laterite-like’ with Fe-Mn oxides length 10 cm, highly altered, yellow 9.5 7.5, yellowish, flat a) 9 6, tabular, yellowish with Fe-Mn oxides b) 5.5 5, altered with Fe-Mn oxides(0.3) 12.5 7.5, ‘fresh’, red 7 4, flat, gritty with Fe-Mn. 5 2.5, flat, black 12 7.5, flat, black to buff, burrows

Position ( S/ E)

Depth (m)

Topography

13/75.7

4275

Seamount

12.6/ 76.1 11.2/73

5300

Seamount

5340

12.4/ 76.8 12.7/ 79.5 11.3/ 73.2 12.4/ 76.3 13/75.7 17/76.9 16/83 14/79.4

5050 5070

Fracture zone/ seamount Flank of seamount Fracture zone

5100

Seamount

4800

Seamount

4275 5100 4872 5172

Seamount Abyssal plain Abyssal plain Seamount

Source: Iyer and Sudhakar (1993a,b). Note: Samples 1–8 from siliceous sediment; 9 and 10 from red clay and 11 from siliceous-red clay, Fe-Mn oxides ¼ Ferromanganese oxides; IONF ¼ Indian Ocean Nodule Field.

110


A

B

0

1

2

cm

3

4

5

Figure 3.16 Specimen of bioturbated zeolitites recovered from the Indian Ocean Nodule Field (IONF). Note the presence of ferromanganese oxides on the lowerspecimens. Scale bar in ‘A’ = 2 cm.

showed exothermic water loss at 157 C and 238 C (Fig. 3.18B) similar to that of the phillipsite–harmotome family (cf. Gottardi and Galli, 1985). As the samples are formed mostly of zeolites and following Deffeyes’s (1959) definition, the term ‘zeolitites’ has been proposed (Iyer and Sudhakar, 1993b). The two possible precursors for the formation of IONF zeolitites are: (1) a continuous layer of volcanic ash and/or altered fragments of pumice, or (2) devitrification of basaltic glass. Ash layers in the IONF (Martin-Barajas and Lallier-Verges, 1993; Vallier and Kidd, 1977), which act as substrates for hardgrounds and pavements (Gupta, 1991), may have been disintegrated by mass movements, erosion by underwater currents, solution and stirring by organisms and tectonic reactivation (Iyer and Sudhakar, 1993b; Menard, 1960) and subsequently were wholly or partly indurated and altered to zeolites. Phillipsite and clinoptilolite, respectively, are the dominant or sole zeolite in silicic tuffs and in mafic ash layers. Interestingly, the IONF zeolitites are associated

111

Volcanics

with both mafic and silicic precursors. They occur in a large field of pumice (600,000 km2 in area) with the pumice fragments sometimes entrapped within the zeolitic slabs, an observation that supports siliceous precursors. Mafic precursors are evidenced by the presence of numerous altered basaltic fragments. Under favourable conditions, the mafic precursors are transformed into zeolites through an intermediate stage of palagonite formation (see earlier section, Iyer et al., 2007b). A

a

b

c

Figure 3.17 (Continued)

112


Figure 3.17 (A) Photomicrographs of polished sections of zeolitites:(a) dendrites and spicules of ferromanganese oxides in a matrix of dominantly zeolitic material; scale bar = 300 mm; (b) cusp formation and initiation of nodule growth with zeolitic material as nucleus; scale bar = 300 mm; (c) a fragment of pumice within the zeolitic material, observe the fine needle-like crystals of phillipsite at the centre; scale bar = 500 mm; (B) Scanning electron microscope (SEM) images of polished zeolititesç(a) platelets of clay minerals; scale bar = 2 mm; (b) well-formed stubby crystals of haromotome (cf. Sheppard and Gude, 1983); scale bar = 4 mm; (c) well-developed prismatic crystals of phillipsite; scale bar = 6 mm; and (d) crystals of zeolites of varied nature; scale bar = 4 mm.

Most of the IONF zeolitites occur near the major fracture zones (Vishnu FZ along 73 E, 75 450 E FZ and Indrani FZ along 79 E) and seamounts. Such locales possibly are conducive for zeolite formation by the alteration of the glass, palagonite and volcanic debris by dilute, low-temperature hydrothermal fluids. The IONF samples are not only typical examples for a process of complete zeolitisation but also the presence of the ferromanganese oxides accounts for later localised induration.

4. Tectonics and Volcanics: Interrelations It is now seen that the majority of the IONF seamounts were formed at the ridge axis between 60 and 45 Ma and were transported to the present sites along with the underlying crust. Many of these seamounts have an east-west ‘bulged’ elongation, which can be attributed to later addition of magma along weak zones after the seamounts moved to the present sites from the MOR (also see Chapter 2) and this is suggestive of multi-episodic volcanism in the basin. The weak zones may

113

Volcanics

A

(041)

PI (140)

(151) (103) (320) (323)

36⬚

(001)

(131)

(321)

(021) (120)

PI (121)

30⬚

20⬚

10⬚

2q B

−3 −4

98.31⬚C

−5 Heat flow (mW)

I −6

II 211.64⬚C

−7 −8 −9 157.10⬚C

238.04⬚C

−10 20

70

120

170

220 270 320 Temperature (⬚C)

370

420

470

Figure 3.18 (A) X-ray profile of zeolitites. PI ¼ Plagioclase impurity; (B) differential scanning calorimetry (DSC) profiles of zeolitites. See text for explanation.

have resulted during mid-plate tectonic reactivation (owing to tensional or compressional stress) of diverse intensity. Volcanic-hydrothermal materials (vhm) as well as altered basalts (zeolitites and palagonites) generally occur near to these ‘bulged’ seamounts and weak zones. These morphotectonic features probably

114


facilitated the circulation of low-temperature fluids, and brought about hydrothermal alteration. High concentrations of vhm in the IONF (metalliferous sediments, volcanic spherules and glass shards), associated with sediments of 10 ka and 425–650 ka age, are likely to have formed in situ from volcanic emanations or hydrothermal solutions. It is possible that the laboratory experiments (interacting thermite melt with watersaturated quartzo-feldspathic sands) have been mimicked by local volcanic eruptions in the IONF. Plausibly, the interaction between the Fe-rich lavas or hydrothermal emanations and the quartzo-feldspathic sands (equivalent to abundant siliceous sediments in the IONF) may have resulted in localised hydrovolcanic events and production of magnetite spherules. However, considering the hydrostatic pressure at a depth in excess of 5000 m, which is well above the critical pressure of generation of a vapour phase, a more detailed investigation is required to explain hydrovolcanic events at abyssal depth. Hence, the presence of a variety of major and minor volcanics in the IONF may imply that this field is not tectonically and volcanically inactive (Iyer and Sudhakar, 1995). This fact may have a strong bearing on the formation, distribution and enrichment of ferromanganese deposits in the IONF (see Chapter 5).

C H A P T E R

F O U R

Sediments

Contents 117 117 124 135 137 139 146 149 151

1. Distribution and Source 1.1. Distribution 1.2. Source 2. Sedimentary Processes 2.1. Dissolution of carbonate 2.2. Bottom water mass and sedimentation 2.3. Depositional environment 2.4. Diagenesis 2.5. Sediment consolidation

Sediments are formed from disintegrated rocks as a result of physical and chemical weathering. The action of various agents, like ice, wind, water and variable temperature, helps fragment the rocks into smaller particles and leach the more soluble minerals. Weathering also results from reaction of seawater with basalt, erupting at the crests of the mid-oceanic ridges and at other submarine volcanic features, contributing in the process considerable amounts of materials to seawater. Another significant contributor has been extraterrestrial matter. Transported either in a dissolved or suspended state, all these materials ultimately form four types of sediment—biogenous and hydrogenous through precipitation, and lithogenous and cosmogenous in a clastic detrital state. Among the biogenous deposits are calcite, aragonite, opal, phosphorites and organics. The hydrogenous types are evaporites, zeolites, manganese nodules and polymetallic sulphides, while silica, feldspar and rock fragments constitute deposits of lithogenous type. The contributors to cosmogenous deposits have been the cosmic spherules, microtektites and minitektites (Chester, 1993; Lisitzin, 1996). In general, mechanical weathering dominates in temperate climates at high latitudes, where water in the form of ice is the chief weathering agent. Chemical weathering (i.e. leaching), on the other hand, is favoured by high rainfall, variable temperature and dominates in tropical areas. Hence, any sedimentary basin reflects the palaeo- and present environmental conditions of sedimentation including the processes and provenances. The distribution of various sediment types in the world’s oceans has been recorded with reliable accuracy (Fig. 4.1; Table 4.1). There have been considerable activities in recent years to understand the sediment distribution and sedimentary processes in the Central Indian Ocean Basin (CIOB), a basin which Handbook of Exploration and Environmental Geochemistry, Volume 10 ISSN 1874-2734, DOI: 10.1016/S1874-2734(07)10004-8

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160W

100

40W

0

40E

100E

60 N

30

0

30

60 S

Ice rafted

Carbonate

Siliceous

Red clay

Terrigenous

Siliceous/red clay

Figure 4.1 Sediment distribution in the world oceans (Kolla and Kidd, 1982; The Open University,1995).

Table 4.1

Distribution (%) of various sediment types in the world oceans

Major type

Minor type

Red clay Siliceous ooze

Diatom ooze Radiolarian ooze Calcareous ooze Foraminiferal ooze Pteropod ooze

Atlantic

Indian

Pacific

All oceans

26 07 – 65 02

25 20 0.5 54 0.1

49 10 05 36 –

38 12 03 47 0.5

Sources: Berzukov (1960) and Stowe (1996).

hosts the sediments of the two largest rivers of the world, the Ganges and the Bramhaputra, and also holds the second richest and second largest manganese nodule field in the world oceans. Hence, the sediments described throughout this chapter, if not specifically mentioned, are essentially those of the CIOB, which also includes the Indian Ocean Nodule Field (IONF).

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Sediments

1. Distribution and Source 1.1. Distribution The nature and distribution of seafloor sediments in the Indian Ocean are principally controlled by five interrelated factors: (1) climatic and current pattern, (2) nutrient and organic production in surface waters, (3) relative solubility of calcite and silica, (4) submarine topography and (5) detrital input. Based on their interactions, four major types of sediments occur—terrigenous, calcareous, siliceous and pelagic (Fig. 4.2). Covering about 35% of the CIOB, calcareous sediments, with a sedimentation rate of 4–6 mm/103 year, are common along the equatorial highproductivity areas and at lesser depths, such as near the seismic and aseismic ridges (the Chagos Ridge, the Ninetyeast Ridge) and the shallow areas of seamount

60E

70

80

90 10 N

0

10 S

20

30

Calcareous Ooze

Siliceous Clay

Calcareous Clay

Brown Clay

Terrigenous Clay

Terri-Siliceous Clay

Mixture of Terrigenous, Calcareous, Siliceous Clay Southern limit of Indonesian Volcanic Tephra

Figure 4.2

Sediment distribution in the Indian Ocean (Udintsev,1975; Kolla and Kidd,1982).

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summits. Siliceous clay (and ooze) and red-brown clay, on the other hand, are the dominant sediment types in the deeper parts of the basins in the CIOB, and cover about 35 and 16% of the surface area, respectively. These sediments are found in areas where the rate of sedimentation is low (

The Indian Ocean nodule field: geology and resource potential

The Indian Ocean Tsunami

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Superpower Rivalry in the Indian Ocean: Indian and American Perspectives

The Indian Ocean (Seas in History)

The Indian Ocean Tsunami of 2004

The Indian Ocean (Seas in History)

Stars and Keys: Folktales and Creolization in the Indian Ocean

Monsoon: The Indian Ocean and the Future of American Power

Monsoon: The Indian Ocean and the Future of American Power

Monsoon: The Indian Ocean and the Future of American Power

Monsoon: The Indian Ocean and the Future of American Power

Coral Reefs of the Indian Ocean: Their Ecology and Conservation

Coral Reefs of the Indian Ocean: Their Ecology and Conservation

Atmosphere-Ocean Interaction (Oxford Monographs on Geology and Geophysics)

Ocean Optics (Oxford Monographs on Geology and Geophysics)

Antarctica As an Exploration Frontier - Hydrocarbon Potential, Geology, and Hazards (AAPG Studies in Geology 31)

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Potential

Classic Ships of Islam: From Mesopotamia to the Indian Ocean

Indian Ocean Studies: Cultural, Social, and Political Perspectives

The Indian Ocean Tsunami of 2004 (Great Historic Disasters)

Potential

Potential

The Indian Ocean in World History (Themes in World History)

Geology and Ecosystems

The Ocean

Ocean

Geology and Ecosystems

The Indian Ocean nodule field: geology and resource potential

The Indian Ocean Tsunami

Basics of Field Geology

Effective Crystal Field Potential

Superpower Rivalry in the Indian Ocean: Indian and American Perspectives

The Indian Ocean (Seas in History)

The Indian Ocean Tsunami of 2004

The Indian Ocean (Seas in History)

Stars and Keys: Folktales and Creolization in the Indian Ocean

Monsoon: The Indian Ocean and the Future of American Power

Monsoon: The Indian Ocean and the Future of American Power

Monsoon: The Indian Ocean and the Future of American Power

Monsoon: The Indian Ocean and the Future of American Power

Coral Reefs of the Indian Ocean: Their Ecology and Conservation

Coral Reefs of the Indian Ocean: Their Ecology and Conservation

Atmosphere-Ocean Interaction (Oxford Monographs on Geology and Geophysics)

Ocean Optics (Oxford Monographs on Geology and Geophysics)

Antarctica As an Exploration Frontier - Hydrocarbon Potential, Geology, and Hazards (AAPG Studies in Geology 31)

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Classic Ships of Islam: From Mesopotamia to the Indian Ocean

Indian Ocean Studies: Cultural, Social, and Political Perspectives

The Indian Ocean Tsunami of 2004 (Great Historic Disasters)

Potential

Potential

The Indian Ocean in World History (Themes in World History)

Geology and Ecosystems

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